Thermodynamic Models of Submarine Hydrothermal Systems
D. V. Grichuk
Faculty of Geology, Moscow State University, Vorob’evy gory, Moscow, 119992 Russia e-mail: [email protected]
Received March 3, 2004
Abstract
—This contribution presents a method for constructing thermodynamic models of convective hydro-thermal systems with an exogenous source of solutions. With the example of a simulation of modern hydrother-mal systems on the ocean floor, the prospects of such a method for the interpretation and forecasting of hydro-thermal mineralization are discussed. A procedure for constructing isotopic chemical models is described usingsulfur isotopes as an example. The role of boiling within oceanic hydrothermal systems in the formation of met-alliferous deposits is evaluated. This paper is intended for geologists who specialize in hydrothermal ore depos-its, geochemists, and marine geologists and can be used by students as a guide for the simulation of ore-formingprocesses.
Preface
This contribution by D.V. Grichuk addresses the development of methods of the thermodynamic modeling ofhydrothermal processes, and application of these methods to the investigation of very intriguing natural objects,ocean-floor hydrothermal systems.
Theoretical modeling is now the main tool of physical geochemistry. The past 50 years have seen considerableadvances in the development of general principles and exploration of some particular reactions to the quantitative inves-tigation of complex multicomponent systems closely correlated with natural prototypes. The Geochemistry Departmentof Moscow State University actively participates in the development of thermodynamic modeling. Studies by Grichukand his colleagues, M.V. Borisov, Yu.V. Shvarov, and others, have greatly contributed to the theory and application of thismethod. Grichuk’s paper is the result of many years of investigation aimed at the creation of a model for hydrothermalsystems with an endogenous source of energy and ore matter and an exogenous source of solutions.
The subject of his study, ocean-floor hydrothermal systems, shows a number of unique features. The most inter-esting among them for geologists is that it provides an opportunity to directly observe the formation of economic-scale ore mineralization. Although the economic value of ore occurrences in mid-ocean ridges is currently esti-mated rather pessimistically, the investigation of oceanic hydrothermal systems sheds light on the formation ofancient massive sulfide deposits. Hydrothermal vents on the ocean floor are very complex objects. Notwithstandingover 25 years of investigative research, there are still some unresolved problems, especially concerning the deepparts of hydrothermal systems. The development of theoretical models allows one to collect diverse pieces of infor-mation to obtain a comprehensive picture, which is clearly demonstrated by the study presented in this issue.
A characteristic feature of Grichuk’s study, which characterizes the current stage of the development of physicalgeochemistry, is that a theoretical model is used as a tool for the solution of geological problems. The conclusionsdrawn in this study regarding the evolution of ore composition owing to metasomatic processes in the interiors ofthe system, the mechanism of formation of zoning in the ore body, and the influence of boiling on ore mineraliza-tion provide new insights into the existing data and considerable progress in the theory of hydrothermal processes.
Head of the Geochemistry Departmentof Moscow State University
academician Vilen A. Zharikov
S160
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
GRICHUK
2.1.1. Infiltration metasomatism and its propertiesin the Korzhinskii theory
......................................S166
2.1.2. Method of the degree of reaction progress by Helgeson
2.1.5. Direct use of consequences of the Korzhinskii theory in algorithms of the simulation of infiltration metasomatism
....S173
2.2. Modeling of Metasomatic Processes in Extended Hydrothermal Systems
.....................S174
2.3. Modeling of Ore Deposition during Cooling of Hydrothermal Solutions
.....................S178
2.3.1. Scenario of closed-system cooling
.............S178
2.3.2. Fractional precipitation in a flow-through system
......................................S180
2.3.3. Scenarios of cooling with mixing
...............S181
2.4. Conclusions
...................................................S182CHAPTER 3. THERMODYNAMIC MODELOF A CONVECTIVE HYDROTHERMAL SYSTEM IN A MID-OCEAN RIDGE................S184
3.1. Geologic Setting and Composition of High-Temperature Oceanic Hydrothermal Systems
.........................................S184
3.1.1. Geologic setting
.........................................S184
3.1.2. Compositions and properties of sulfide edifices and hydrothermal solutions
.....S187
3.2. Logical Scheme of the Process and Its Experimental Substantiation
....................S209
3.3. Modeling of Modern Ore Formationin the Ocean: Previous Work
...............................S220
3.4. Model Description
........................................S224
3.4.1. Geological model
.......................................S224
3.4.2. Physicochemical model
..............................S226
3.4.3. Software for thermodynamic modeling
......S229CHAPTER 4. SIMULATION OF GEOCHEMICALPROCESSES IN THE HYDROTHERMAL SYSTEM OF A MID-OCEAN RIDGE................................S231
4.1. Results of Simulation for the DownwellingLimb of a Convection System
...............................S231
4.1.1. Model of a short-lived hydrothermal system
............................................S231
4.1.2. Chemical evolution of the system during the development of the hydrothermal process (model of a long-livedhydrothermal system)
6.4. Thermodynamic Model of Boilingin Oceanic Hydrothermal Systems
.......................S287
6.4.1. Isothermal boiling of solutionin contact with a rock: a modelfor the focus of a hydrothermal system
................S287
6.4.2. Phase separation in contactwith a rock (open system)
.....................................S291
6.4.3. Adiabatic cooling of two-phase fluid
..........S292
6.4.4. Condensation of two-phase fluidduring rapid discharge and ore deposition
..........S296
6.4.5. Verification of the model ...........................
S298
6.5. Conclusions
...................................................S298CHAPTER 7. POSSIBLE ROLE OF MAGMATIC FLUID: A COMBINEDEXHALATION–RECYCLING MODEL ............S298
7.1. Problem Formulation
....................................S298
7.2. Characteristics of the ExhalationConvective Model
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S161
INTRODUCTION
This paper summarizes the investigations that havebeen carried out by the author in 1979–2003 at theGeochemistry Department of Moscow State University.When this work was initiated, the thermodynamic mod-eling of geochemical processes had already developedinto a branch of the geosciences. Efforts of outstandingresearchers, R.M. Garrels, H.C. Helgeson, I.K. Karpov,I.L. Khodakovskii, B.N. Ryzhenko, and others, pro-vided a theoretical basis for the method, algorithms,and program codes for the calculation of equilibria inmulticomponent systems. The application of thermody-namic modeling to the investigation of geologic objectsrequired the development of simulation techniques ade-quately reflecting the complexity of geologic pro-cesses.
The goal of this work was to develop methods forthe thermodynamic modeling of convective hydrother-mal systems, devise a model for a hydrothermal ore-forming system in the oceanic crust, and characterize,with the aid of this model, hydrothermal ore-formingprocesses in a submarine environment.
I am very grateful to academician V.A. Zharikov,head of the Geochemistry Department, for his help andsupport during the entire period of this study. Of partic-ular significance over the years was the creative coop-eration of my colleagues B.N. Ryzhenko, M.V. Borisov,Yu.V. Shvarov, S.G. Krasnov, and M.Yu. Korotaev. Theauthor acknowledges the assistance, valuable discus-sions, and advices by Yu.V. Alekhin, L.A. Bannikova,E.N. Baranov, L.V. Dmitriev, A.V. Zotov, A.I. Krivtsov,A.Yu. Lein, V.I. Mal’kovskii, A.A. Pek, V.S. Savenko,N.M. Sushchevskaya, and A.A. Yaroshevskii. Assis-tance by A.Yu. Bychkov, A.V. Tutubalin, E.E. Abram-ova, and G.L. Mel’nikova is greatly appreciated.
Various stages of this study were financially sup-ported by grants of the Special Federal Program
WorldOcean
, the program
Universities of Russia
(Geomodelproject UR 09.003.003), the International ScientificFoundation, the Russian Foundation for Basic Research(project nos. 94-05-17301a, 96-05-64887, 99-05-64868, and 02-05-64282), the Federal Program for theSupport of Leading Scientific Schools (project nos. 96-15-98338 and NSh-491.2003.5), and contracts of theMinistry for Natural Resources of the Russian Federa-tion.
CHAPTER 1. GENERAL FORMULATION OF THE PROBLEM
1.1. Methodology of Thermodynamic Modeling
Recent advances in geology, as well as in other sci-ences, have arisen through the introduction of a newmethodological approach, which can be termed themodel approach. For a long time, geology developed byway of empirical generalizations. The expansion of ourknowledge of nature and, more importantly, the intro-duction of methods of exact sciences (including physi-
cal chemistry) disclosed the limited possibilities of theempirical approach. This motivated researchers in geol-ogy, in general, and geochemistry, in particular, to pur-sue studies of another type using theoretical models.
Problems related to the utilization of theoreticalmodels have been discussed in a number of publica-tions on the philosophy of science. Without going intodetail, the essence of the model approach can be brieflypresented as follows: this method replaces a studyobject by its model, which is simpler and more accessi-ble for investigation and corresponds to the object insome principal aspects; the results of the investigationof the model are transferred to the properties of theobject.
Two important points must be noted. First, an objectand its model cannot be exactly isomorphic; otherwise,model substitution for the object would not provide anyadvantages. The choice of key aspects depends on thepurposes of an investigation and is based on subjectivejudgments, and it is not known
a priori
if they are suf-ficient for goal attainment. Therefore, a necessary stepis to check their correspondence to the object, i.e., toverify the model. Second, the main stage of research isnot the construction of a model but the investigation ofits properties. The model is a tool rather than the goalof the work. When the model approach was being intro-duced into geology, this necessary property was some-what blurred, but now such a perception of models is ofspecial significance. A model must have predictivecapabilities, and some of its consequences must corre-spond to the as yet unknown properties of its naturalprototype.
The general scheme of a study utilizing the modelapproach is depicted in Fig. 1. The results of an inves-tigation of the properties of a natural object or objects(prototype) are used to perform a schematization, i.e.,the properties of the prototype essential for furtherinvestigations are distinguished, and a logical scheme isconstructed from them. A theoretical model is createdon the basis of the logical scheme using the tenets ofexact sciences. The theoretical model is explored, andits properties (model consequences) are determined.Some of the consequences are used to assess the effi-ciency of the model (verification), and others can beemployed to predict unknown properties of the naturalprototype.
It is obvious from this scheme that at least some ofthe model consequences must be suitable for a compar-ison with the object during model verification. How-ever, if all the consequences of a model are used for ver-ification, the predictive capacity of the model vanishes,and the model is useless. Note that the process of veri-fication is two-sided; the results of simulation may indi-cate a need for a refinement of some properties of thenatural prototype. It is even possible that some essentialparameters of the natural object will be unraveled only
S162
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
GRICHUK
during the construction of its theoretical model (e.g.,Bychkov, 1995).
Based on the above considerations, the problemsaddressed in model investigations can be divided intotwo groups: type I problems are external with respect tothe method of modeling, and type II problems are inter-nal issues of the model.
Type I comprises the following problems:—prediction of the unknown properties of natural
objects and—establishing causal and correlation relationships
between the known properties of objects.Type II comprises the following problems:—choosing the most adequate logical scheme for
the object from several alternative hypotheses,—demonstration of the plausibility of the accepted
logical scheme or revealing contradictions in it, and—demonstration of the internal consistency and
efficiency of the theoretical model.The problems are listed in order of descending sci-
entific importance. However, in terms of methodology,in a particular study, problems should be consideredfrom simple to complex, or in ascending order in theabove lists.
Theoretical models in modern geochemical investi-gations are primarily related to the application of quan-titative physicochemical methods, and the progress inthe development of the model approach during the pastdecades was related to the utilization of computer tech-niques. Therefore, three “layers” can be distinguishedwithin the majority of thermodynamic models ofgeochemical processes (
Methods of…
, 1988) (Fig. 1):—a geological model (logical scheme),—a physicochemical model and—a mathematical model (a means of obtaining con-
sequences).
The geological model defines the spatial and tempo-ral scales and
P
–
T
conditions of the process; mattersources and their mineral and chemical compositions;the mechanisms and characteristics of mass transfer;and the chemistry, mineralogy, and spatial distributionof the products of the process.
The physicochemical model describes the chemicalcomposition of the geological model in terms of a phys-icochemical system. If the methods of equilibrium ther-modynamics are utilized, the physicochemical modelincludes the thermodynamic properties of newlyformed compounds, equations for the calculation ofthermodynamic equilibria, and expressions describingreaction kinetics and the dynamics of mass transfer.
The mathematical model provides a tool for thequantitative solution of physicochemical model equa-tions (computation algorithm) and a computer code forits implementation.
Although such a three-layer division is obvious andeven trivial, each of the constituents of the thermody-namic model uses specific laws and methods and intro-duces its own set of simplifications and approximationsand corresponding sources of errors. It should be notedalso that within-layer approximations (for instance, theuse of the Debye–Hückel equation for the calculationof activity coefficients) are often universally acceptedand even unified, whereas across-layer transitions areusually subjective and specific in each investigation.
The application of equilibrium thermodynamicmethods, which deal with the states of a system, for theconstruction of models for natural processes results inan intrinsic contradiction. This contradiction was firstresolved by Helgeson (1968). This paper is still one ofthe most cited geochemical publications. Using the par-tial equilibrium concept of P. Barton, Helgeson pro-posed consideration of a nonequilibrium process as asequence of equilibrium states of the system, whosecomposition changes depending on the progress of anirreversible reaction. The Helgeson approach (methodof the degree of reaction progress) was extended and
Mathematicalmodel
Physicochemicalmodel
Geologicalmodel
Predictions
Dataon natural
object
Verification
Examination of
Schematization
model properties
Fig. 1.
Methodology of a simulation study.
Consequences
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S163
developed by other authors (e.g., Karpov, 1981). It isinherently applicable to the development of a process intime and ignores the movement of materials in space(this issue is discussed in more detail in Chapter 2). Thespatial variability of a process can be described usingthe principle of local equilibrium proposed by Korzhin-skii (1951, 1969). In this approach, the variability of aprocess in space and time is approximated by a series ofequilibrium states of the systems whose compositionsare interconnected by both kinetic and dynamic rela-tionships, which are the conditions of mass transferbetween domains of the geological model. The modelsof this class are referred to as equilibrium dynamicmodels (
Methods of…
, 1988). Such models are used inthis study.
In the past three decades, the thermodynamic mod-eling of geochemical processes has been extensivelydeveloped in Russia and other countries. Considerableprogress in this branch of geochemistry was related tothe efforts of Helgeson and his colleagues (Helgeson,1968; Helgeson
, 1978), B.N. Ryzhenko, Vikt.L. Barsukov,M.V. Borisov, R.P. Rafal’skii, B. Fritz (1975, 1981),and many other researchers. The success of their stud-ies was connected with the creation of efficient com-puter programs, including SELECTOR (Karpov
et al.
,Siberian Institute of Geochemistry), GIBBS(Yu.V. Shvarov, Moscow State University), EQ3/6(T.J. Wolery, Lawrence Livermore National Labora-tory), and CHILLER (M. Reed, US Geological Sur-vey). The investigations relied on the development ofbanks and bases of thermodynamic data embracing aconsiderable portion of geologically important sub-stances and a wide range of conditions. The mostwidely used databases are SUPCRT92 (Johnson
et al.
,1992), DIANIK (Vernadsky Institute of Geochemistryand Analytical Chemistry), THERMINEOS (SiberianInstitute of Geochemistry), UNITHERM (MoscowState University), and some others.
Thermodynamic modeling is used in moderngeochemistry for the description of diverse processes,from weathering to magmatism and protoplanetary dif-ferentiation. This study focuses on the methods of ther-modynamic modeling of hydrothermal processes by theexample of modern mid-ocean ridge systems.
1.2. Correlation of Ancient and ModernHydrothermal Processes
The first ore-forming seafloor hydrothermal systemwas discovered in 1963 in the axial part of the Red Seaby an American expedition of the R/V
Atlantis-II
(
HotBrines…
, 1969). This system (and similar adjacent sys-tems, which were later discovered) showed some verydistinctive features owing to the influence of Mioceneevaporates intersected by the Red Sea rift, and the ques-tion of its ancient analogues remains to be solved. The
discovery of this system and a series of subsequentstudies developing the hypothesis of the hydrothermal–sedimentary origin of massive sulfide ores as a result ofseawater circulation in the hot crust (Ohmoto and Rye,1974; Sakai and Matsubaya, 1974; Lister, 1972;Spooner
et al.
, 1977; Solomon and Walshe, 1979) havegenerated considerable interest in this problem.
Evidence for hydrothermal activity was found in theMid-Atlantic Ridge (MAR), the East Pacific Rise(EPR), and the Galapagos Spreading Center (GSC).The first high-temperature hydrothermal system wasfound at 21
°
N on the EPR at the end of 1979 (
RISEProject…
, 1980). During subsequent American,French, Russian, and international expeditions, newhydrothermal systems have been found on the oceanfloor almost each year. Up to now, such occurrenceshave been reported from many localities in the EPR,MAR, the system of the Juan de Fuca Ridge, and back-arc basins of the southwestern Pacific Ocean. Althoughit soon became clear that economic ore deposits couldhardly be discovered on the ocean floor, the impact ofthe data obtained on the theory of hydrothermal pro-cesses, marine geology and geochemistry, and evenbiochemistry and ecology (related to the discovery ofpeculiar populations of organisms near hydrothermalvents) was so significant that the number of publica-tions relevant to this problem has long exceeded 1000.The geology and geochemistry of oceanic hydrother-mal systems were addressed in a number of books; the-matic issues of the
Marine Mining
, the
Canadian Min-eralogist,
the
Economic Geology
, and the
Journal ofGeophysical Research
; and several special symposia.Oceanic hydrothermal systems are among the mostcomprehensively studied geologic objects.
These studies demonstrated a close resemblance ofmodern sulfide occurrences on the ocean floor toancient massive sulfide ores. The factual informationcollected during this work supported the possibility offormation of hydrothermal–sedimentary ores by thethermal convection of surface (marine) waters in crustalrocks near a magma chamber, which serves as a heatsource (recycling model, Krivtsov, 1981). Many char-acteristics of oceanic sulfide ores were found in ancientdeposits (Zaikov, 1991; Maslennikov, 1999). Some dif-ferences were revealed simultaneously, primarily ingeodynamic settings: almost all seafloor sulfide occur-rences are related to spreading zones in mid-oceanridges or back-arc basins, whereas an island-arc envi-ronment is reconstructed for ancient copper massivesulfide deposits (
Hydrothermal Sulfide…
, 1992).
Thus, the reason for the compositional similarity ofthe ores is related to the similar physicochemical mech-anisms and conditions of ore formation, which areprobably universal. Such a convergence of ore forma-tion mechanisms provides an opportunity to use thedata on oceanic hydrothermal systems for the construc-tion of a model of ore-generating processes, whose
S164
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
GRICHUK
application will be much wider than spreading zoneenvironments.
The following problems were considered in thisstudy
—development of methods for the simulation ofinfiltration metasomatic processes and elaboration ofnecessary thermodynamic and program tools;
—construction and exploration of a thermodynamicmodel of a hydrothermal system in the oceanic crustand verification of this model;
—development of methods for the simulation ofhydrothermal–sedimentary ore formation and construc-tion on this basis of a model for the formation of a mas-sive sulfide ore body;
—development of the thermodynamic modeling ofisotopic chemical systems and application of thismethod for the investigation of the sulfur isotopicmodel of ore formation in the ocean; and
—development of simulation procedures, construc-tion of a model, and assessment of the influence of boil-ing within a hydrothermal system on the processes ofore formation.
CHAPTER 2. METHODS OF THERMODYNAMIC MODELING OF HYDROTHERMALAND METASOMATIC PROCESSES
Problem formulation. Water–rock interaction is ofprime importance for the geologic processes involvingaqueous solutions, both hydrothermal and low-temper-ature. This phenomenon is responsible for the forma-tion of the chemical composition of metalliferoushydrothermal solutions and hydrothermal and metaso-matic ore deposition. Therefore, the development ofmethods for the simulation of water–rock interaction iscrucial for the construction of thermodynamic modelsfor hydrothermal processes.
The methods of modeling are currently developed intwo directions. One of them is aimed at simulation pro-cedures involving the dynamics of mass transfer andkinetics of chemical reactions (macrokinetic approachof Zaraiskii et al., 1989). Considerable progress hasbeen achieved in the development of the theoreticalbackground of this approach (Lichtner and Balashov,1993; Steefel and Lasaga, 1994). Unfortunately, as wasshown by Rafal’skii (1993), the paucity of data onkinetic constants hampers the wide use of macrokineticmodels. The application of this approach to modelswith sufficiently large numbers of components, allow-ing sophisticated geological interpretation, is a researchtopic of the distant future.
Another direction concerns the construction ofmodels combining the dynamics of mass transfer andimitation of the kinetics of chemical reactions througha series of equilibrium states of the system (equilibriumdynamic approach according to Methods of…, 1988).This approach was employed in the majority of modernstudies on the simulation of hydrothermal processes. In
accordance with the principles of local equilibrium(Korzhinskii) or partial equilibrium (Barton), the equilib-rium dynamic approach implies that the development of aprocess in space and/or time is considered as a series ofequilibrium states of chemical systems, whose composi-tions are defined by dynamic and kinetic relationships.
The calculation of equilibrium dynamic modelsincludes two phases: (a) prescribing a series of states ofthe systems (dynamic phase) and (b) calculation of theequilibrium states of these systems (equilibriumphase). There are sophisticated methods for the calcu-lation of equilibria in multicomponent multisystems(Karpov, 1981; Reed, 1982; Methods of…, 1988).There are also highly efficient computer programs(SELECTOR, HCh, EQ3/6, CHILLER, etc.) allowingequilibrium calculations for systems of any complexity.The input thermodynamic information provided bymodern thermodynamic data banks (SUPCRT92,UNITHERM, etc.) is sufficient to obtain important geo-logical results.
The main problem with this method is currentlyrefinement of the dynamic parts of the models. Sincethe dynamic characteristics of a model depend on theproperties of the geologic process considered, it ishardly possible to create a universal method appropri-ate for all processes. Studies conducted at theGeochemistry Department of Moscow State Universitydemonstrated that different methods (scenarios) have tobe used for the simulation of processes occurring invarious parts of hydrothermal ore-forming systems. Inthis chapter we evaluate methods for the simulation ofseveral geologic situations most common in currentmodel studies: (a) formation of infiltration metasomaticcolumns, (b) formation of ore-bearing solutions inextended hydrothermal systems with prevailing frac-ture percolation, and (c) ore deposition during coolingof hydrothermal solutions.
Some characteristic features of the methods of equi-librium dynamic simulation of hydrothermal and meta-somatic processes are illustrated by the example of sea-water interaction with tholeiitic basalts. This interac-tion has been comprehensively studied by experimentaland computation techniques (Chapter 3). In the simula-tion of metasomatic processes, this problem yieldsrather complicated results and displays some relation-ships that are rarely observed in problems with simplermineral compositions, for example, in the simulation ofacid leaching of granitoids. Ore deposition during cool-ing is examined in oceanic hydrothermal systems,which provide typical examples of hydrothermal–sedi-mentary ore formation (Chapters 3, 4).
The application of various methods is illustrated inthis chapter by calculations in the 15-element multisys-tem H–O–K–Na–Ca–Mg–Fe–Al–Si–C–S–Cl–Cu–Zn–Pbincluding 45 minerals of fixed composition and55 solution species.1 This system is described in detail
1 Mineral symbols used in the diagrams of this chapter are given inTable 18.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S165
in section 3.4.2. For the calculation of isothermal mod-els (T = 350°C and P = 500 bar), the initial solution wasrepresented by seawater modified to reach an equilib-rium state under these conditions (accounting for anhy-drite and brucite precipitation from seawater undersuch conditions). The models of hydrothermal solutioncooling used reconstructed smoker solutions as an ini-tial hydrothermal solution. The compositions of rocksand solutions used in calculations are shown in Table 1.
Classification of chemical reactors. The equilibriumdynamic models considered here have much in com-mon with apparatuses used in chemical processes. Thetheory of chemical reactors is thoroughly developed inchemical engineering (Denbig, 1968; General Chemi-cal…, 1977) and some of its aspects can be used for thesimulation of geologic processes. In particular, the clas-sification of chemical reactors is important for the fol-lowing discussion.
Chemical reactors are classified with respect to anumber of criteria, including method of filling, charac-ter of movement of the reaction medium, thermalregime, etc. By the method of filling, chemical reactorsare divided into periodic, which are filled before the
beginning of the chemical process and emptied after itscompletion, and continuous (flow) reactors, whichinvolve the input and output of substances in the courseof the process. By the character of material movement,the continuous reactors are subdivided into mixed andreplacement reactors. The ultimate cases are referred toas ideal mixing and ideal replacement (plug) flow reac-tors. Real replacement reactors differ from the idealpattern for many reasons: along-axis diffusion, hydro-dynamic dispersion, hydrodynamic wall effects, etc. Bythe thermal regime, isothermal, polythermal, adiabatic,and other types of reactors are distinguished. There is aclass of composite reactors, whose construction iseither represented by cascades of reactors or separatedinto steps (cells, rectification plates, etc.). Figure 2 pre-sents the principal structure of main reactor types.
Processes in each type of reactors are described by aspecific system of equations. It is supposed (Principlesof…, 1991, p. 79) that continuous replacement flowreactors (Fig. 2c) can be approximated by a cellularmodel, as a cascade of ideal mixed reactors (Fig. 2d)with a great number of cells.
Table 1. Compositions of the basalt, alaskite granite, and solutions that were used for the numerical simulation of infiltrationmetasomatic columns and ore deposition during cooling
Component Basalt Alaskite granite Seawater Seawater modified by heating
Hydrothermalsolution
H 0 0 0.002114 0.01535 0.117295
O 27.32542 30.4 0.120401 0.098189 0.060188
K 0.0382 2.3 0.0099 0.009904 0.028856
Na 0.7936 – 0.468 0.468079 0.467921
Ca 2.0837 – 0.0103 0.00539 0.024077
Mg 1.9606 – 0.0532 0.046397 1.0 × 10–7
Fe 1.4364 – 0 0.000154 0.001175
Al 2.9756 2.3 0 2.51 × 10–8 8.11 × 10–6
Si 8.4231 12.9 0 0.000027 0.013992
C 0.0088 – 0.002327 0.002328 0.006694
S 0.0209 – 0.02823 0.023114 0.008247
Cl 0 0 0.5459 0.5459 0.5459
Cu 0.00116 – 0 2.16 × 10–7 1.3 × 10–6
Zn 0.0011 – 0 2.1 × 10–7 0.000025
Pb 0.000012 – 1.01 × 10–7 1.0 × 10–7 3.0 × 10–7
H2O – – 55.51 55.49722 54.59447
Note: The composition of modified seawater was obtained by calculating the equilibrium 1 kg seawater +0.1 g basalt at 350°C and 500 bar.The composition of the hydrothermal solution was calculated within the model of an oceanic hydrothermal system at Tmax = 350°C,P = 500 bar, and R/W = 0.496 (Section 4.1.1).
S166
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Note also that a steady-state regime in continuousreactors differs from start and stop regimes, which aredescribed by different systems of equations.
2.1. Methods of Simulation of Infiltration Metasomatic Processes
2.1.1. Infiltration metasomatism and its propertiesin the Korzhinskii theory
In modern petrology and geochemistry, the interpre-tation of infiltration metasomatism is based on the the-ory of metasomatic zoning developed by Korzhinskii(1969). He postulated a few main features of infiltrationmetasomatism (logical scheme) and analyzed the sys-tem of equations (theoretical model) describing this
scheme. This analysis revealed a number of conse-quences that must control the properties of metaso-matic columns produced by fluid infiltration. If a ther-modynamic model of hydrothermal and metasomaticprocesses is aimed at reproducing metasomatic zoning,and its parameters correspond to the postulates of theKorzhinskii theory, its consequences must be satisfiedin the numerical model.
The following conditions were postulated byKorzhinskii for the development of the theory.
I Solution percolates through a permeable homoge-neous medium. Solution movement is laminar and cor-responds to plug replacement, and there is no hydrody-namic dispersion.
(a) (b)
(Ò)
(d)
Fig. 2. Classification of chemical reactors with respect to the method of filling and the character of reaction medium movement.(a) Periodic reactor, (b) continuous (flow-through) reactor with ideal mixing, (c) continuous (flow-through) reactor with idealreplacement, (d) cascade of reactors (step reactor) with ideal mixing.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S167
II The solution reacts instantaneously with solidphases, and all portions of solid phases are equallyaccessible for the reaction.
In addition, it was implicitly assumed that theamounts of dissolved components per unit volume ofthe porous medium are negligible compared with theircontents in the solid phase.
These conditions implied the following traits of anisothermal metasomatic column without minerals ofvariable composition (this case is the most appropriatefor testing the methods of thermodynamic modeling).
1 At the onset of infiltration, a metasomatic columnis immediately formed in its complete and final formand only subsequently expands.
2 Metasomatic zones grow linearly with time. Thepropagation rate of each boundary is higher than that ofthe preceding one, which provides the stable structureof the column.
3 The percentages of minerals are constant over thewhole length of each metasomatic zone.
4 The composition of the pore solution changesabruptly at the boundaries between zones and is con-stant within any metasomatic zone.
Such a peculiar dynamic regime was referred to asthe conditionally steady-state system (Korzhinskii,1979). Its properties are constant when scaled by x/t,where x is the distance from the beginning of the col-umn, and t is the time.
The simulation of infiltration metasomatic zoning, ifit satisfies conditions I and II, must yield a column withcertain properties specified by consequences 1–4.2 Thisprovides an opportunity to test the internal consistencyand applicability of the methods of numerical modeling.
2.1.2. Method of the degree of reaction progress by Helgeson
A computer method for the thermodynamic model-ing of irreversible water–rock interactions was first pro-posed by Helgeson (1968). He applied the principle ofpartial equilibrium by Barton. The global parameter ofthe irreversible process was represented by the degreeof progress (ξ) of the reaction limiting the rate of theentire process (this is usually a reaction with initialminerals or initial rock on the whole), whereas all otherreactions were regarded as equilibrium ones and could,consequently, be calculated by the methods of equilib-rium thermodynamics. The rate-limiting reactionchanges the composition of the equilibrium portion ofthe system in accordance with the stoichiometry of theinitial substances in it. This relation was explicitlydescribed by Karpov (1981, pp. 124–126). The quantityξ is an implicit function of time; this function is monot-onous, which allows us to calculate the chemical path-
2 The conditions and consequences were not numbered byKorzhinskii (1969), but was done in this paper to facilitatedescription.
way of the process in mi versus ξ coordinates, where miare the amounts of dissolved substances and mineralproducts of the process. Figure 3 shows a pathway ofseawater–basalt interaction calculated by this proce-dure for 350°C and 500 bar.
Analysis of the method of thermodynamic modelingthat was described in detail by Helgeson (1979) showsthat it deals with systems closed with the respect to thesolution and does not account for any spatial movementof material. In terms of equilibrium dynamic models,the Helgeson method can be considered as a series ofequilibrium states of a multisystem whose compositionis changed by the continuous addition of initial rock. Interms of the theory of chemical reactors, it correspondsto a periodic (i.e., not flow-through) reactor (Fig. 2a).The Helgeson method is thus not designed for the cal-culation of metasomatic processes, which include masstransfer in space as one of the most important proper-ties. The results of calculations shown in Fig. 3 do notcorrespond to metasomatic zones but rather describemineral associations changing in time within a fixedreaction volume, for instance, in an autoclave. There-fore, condition I of the Korzhinskii theory is not met insuch a model, and the results of application of thismethod are not consistent with all its consequences. Itcan be clearly seen in Figs. 3a and 3b that the propor-tions of minerals and the composition of the solutionvary with ξ within the stability field of a certain mineralassemblage. The interpretation of this method sensulato as a means to reproduce spatial and temporal rela-tions in metasomatic processes may lead to controver-sies and even, as will be shown below, erroneous con-clusions. Attempts of such expanded interpretationsappeared in the literature and even gained some popu-larity (Karpov, 1981; Kashik, 1989).
This problem was first explored by Fritz (1975),who pointed out the main factor hampering the applica-tion of the Helgeson method to spatial problems. In thismethod the products of previous reaction steps remainin the sphere of interaction and can react during the fol-lowing steps with the altered solution (back-reaction),whereas the mineral products of reactions are leftbehind and do not further react with the solution portionthat formed them in the infiltration metasomatic pro-cess. The importance of this difference depends on thechemistry of a particular system, and this problem willbe discussed below.
2.1.3. Method of a step flow-through reactor
In order to adapt the Helgeson approach to the sim-ulation of infiltration metasomatism, Fritz (1975) pro-posed a method referred to as systeme reversale. Hemodified the algorithm of the degree of reactionprogress method. At each step of calculation, the bulkcomposition of the system changed not only at theexpense of the limiting reaction proportionally to ξ, butalso by the removal from the system of the solid phasesformed during the previous step of calculation. In such
S168
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
a way, the solid and liquid products of interaction wereseparated, which is characteristic of an infiltration pro-cess. With such a separation of reaction products, theFritz method does not correspond to a periodic reactor ora continuous completely mixed reactor (Figs. 2a, 2b). Thesystem of differentiation equations used by Fritz (1975)is analogous to that proposed by Helgeson (1968) andimplies infinitely small ξ increments, which corre-sponds to a process in a continuous flow reactor(Fig. 2c). A numerical implementation of the methodwith finite, although small, ξ increments is equivalentto a step flow reactor (Fig. 2d). It is essential that amaterial of fixed composition (rock) is added to the sys-tem at each ξ increment. Hence the process simulatedin this method corresponds to the movement of a single(first) portion of solution through a step reactor filledwith fresh rock, i.e., to the regime of an incipient flowreactor.
The results obtained by such a method for the modelof infiltration metasomatism in the seawater–basalt sys-tem are shown in Fig. 4. A comparison of Figs. 3 and 4reveals some differences between the model results
(ignoring the fact that the variable ξ has a differentphysical meaning in the Helgeson and Fritz methods,and they are, strictly speaking, not comparable). Themost important difference is related to the ξ range ofanhydrite stability, which extends in the Helgesonmethod up to ξ > 200 g/kg. In addition, the Helgesonmethod implies much higher sulfur contents at high ξ.This discrepancy is related to the fact that sulfur is rap-idly removed from the system as a solid product (anhy-drite) in the case considered by the Fritz method,whereas it is retained in the reaction sphere in the Hel-geson approach. There are also differences in thesequence of mineral assemblages (in the example con-sidered, the Helgeson method yields a larger stabilityfield of sericite, Fig. 3a, but produces no hematite andmagnetite, which are predicted by the Fritz method),and in some minor details. In general, the fields ofphase assemblages calculated by the Helgeson methodare more extended along the ξ axis compared with theFritz method.
The two methods must yield identical results undercertain conditions, which can be easily estimated from
(a)
(b)
100
80
60
40
20
0
%
0.05
0.04
0.03
0.02
0.01
01 11 21 31 41
Rock/water, g/kg
mol/
kg
K
Ca
Mg
Fe
Si
S(VI)
S(II)
Act80
Tr
Chl50
Cch
Srp
Tlc
Ep
Anh
Prl
Ser
Ab
Qtz
Act80
Ep
Anh
Ab
Ser
Qtz
Cch
Chl50Tr
Tlc
Srp
Prl
Fig. 3. Calculation of seawater–basalt interaction by the Helgeson method at T = 350°C and P = 500 bar. (a) Mineral assemblagesand (b) solution composition.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S169
the analysis of the reason for the aforementioned dis-crepancies. The Helgeson and Fritz methods are identi-cal with respect to solution composition and sequenceof mineral assemblages when all the minerals occurringin the initial rock dissolve congruently and the compo-nents introduced with the solution do not precipitate(canonical column according to Shvarov et al., 2000).Under such conditions, nothing is lost from the solutionmoving through the flow reactor as minerals leftbehind. However, even in such a case, the percentagesof minerals calculated by the Helgeson and Fritzmethod are different, because the mineral compositionin the former approach is summed over all the previouszones of the column for the given step.
Metasomatic zoning produced by acid leachingapproaches the canonical column. Figure 5 presents anexample of calculations for the interaction of granite(chemical composition of the rock is given in Table 1)with sulfuric acid solution. This example almost com-pletely satisfies the aforementioned necessary condi-tions, but it can be clearly seen that the percentages of
minerals in the assemblages calculated by the Helgesonmodel are variable (Fig. 5a), and the muscovite stabilityfields have different sizes in Figs. 5a, 5c.
Thus, there is a simple criterion (necessary but notsufficient, as will be shown below) for the applicabilityof the Helgeson method to simulate at least the qualita-tive structure of an infiltration metasomatism column:the absence of incongruently dissolving minerals andphase replacement in the model column. The majorityof published model columns obtained by the Helgesonmethod for weathering and hydrothermal alteration donot meet this criterion.
The Fritz method complies with condition II of theKorzhinskii theory and does not violate condition I.However, it can be readily seen that the results of calcu-lations shown in Figs. 4a, 4b, and 5d are not consistentwith consequences 3 and 4: the proportions of mineralsand the compositions of solutions change within indi-vidual metasomatic zones, i.e., the columns are notconditionally steady-state. For instance, an examplewith alaskite granite (Fig. 5d) shows distinct variations
(a)100
80
60
40
20
0
%
Ab
Hem
Qtz
Cch
Chl50Act80
Tlc
Srp
Prl Mag
Anh Ep
Chl75
Ep60
Ep75
(b)0.05
0.04
0.03
0.02
0.01
01 11 21 31 41
Rock/water, g/kg
mol/
kg
K
Ca
Mg
Fe
Si
S(VI)
Act80
Mag
Chl50
CchSrp
Tlc
Ep
AnhPrl
Hem
AbQtz
Chl75
Ep50
Ep75
Fig. 4. Calculation of seawater–basalt interaction by the Fritz method at T = 350°C and P = 500 bar. (a) Mineral assemblages and(b) solution composition.
S170
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
in pH and K concentration in the solution within thefield of the quartz–kaolinite assemblage. Consequence2 (column growth with time) is in principle not repro-duced by this method. The physical presentation of themethod as a movement of a portion of solution througha rock suggests that consequence 1 (instantaneous for-mation of the whole metasomatic column) is also notprovided. The fundamental reason for this discrepancyis related to the fact that condition II of the Korzhinskiitheory implies instantaneous equilibrium between thefirst infinitely small portion of solution and the initialrock (frontal zone of the column), i.e.,
(1)
where R/W is the mass ratio of the reacting rock andwater. The use of the Fritz method in its analytical formis equivalent to
(2)
and for numerical solutions with finite ξ increments:
(3)
R/W( )ξ 0→lim ∞,=
R/W( )ξ 0→lim const,=
R/W( )ξ 0→lim 0.=
In the simulation by the Fritz method, a movingbatch of solution reacts with new rock portions andaccumulates components partitioning into solution (forinstance, potassium, in the calculations illustrated byFigs. 4b and 5d), which is in conflict with consequence 4.This results in variability in the proportions of mineralswithin an individual metasomatic zone (contrary toconsequence 3). Thus the Fritz method does not repro-duce any consequence of the Korzhinskii theory.
Note finally that the physical presentation of a flowreactor was not used in Fritz’s study. A step flow reactorwas first explicitly used for the thermodynamic model-ing of the hydrothermal process by Grichuk andBorisov (1983). However the reactor considered bythese authors was a polythermal one, and condition 1,plug replacement, was not met in the problem formula-tion.
2.1.4. Method of multiwave step flow reactor (MSFR)
An obvious limitation of the Fritz method is thatinteraction between a single portion of solution withfresh rock (initial regime) is simulated. One way to
–7–6
log(m
ol/
kg)
log(rock/water)
–5
–3
–1
1
3
5
7
0
(b)
–5 –4 –3 –2 –1
pH
(a)
20
40
60
80
100
Min
eral
, %
–6log(rock/water)
(d)
–5 –4 –3 –2 –1
(Ò)
pH
K
Al
Si
Kln
Mu
Qtz
Fig. 5. Interaction of an alaskite granite with hydrochloric acid solution at T = 250°C, P = 500 bar, and 0.001 mol/kg HCl. (a) Hel-geson method, mineral assemblages; (b) Helgeson method, solution composition; (c) Fritz method, mineral assemblages; and(d) Fritz method, solution composition.
Kln Kln
Qtz Qtz
Mu
McMc
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S171
overcome this limitation is to use the method of a mul-tiwave step flow reactor (the term wave is used synon-ymously with the term solution batch by Heinrich,1990). The idea of simulating metasomatic alterationsby passing many portions of solution through a systemof cells (steps of a flow reactor) containing certainamounts of rock was first used by Fouillac et al. (1977).This method was referred to as systeme renouvelle.These authors modified Helgeson’s algorithm repre-senting a weathering profile as an array of 20 cellsthrough which solution was transported. They assumedthat only part of the solution occurring in a given cellwas moved to the next cell during the next calculationstep (the fraction of solution transported was specifiedby the parameter ζ), and the ratio ζ/ξ was used as a sim-ilarity criterion in their modeling. Fouillac et al. (1977)did not discuss the physical meaning of this feature oftheir approach, but it is reasonable to suggest that thevariable ζ reflects the influence of hydrodynamic dis-persion. In terms of the theory of chemical reactors,such a method corresponds to a cascade of partiallymixed reactors. Fouillac et al. (1977) demonstrated thatthe results obtained in their systeme renouvelle arequalitatively different from those of the Helgeson andFritz approaches. Since the introduction of hydrody-namic dispersion violates condition I of the Korzhinskiitheory, it is impossible to test the applicability of thismethod by the agreement with the consequences of thetheory. It should also be noted that the use of a logarith-mic scale for ζ/ξ in a cell array violates the condition offlow continuity and cannot be physically interpreted.3
3 The study of Foullac et al. (1977) was unfortunately not contin-ued. According to the Science Citation Index, it has been cited16 times between its publication in 1977 and 1995, and always insections containing overviews of previous work.
A multiwave flow reactor was first explicitly used in athermodynamic model by Grichuk (1988), who simu-lated convection in hydrothermal systems rather than ina metasomatic column (Section 2.2).
The MSFR method simulates a metasomatic col-umn by passing many solution portions (waves)through the reactor, and, in contrast to the Fritz method,it has to be free of the effect of accumulation of easilymobilized components in the solution. Moreover, sincethe column is divided into an increasing number ofsteps affected by a great number of waves, the model isreminiscent of a continuous reactor (General Chemi-cal…, 1977), which must give better agreement withthe Korzhinskii theory.
Figures 6 and 7 present the results of MSFR appli-cation to the calculation of an infiltration column in theseawater–basalt system. Figure 6 shows that the move-ment of the boundaries of metasomatic zones is approx-imately linear with time (i.e., with the number of solu-tion portions passing through the column) (conse-quence 2). The calculation produced continuous growthof a series of eight metasomatic zones.4 As can be seenfrom Fig. 7b, the composition of the solution changesat the boundaries of metasomatic zones, which is inagreement with consequence 4 of the Korzhinskii the-ory. However, the calculation of mineral composition(Fig. 7a) revealed, in this case, unexpected variations inthe proportions of minerals within a zone. A compre-hensive analysis of these results showed that such vari-ations are artifacts of the computation method.
4 There is an additional zone, Chl + Qtz + Hem + Anh + Tr, but it isrepresented in all waves by no more than one reactor and is prob-ably a ghost zone.
80
70
60
50
40
30
20
10
Wav
e no.
100 20 30 40 50 60Reactor no.
Zone
Ä
B
C
D
E
F
G
H
H
G
FEDCBÄ
0
Fig. 6. Growth of metasomatic zones with time calculated by the method of a multiwave flow step reactor (MFSR) for the interactionof seawater with basalt at T = 350°C, P = 500 bar, 50 steps, and 80 waves (solution portions). Symbols are calculated points. Meta-somatic zones: (A) Cch + Srp; (B) Cch + Qtz + Prl; (C) Chl50 + Qtz + Prl + Hem; (D) Chl50 + Qtz + Hem + Anh; (E) Chl50 + Qtz +Hem + Prl + Anh; (F) Chl75 + Qtz + Ep + Act80 + Ab; (G) Chl75 + Ep + Act80 + Wai; and (H) Chl75 + Ep + Act80 + Ab + Wai.
S172
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
0 0.6
mol/kg
Rock
/wat
er (
g/k
g)
0.0
5
25.6
(b)
01
mol/kg
Wav
e no.
0.0
5
21
(e)
01
mol/kg
Wav
e no.
0.0
5
121
(c)
6.9
13.1
19.4
Rock
/wat
er (
g/k
g)
0.0
4
0.0
3
0.0
2
0.0
1
K Ca
Mg
Fe
Si
S(V
I)
S(I
I)
0.0
4
0.0
3
0.0
2
0.0
1
21
41
61
81
101
K Ca
Fe
Si
11
0.0
4
0.0
3
0.0
2
0.0
1
(d)
K Ca
Mg
Fe
Si
S(V
I)
(a)
100
80
60
40
20 0
%
0.6
25.6
6.9
13.1
19.4
Rock
/wat
er (
g/k
g)
0.2
2.8
5.2
7.8
10.0
100
80
60
40
20 0
%
Act
80
Tr
Chl7
5
Chl5
0
Cch
Srp
Ep50
Ep75
Ep
Anh
Prl
Ab
Qtz
Hem
Qtz
Ab
Srp
Prl
Ep
Ep75
Ep60
Chl7
5
Act
80
Tr
Chl5
0
Cch
Hem
Chl5
0
Cch
Srp
Anh
Prl
Qtz
Hem
Chl5
0
Cch
Anh
Qtz
Prl
Hem
Srp
Anh
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S173
The reason for these variations lies in the fact thatreactor steps have fixed dimensions (i.e., the masses ofinitial rock) in the MFSR method. The mineral andchemical compositions of a step containing a boundarybetween metasomatic zones appear to be combinedfrom the compositions of the contacting metasomaticzones. The solution passing across the boundary shouldbe out of contact with the minerals of the previous zone,but equilibrium is calculated for the bulk compositionof such a reactor step, which results in the occurrenceof a back-reaction (though small-scale). The calcula-tion of equilibrium for such a boundary step will yielda spurious result in two cases: (a) if incongruent disso-lution or replacement of phases takes place on theboundary, and (b) for a dissolution boundary, if theamount of the dissolved phase is not sufficient for thesaturation of the given amount of water (in such a case,the boundary will move to the next reactor step in thefollowing stage of calculations). This feature causesoscillations in the composition of a solution passingthrough a fixed cross-section of the reactor, even if thiscross-section stays within a single metasomatic zone(Fig. 7c). In other words, although the composition ofeach model water portion passing through a metaso-matic zone remains constant, the adjacent water portionoccurring simultaneously in the same zone may have adifferent composition.
An increase in the number and a decrease in the sizeof steps (the fraction of boundary steps yielding spuri-ous results should diminish) do not in fact improve thestability of calculated results, because the probability ofsolution undersaturation increases owing to reason (b),and reason (a) is not fully eliminated. As can be seenfrom Figs. 7d and 7e, in such a case variations in thesolution composition decrease, but variations in theproportions of minerals in the zones become more pro-nounced. This may even produce spurious repeatedzones (intercalation of monomineralic clinochlore andclinochlore + serpentine zones in the left part ofFig. 7d). An increase in the prescribed mass of rock perstep reduces the variability in mineral composition(Fig. 7b), but this lengthens the computation time forthe following reason. The MSFR method may providespurious mineral assemblages for initial waves, if sev-eral metasomatic zones appear to be combined in a sin-gle step of the reactor. Only after the passage of severalwaves of solution will the column expand, these spuri-ous assemblages vanish, and the movement of zoneboundaries become linear in time (Fig. 6). An increasein the mass of rock per step protracts the initial non-steady-state period of calculations, sometimes very sig-
nificantly (calculations may appear physically impossi-ble on a personal computer).
Thus, the use of the MSFR method for the simula-tion of an infiltration metasomatic column fulfils conse-quence 2 and, approximately, consequences 3 and 4.The computational scheme is not sufficiently efficient,because a great number of waves must be calculated.These disadvantages are intrinsic properties of themethod and cannot be avoided. The only regulatedparameter of the method is the initial mass of rock perstep. An increase in this mass is physically equivalentto an enhancement of hydrodynamic dispersion, whichviolates the condition of plug replacement (condition Iof the Korzhinskii theory), and a decrease in the massleads to the limiting transition expressed by Eq. (3),which violates condition II.
2.1.5. Direct use of consequencesof the Korzhinskii theory in algorithms
of the simulation of infiltration metasomatism
The properties of infiltration metasomatic columnsdescribed by the Korzhinskii theory can be directlyused to construct an algorithm for equilibrium dynamicmodeling. For instance, Ivanov and Borisov (1980)used the qualitative stability of zone sequence in a col-umn as a criterion in the reconstruction of the composi-tion of a metasomatic solution.
Alexeyev (1985) proposed an approach to the calcu-lation of metasomatic columns by means of the thermo-dynamic simulation of reactions at zone boundaries anddetermination of the rate of zone boundary movementas the main parameter of comparison according to con-sequence 2 of the Korzhinskii theory (method ofboundary reactions, Shvarov et al., 2000). In fact, themethod proposed by Alexeyev represents a metaso-matic column as a step flow reactor (though the termwas not used in his study), but reactions are calculatedonly in those elementary volumes (steps) where theboundaries of metasomatic zones are situated at thegiven step of the simulation. In such a way, conse-quences 3 and 4 are automatically satisfied. The step-wise method of computation employed by Alexeyev(1985) is formally not consistent with consequence 1.
The application of this method was illustrated by theexample of acid leaching of alaskite granites, whichgives rise to a simple metasomatic column with three orfour metasomatic zones depending on the compositionof the initial solution. The simulation of more complexsystems leads to some problems that were not consid-ered by Alexeyev (1985). For instance, if a solution of
Fig. 7. Calculation of seawater–basalt interaction by the method of a multiwave flow step reactor (MFSR) at T = 350°C and P = 500 bar.(a) Mineral assemblages, metasomatic zones for the case with 50 g of rock on reactor steps and wave no. 80; (b) changes in thecomposition of the 80th solution portion during its passage through the metasomatic column; (c) compositions of solution portionspassing through the 40th step of the reactor, the mass of rock in each reactor step is 50 g; (d) mineral assemblages for the case with5 g of rock in each reactor step and wave no. 20; and (e) compositions of solution portions passing through the 40th step of thereactor for the case with 1 g of rock in each reactor step.
S174
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
complex composition reacts with an excess amount ofinitial rock, the first step of calculations by the Alex-eyev method often does not provide the correct set ofminerals in the back zone (compare with the above-described method of the degree of reaction progress atξ > 1, Fig. 3). As a result, the subsequent calculation ofthe structure of the whole column will be erroneous,and Alexeyev did not provide a way to recognize sucha discrepancy.
The method of boundary reactions was furtherdeveloped by Shvarov et al. (2000). These authorsshowed that a zone boundary reaction can be written as
(4)
where L and S denote the compositions of the solutionand solid phase, respectively, and the subscripts i andi + 1 refer to the zone numbers, back and frontal withrespect to the given boundary. Reaction (4) reflects thecondition of formation of the back zone during theinstantaneous reaction of Li with the mineral substancesof the frontal zone.
Shvarov et al. (2000) showed that there are twotypes of boundaries between zones in a metasomaticcolumn and reactions on these boundaries must be cal-culated differently.
(1) If the mineral assemblage of Si + 1 includes all theminerals of Si (congruent dissolution), then Li + 1 is inequilibrium with Si + 1 and Si simultaneously (conditionof equilibrium downflow). It can be easily shown thatthe number of minerals in Si + 1 is higher by one thanthat in Si. The boundaries that comply with this condi-tion are referred to as simple, and the columns includ-ing only simple boundaries are termed canonical col-umns. In the case of a simple boundary, the calculationof equilibrium for the chemical system composed of Liand Si + 1 (excess solid phase) yields a solution corre-sponding to Si + 1. A comparison of the compositions ofphases defines the parameter t:
(5)
The reciprocal parameter (i.e., 1/t) is the ratio of therate of zone boundary movement to the velocity ofsolution flow.
(2) If the mineral assemblage of the back zone con-tains at least one mineral missing in the frontal zone(indicating incongruent dissolution), the zone boundaryis referred to as a reaction boundary, and the columnwith such boundaries is noncanonical. The reactionboundary is associated with two (or more) coupledreactions. The mineral products of one of these reac-tions are initial compounds for another reaction. Forsuch a boundary, the calculation of equilibrium in the(Li + excess Si + 1) system results in extra mineral spe-cies and an incorrect t value.
tLi Si 1+ tLi 1+ Si,+ +
tSi Si 1+–Li Li 1+–---------------------.=
The following algorithm was proposed for the cal-culation of a canonical column composed of n zones(S0 = 0):
.
(1) Equilibrium is calculated for the system (L0 +excess Sn). The t value of the disappearance of the firstmineral from the phase assemblage of the frontal zoneis determined. The obtained solid phase is regarded aszone (n – 1). This operation is repeated for zone (n – 1)and all the following zones to obtain a first approxima-tion for the mineral assemblages of all zones of the col-umn.
(2) Passing from the back zone to the frontal one, thecomposition of the solution is calculated for each zone.
(3) Passing from the frontal zone to the back one, theamounts of minerals in the mineral assemblages ofzones are calculated.
The appearance of new mineral assemblages ischecked in stages (2) and (3). If they do form, the cal-culation returns to the previous step in order to refinethe zone sequence.
The correctness of the resulting column is generallycontrolled by t values. The following condition must bemet for all boundaries within the column:
ti ≤ ti + 1. (6)
If this condition is not met, the column is not canon-ical, i.e., there are reaction boundaries. The violation ofcondition (6) implies that there is a ghost zone in thecalculated model column (Lichtner and Balashov,1993), whose back boundary moves faster than thefrontal one. Shvarov developed a special procedure forthe calculation of such boundaries taking into accountthe incompleteness of reactions occurring on the backboundary of a ghost zone. It should be noted that theoverwhelming majority of infiltration metasomatic col-umns contain reaction boundaries and are not canoni-cal.
Shvarov et al. (2000) and Grichuk and Shvarov(2002) compared the results of calculations of metaso-matic columns and showed that, with respect to compu-tational efficiency, the method of boundary reactions isfar superior to the MSFR method. It is devoid of arti-facts related to the stepwise treatment of the process inthe MFSR method. This can be illustrated by the exam-ple presented in the paper. In particular, even for waveno. 1000, the column calculated by the MFSR methodcontained an extra zone (probably a ghost zone), whichwas lacking in the strict solution obtained by themethod of boundary reactions.
2.2. Modeling of Metasomatic Processes in Extended Hydrothermal Systems
With respect to the character of processes, extendedhydrothermal systems with percolation through frac-
L0 L1 L2 … Ln
S0 S1 S2 … Sn
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S175
ture networks are remarkably different from the above-described process of formation of infiltration metaso-matic columns. The most fundamental difference is thatconditions I and II of the Korzhinskii theory are notapplicable for large-volume systems. Indeed, percola-tion through fractures in rocks with a strongly varyingpermeability gives rise to significant hydrodynamicdispersion, and plug replacement (condition I) is notrealized. The velocity of solution movement in frac-tures can be much higher than the rate of penetrationinto the blocks of rock and metasomatic alteration ofthe inner parts of the blocks. This results in the forma-tion of relicts of unaltered rocks and concentric zoningof metasomatic alterations in the blocks (violation ofcondition II). Therefore, metasomatic processes in suchhydrothermal systems are in fact a superposition ofinfiltration metasomatism in the direction of solutionflow and diffusion metasomatism in the perpendiculardirection, complicated by strong hydrodynamic disper-sion. This situation obviously cannot be reduced to theformation of an ideal infiltration metasomatic column.The rigorous analysis of such processes presents con-siderable difficulties, and, when necessary, the simula-tion of such geologic processes is carried out using var-ious approximations.
A specific class of problems of simulating processesin extended systems is the model of formation of ore-bearing solutions in convective hydrothermal systemswith an exogenous source of solution (atmospheric ormarine) and an endogenous source of energy. In such acase, the metal loading of solution is controlled bywater–rock interaction in a thermally heterogeneousmedium with prevailing fracture percolation. The best
studied example of such problems is the model of hydro-thermal systems in the oceanic crust (Section 3.1.4); sim-ilar characteristics display the models of epithermaland stratiform deposits (Heinrich et al., 1995; Ilchikand Barton, 1997) and even of the greisen process(Korotaev et al., 1992).
Grichuk and Borisov (1983) and Grichuk (1988)proposed use of the MSFR method for the simulation ofsuch systems. A convective hydrothermal system is pre-sented in this method as a polythermal step reactor(Fig. 8 analogous to Fig. 2d), in which a partial equilib-rium is reached in each step between the solution andthe rock. The solution flow path is divided into stepswith a fixed temperature increment (10–50°C). Aparameter of partial equilibrium has to be prescribedfor each step. This parameter specifies irreversibleinteractions between the solution and rock and is repre-sented by the rock/water ratio (R/W). This is, in fact, asimilarity criterion of the model, equal to the mass pro-portion of fresh rock reacting with the given solutionportion at the given step to the mass of this portion.
The mass of the water portion is given by the expres-sion
(7)
where S∅ is the effective section, v is the velocity ofsolution movement, ρw is the density of the solution, Δtis the time of passage through the reactor step (temper-ature zone along the flow path), and Q is the solutiondischarge rate. The mass of reacted rock is
(8)
W S∅vρwΔt QΔt,= =
R v rS frΔt,=
100°C
225°C
350°C
150°C
200°C
250°C350°C
375°C
AB
C
C
Magma chamber
Fig. 8. Simplified model of a hydrothermal system as a flow step reactor: (A) region of the downwelling limb, (B) region of theupwelling limb, and (C) zone of ore deposition.
S176
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
where Sfr is the total surface of fractures within the rockblock (reaction step), and vr is the specific rate of reac-tion between the solution and fresh rock. It can be eas-ily seen that the R/W value is not directly dependent onΔt and the models utilizing this parameter can be scaledby time.
The parameters of Eqs. (7) and (8) cannot yet be pre-cisely determined, especially vr, and they cannot beused as a basis for the construction of quantitative mod-els (Rafal’sky, 1993). Grichuk and Borisov (1983) andGrichuk (1988) proposed a method to circumvent thisdifficulty. The R/W values can be constrained by indi-rect geochemical methods, in particular, from isotopicsystematics (Spooner et al., 1977; Norton and Knight,1977; etc.). In particular, for modern oceanic hydro-thermal systems, a procedure was developed for theestimation of the integral ΣR/W value that was reachedin the hydrothermal system as a whole from the concen-trations of mobile elements (see Section 3.2.1 for moredetail). For the hydrothermal systems studied, thesevalues are between 0.5 and 2.0 kg of rock per one kilo-gram of solution. In the context of the MSFR method,a drawback of such estimates is that the ΣW/R valuemust be divided in the model between the contributionsof each step. Moreover, the ΣW/R of the system canvary with time.
The following simplifying assumptions can be madefor the model of a convective hydrothermal system:(a) the model is steady-state with respect to the positionof temperature zones and velocities of solution move-ment; (b) the times of passage of temperature zones areapproximately equal, and the time of solution passagethrough the ith step, Δti, is proportional to ΔTi; and(c) the permeabilities of rocks do not vary significantlyin space and time. Given these assumptions, the param-eters Q and Sfr can be taken to be equal in the reactorsteps and time-invariant (although their exact values arenot known). The only time- and temperature-dependentparameter appearing in the expression for R/W is thenthe specific rate of reaction with fresh rock, vr.
The temperature dependence of R/W can be writtenin the notation used in this work as
(9)
where T0 and Tf are the temperatures of the beginningand end of the process in the descending flow, ΔT is the
temperature increment between the reactor steps, isthe average rate of the reaction in the temperature inter-val corresponding to step i, and n is the number of stepsin the reactor. The temperature dependence of the rateof the solution–rock reaction can be specified by theempirical equation of Wood and Walther (1983) andWalther and Wood (1989), who demonstrated that,under far from equilibrium conditions, the rate of alu-minosilicate dissolution in near neutral solutions nor-
ΣR/WS fr
Q------ v r T( ) Td
T0
T f
∫S fr
Q------ v ri T( )ΔTi[ ],
i 1=
n
∑= =
v ri
malized to the number of oxygen atoms in the formulais governed by an Arrhenius-type equation:
logk = – 6.85 (g-atom O/cm2 s), (10)
where k is equivalent to vr in our notations. The firstterm of Eq. (10) corresponds to the quantity2.301Ea/RT in the Arrhenius equation, which gives anactivation energy of 10.5 kJ/g-atom O. The numericalintegration5 of this expression over temperature inter-vals allows us to obtain the vri values and then distrib-ute the value ΣR/W between the reactor steps using thesimple expression
(R/W)i = Σ(R/W) . (11)
The dependency corresponding to Eq. (11) is shownin Fig. 9. A more rigorous hydrodynamic calculation ofthese relations for a convective system (Tutubalin andGrichuk, 1997a) without assumptions (a) and (b)yielded rather similar results.
Time dependence of R/W. A metasomatic zone prop-agates linearly with time normal to the fracture sur-faces, if its growth is controlled by the rate of a surfacereaction, as was supposed above. However, it is obviousthat along with the expansion of the layer of metasom-atized rocks, diffusion mass transfer in this layer willplay an increasing role and the process of interactionwith fresh rock will be slowed down. The theory of dif-fusion metasomatism (Korzhinskii, 1969) implies aparabolic growth law: the width of metasomatic zonesincreases proportionally to the square root of time, τ. Ingeneral it is impossible to predict a transition from alinear growth regime to a parabolic one. For a model ofa hydrothermal system of considerable volume and dis-charge rate with a Σ(R/W) of about 1, it can be sug-gested that the linear stage is already completed afterthe passage of a few solution portions. Therefore, thegrowth rate of the layer of metasomatic rocks near frac-ture walls will decrease:
(12)
where is the rate in the initial moment of the pro-cess, and x is the thickness of the altered layer. Thestepwise approximation of the interaction process usedin the MSFR method implies that an increase in R/W ineach reactor step can be described for the jth wave bythe simple expression
(R/W)j = (R/W)1( – ). (13)
5 An Arrhenius equation of the form v = k0exp(–Ea/RT) cannot beintegrated analytically.
2900T
------------
v ri
v ri
i 1=
n
∑---------------
v r∂x∂τ------ v r
0∂ τ( )∂τ
-------------- v r0 1
τ------,= = =
v r0
j j 1–
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S177
A comparison of Eqs. (8) and (12) shows thatEq. (13) assumes a constant interaction surface area,which is equal to Sfr. This is true for a system of plane-parallel fractures, whereas other geometries of rockblocks (spherical jointing or hexagonal systems of frac-tures due to thermal contraction) result in a more rapidvr decrease with time than that predicted by Eqs. (12)and (13).
The limiting case is the exhaustion of fresh rockat a reactor step when R/W falls to zero. Equations(12) and (13) do not account for such a possibility, andany simulation algorithm must include a special con-straint on the total volume of rock that is allowed toreact with solution in the given reactor step, which fol-lows from the relation
(14)
where ncr is the number of the solution portion underwhich fresh rock is exhausted; Φ is the fracture poros-ity; and ρr and ρw are the densities of the rock and water,respectively. For ΣR/W values of 0.5–2.0 measured inoceanic hydrothermal systems (Von Damm, 1990) anda fracture porosity of basalt of about 3% (Hyndman andDrury, 1976), fresh rock is exhausted in the high-tem-perature reactor steps at a wave number of n × 1000,i.e., in long-lived model systems. This is in agreementwith observations in ophiolitic sequences (Richardsonet al., 1987; Nehlig et al., 1994), where the completealteration of rocks to epidosite occurs at the boundaryof dike complexes with isotropic gabbro marking theposition of the roof of a magma chamber (i.e., the hot-test and chemically most active part of the convectivesystem), whereas the overlying rocks are weaklyaltered.
Limitations of the model. The above-described pro-cedure of the simulation of a hydrothermal system as a
ncr R/W( ) j
j 1=
ncr
∑ R/W( )1 ncr1Φ----
ρr
ρw
------,≤=
polythermal step flow reactor involves implicit assump-tions which must be taken into account during the inter-pretation of computational results.
(1) The model assumes that the layer of metaso-matic rock is uniform, whereas it is a diffusion metaso-matic column according to the theory of diffusionmetasomatism. Ignoring this phenomenon impliesequal migration abilities for all the components that didnot form solid phases under the given ΣR/W values inthe model, whereas the natural situation will be morecomplicated with differentiated mobility of elements inthe direction normal to the fissure channels. Since themethod of determining ΣR/W from the compositions ofnatural solutions (Von Damm et al., 1985) is based onthe most mobile alkali elements, model calculations foroceanic hydrothermal systems may overestimate theremoval of some other elements from rocks (S, Pb, andZn); the model implicitly implies their removal up tothe frontal zone of the diffusion column, which may notcorrespond to the natural situation.
(2) The model suggests that fresh rock reacts withsolution as a homogenous material. This is imposed bythe use of Eq. (10), which was obtained for a nonequi-librium state solution–aluminosilicate interaction.However, as a result, the mineral composition of theinitial rock and, more importantly, the composition ofsolution do not affect the rate of this reaction.6
(3) The (R/W)1 value used in Eq. (13) concerns a dif-fusion-controlled process, and its distribution with tem-perature may not be identical to that defined byEq. (10), which was obtained for a process controlledby a surface reaction. The temperature dependency inthe Wood–Walther equation corresponds to an activa-
6 It should be noted that the rates of solution reactions with othermineral classes (oxides, sulfides, carbonates, and sulfates) arehigher than for aluminosilicates (Rafal’sky, 1993), and their inter-action with water will be controlled to a large extent by armoringwith an aluminosilicate matrix.
200150 250 300 350 400 450Temperature, °C
0
0.5
1.0
1.5
2.0
2.5ΣR
/W, kg/k
g1
2
3
4
(a) (b)
150100 200 250 300 350 400 450Temperature, °C
123
Fig. 9. Temperature dependencies of ΣR/W. (a) (1) Calculated by Eq. (11) in an approximation of continuous movement (normalizedto 1 kg/kg at 370°C); (2) calculated for a diffusion-controlled process at Ea = 16 kJ/g-at; (3) minimum and (4) maximum ΣR/Wvalues obtained in the hydrodynamic model of Tutubalin and Grichuk (1997a). (b) Dependencies of ΣR/W on T used in the thermo-dynamic models of (1) Grichuk et al. (1985), (2) Bowers and Taylor (1985), and (3) model calculations of this study.
S178
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
tion energy (Ea) of 10.5 kJ/g-at O for a heterogeneousreaction. This value is comparable to the activationenergies of ion diffusion in aqueous solutions (16–19 kJ/at; Erdey-Gru, 1974) and the activation energiesof diffusion-controlled reactions in an aqueous environ-ment (<20 kJ/mol; Lasaga, 1981).
The replacement of the constant in the integrand ofEq. (9) by the value characteristic of a diffusion processdoes not significantly change the R/W dependency(Fig. 9a). The deviations in the distribution of R/W willbe smaller than other model uncertainties, and account-ing for these features in Eq. (11) makes the model morecomplicated but does not improve its quality.
(4) The procedure accepted in the model for thedivision of the solution flow pipe into temperature seg-ments assumes a complete mixing within any segment,owing to hydrodynamic dispersion and the absence of
more considerable processes of solution mixing withdifferent thermal prehistories within the hydrothermalsystem. These conditions are not always satisfied innatural convective systems. While the former conditioncan be optimized by choice of the temperature incre-ment in the model (in fact, the size of the rock block isdistinguished as a reactor step), the second conditionrequires a hydrodynamic model of the process and thedivision of the thermodynamic model into a spatialmesh (Tutubalin and Grichuk, 1997b).
(5) Similar to the above-described isothermal col-umn (Fig. 7), the polythermal model based on theMSFR method is prone to artifacts, including the vari-ability in solution composition at a change in the com-position of mineral assemblages in previous reactorsteps.
The above-described method is used in this paperfor the simulation of convective hydrothermal systemsin the oceanic and island-arc crust (Section 4.1).
2.3. Modeling of Ore Deposition duringCooling of Hydrothermal Solutions
Modeling of ore deposition is of special interest forresearchers, because it is closely connected with prac-tical geological tasks. A considerable number of studieson the thermodynamic modeling of ore formation havebeen published up to now. They make use of variousapproaches, from the calculation of phase diagrams forore minerals and assemblages (Garrels and Christ,1965) to dynamic models (Averkin, 1987). Analysis ofall the proposed methods of modeling is not possible inthis contribution, and, similar to the previous section,we restrict ourselves to the models of the dynamic equi-librium class.
All the models that interpret cooling-related oredeposition in the context of the local equilibriumapproximation can be divided into two groups depend-ing on whether the system considered is open or closed.Open-system models can be further subdivided intotwo subgroups: without input of substances into thesystem (its composition can be changed by massremoval during spatial separation of solution and pre-cipitated phases) and with such addition (for instance,systems involving solution mixing). Modern methodsof thermodynamic calculations allow one to obtainrather precise (in the physicochemical sense) solutionsfor such models. As will be shown below, a major chal-lenge is the geological interpretation of model results,i.e., the determination of natural processes that can becorrelated with a particular model scenario. The princi-pal schemes of several scenarios considered in this sec-tion are shown in Fig. 10.
2.3.1. Scenario of closed-system cooling
In cooling scenarios a series of equilibrium states iscalculated for a system of given composition at decreas-ing temperature and pressure. The main property of the
(b)
(d)
(c)
T3
T2
T1
T2 � T1
T1
T3
T2
T1
Seawater
Fig. 10. Cartoons showing the construction of models forore deposition. (a) Model of slow cooling in a closed sys-tem; (b) model of rapid cooling; (c) model of slow coolingin a flow-through system; and (d) model of sequential cool-ing with mixing.
(a)
T ≠ const
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S179
closed-system scenario is that the state of the system ineach computation step is independent of the characterof previous equilibrium states of the system under dif-ferent equilibrium parameters.
The closest prototype of such a process is the calcu-lation of slow cooling of a gas–liquid inclusion in aninert crystal (Fig. 10a). If some phases become unstablein response to changes in conditions (for instance, atemperature decrease), they disappear from the equilib-rium assemblage. When a series of states is considered,it is equivalent to the occurrence of back-reactions, i.e.,replacement of previously formed minerals.
Figure 11a illustrates the calculation of the compo-sition of solid phases in a closed system of given com-position (Table 1, hydrothermal solution) as a functionof decreasing temperature. It can be clearly seen in thisdiagram that, when certain temperatures are reached,previously precipitated minerals are replaced (pyrrho-tite is replaced by pyrite at 175°C, and albite is replacedby pyrophyllite at 270°C). Since the system is closed, itcannot be regarded as a sequence of minerals depositedfrom a solution moving along a fracture or a channel.
Moreover, the correspondence of such a model to somereal situation is dubious, because the reactions ofrecrystallization of previously formed minerals areinhibited at low temperatures. In contrast to the equilib-rium model, the ore phases formed in nature can be“quenched” and will not be subsequently altered. Thus,the closed-system cooling scenario has limited applica-tions for the simulation of processes in the interiors ofhydrothermal systems.
The closed-system cooling scenario was used byBowers et al. (1985) for oceanic hydrothermal systems.Although, owing to the geochemical characteristics ofthe initial composition, the results of this study wererather reasonable, some details clearly show that theuse of a closed-system scenario is not justified for flow-through systems. For instance, the calculation of solu-tion composition in the Guaymas Basin hydrothermalsystem yielded an increase in Mn concentration owingto back-reactions during solution cooling at tempera-tures below 200°C.
Surprisingly, the closed-system cooling scenario hasfound application for simulation of the process of very
(a)
(b)
0.14
0.12
0.10
0.08
0.06
0.04
0.02
0350 320 290 260 230 200 125 50
Temperature, °C
Min
eral
, g/k
g s
olu
tion
Prl
Ab
Qtz
Ccp
Sp
Py
Qtz
Prl
Ab
200°C 150°C
Py
Po
Sp
Ccp
Qtz
Prl
100 mg/kgsolution
Prl
Py
Qtz
Po
Sph
Ccp
Prl
Py
Po
Sph
Ccp
Fig. 11. Calculation of ore deposition during cooling at a feeder reactor temperature of 350°C and P = 500 bar. (a) The mineralogyof cooling products in a closed system of a given composition as a function of temperature. (b) Rapid cooling in a closed system asa function of quenching temperature (quenching pressure of 200 bar). The circle area is proportional to the mass of precipitate.
Po
S180
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
fast hydrothermal and sedimentary ore deposition,when the temperature of a hydrothermal solutiondecreases abruptly during its venting on the ocean floor.If the ore phases are rapidly cooled and quenched andthey have no time to react with the components of bot-tom water, the closed-system model (Fig. 10b) yieldsthe desired assemblage. Such a scenario was used byGrichuk et al. (1985) for the calculation of ore deposi-tion in smokers.
The model of rapid cooling has its own stumblingblocks. It is necessary to define final quenching temper-ature, and this temperature is probably different for dif-ferent reactions. Some thermodynamically stable min-erals may be missing in similar natural associations,and metastable phases may occur (for instance, amor-phous silica instead of quartz). Thus, the model of rapidore deposition requires a priori natural information.
The average quenching temperature can be approx-imately estimated by calculations similar to thoseshown in Fig. 11a. The best agreement with smokeemitted by seafloor smokers, i.e., a suspension ofhydrothermal minerals formed during the rapid dis-charge of hot solutions into cold seawater, was attainedat a quenching temperature of 150–200°C. In order toemphasize the discontinuous character of solution cool-ing, pie diagrams of quench products appeared to beespecially convenient (Grichuk et al., 1985). These dia-grams allow presentation of additional dependenciesfrom the parameters of the model. Figure 11b shows thecomposition of hydrothermal precipitates as a functionof quenching temperature. Similar results are describedin Section 4.3.1.
2.3.2. Fractional precipitation in a flow-through system
If the natural prototype of the cooling model is aflow-through system, the principle of a flow reactormust be used in the model. Considering the available
computer programs for numerical simulation, itappears reasonable to use a step flow reactor instead ofa continuous one (Fig. 10c). If the reactor contains asufficiently large number of steps, the influence ofback-reactions will be small and the model results willbe quite adequate.
Figure 12 displays the simulation of solution cool-ing in a flow-through system. It can be seen that themajor portion of ore phases (pyrrhotite with minorsphalerite) is deposited within 300–200°C. Althoughthis diagram appears similar to Fig. 11a, the results aresignificantly different. The main ore phase during frac-tional precipitation is pyrrhotite rather than pyrite,which was obtained in the closed-system model(Table 2).
The temperature increment used in calculations isan indefinite parameter of the step flow reactor model.In practice, a decrease in the temperature incrementlengthens computations and makes the preparation ofinput data and the interpretation of model results morelaborious and time-consuming. The choice of incre-ment value is a compromise between computation timeand the quality of results. Table 2 compares the resultsof calculations obtained with different temperatureincrements. The similarity in the qualitative and quan-titative compositions of the bulk precipitate was used asa criterion. These results show that there are no real dif-ferences between the models with increments smallerthan 10°C. It was found that calculations with reducedtemperature steps are reasonable in the intervals of themost extensive ore precipitation. The calculation proce-dure can be optimized using a variable temperatureincrement. The results of such calculations are pre-sented in the next to last column of the table.
The last column of the table presents the results ofsimulation of closed-system cooling. A comparisonwith the models of step flow reactors reveals consider-able differences between the results of ore precipitationobtained in various model scenarios. This suggests that
0
0.002
0.004
0.006
0.008
0.010
Min
eral
s, g
/kg s
olu
tion
320350 290 260 230 200 125 50
Ab
Qtz
Ccp
Sp
Po
Py
Ab
Sph
Po
Qtz
Py
Temperature, °C
Fig. 12. Results of the simulation of fractional ore deposition during solution cooling in a flow-through system with a temperatureincrement of 10°C.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S181
it is inappropriate to use closed-system models for theproblems imitating flow-through systems. Among suchproblems is, for instance, the simulation of formation ofmineralized stockworks in the ascending channels of ahydrothermal system.
2.3.3. Scenarios of cooling with mixing
The mixing of solutions of different compositions isone of the main mechanisms of ore deposition in hydro-thermal systems. Since the temperatures of mixed solu-tions are in general different and the hotter solution isusually more metalliferous, temperature is also a factorof ore deposition during mixing.
The analysis of previously published results showsthat several scenarios have been invoked to simulatesuch processes. They differ in the style of introductionof the second (cold) solution and the way they accountfor metastable reactions.
The first thermodynamic model of ore depositionduring mixing was calculated by Janecky and Seyfried(1984). These authors addressed the problem of repro-ducing processes accompanying smoker discharge intobottom seawater. In contrast to the above-describedscenario of rapid discharge, they assumed that the sub-stances of hydrothermal solutions had enough time toreact chemically with seawater (completely or par-tially). Taking temperature as the main parameter of themodel and using heat balance conditions, these authorsobtained an equation for the calculation of mixing pro-portions of hydrothermal solution (mHS) and seawater(mSW). If the average heat capacities of these solutionsare not significantly different, this equation is written as
(15)mSW/mHS
THS Tmix–Tmix TSW–-------------------------,=
where Tmix is the temperature of the mixture. The exam-ination of their simulation method shows that it is amodification of the method of the degree of reactionprogress by Helgeson, in which the main component ofthe system is 1 kg of hydrothermal solution and coldseawater is gradually added to it. Thus, it is a peculiartitration procedure of hydrothermal solution withsimultaneous cooling. However, the Janecky–Seyfriedmodel does not involve the removal of solid reactionproducts. Therefore, some precipitated minerals couldbe dissolved and replaced by other phases during fol-lowing cooling steps. Figure 13 shows the results ofsimulation by this scenario for the mixing of smokersolution and cold seawater. It can be seen that, accord-ing to these calculations, newly formed anhydrite is notretained upon further cooling and talc is partly replacedby dolomite. Chalcopyrite precipitates at 350°C anddisappears at lower temperatures. Bornite forms and issubsequently replaced by chalcocite. At temperaturesbelow 100°C, the amounts of copper-bearing mineralsare the same as at the beginning of the process. Theconfinement of copper mineralization to the hottestparts of hydrothermal systems, which is among themost persistent features of oceanic ore formation, is notdescribed by such a scenario. This method cannot inprinciple reproduce spatial differentiation of ore com-ponents.
Since the nature of the aforementioned flaws in thisscenario is known, the mixing model can be easilymodified to be applied to the sequential mixing sce-nario (Fig. 10d). In this scenario, at each cooling stage,small batches of cold solution are added to the previ-ously formed mixture and the solid products of mixingare swept from the system. In essence, this scenario isa nonisothermal case of the Fritz method. The addition
Table 2. Simulation of the fractional deposition of ore matter during cooling of a hydrothermal solution from 350 to 50°C(mg/kg of solution); the composition of the initial solution is shown in Table 1
MineralCooling step
Closed system50°C 20°C 5°C optimized*
Pyrite 230.8 230.9 230.9 230.9 819.6
Pyrrhotite 601.9 601.7 601.5 601.5 0
Galena 0.327 0.327 0.327 0.327 0.327
Sphalerite 24.7 24.7 24.7 24.7 24.7
Chalcopyrite 1.25 1.25 1.25 1.25 1.25
Quartz 10582 10581 10581 10581 10588
Albite 6.6 6.8 7.1 7.1 0
Pyrophyllite 0.77 0.64 0.52 0.52 4.06
Actinolite-80 0.024 0.024 0.024 0.024 0
*The optimized increment is 10°C within 350–200°C and 25°C within 200–50°C.
S182
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
of cold solution is described by an equation identical toEq. (15). For the ith step of cooling, we have
(16)
where Δmi is the portion of seawater added to the solu-tion. Figure 14a shows the results of calculations for thescenario of sequential mixing of hydrothermal solutionand seawater. A comparison with the Janecky–Seyfriedmodel (Fig. 13) shows that the compositions of sedi-ments are different both quantitatively and qualitatively(in the association and temperature intervals of precip-itation).
One important merit of Janecky and Seyfried’s(1984) study was the evaluation of the role of metasta-ble states for such a rapid process as hydrothermal vent-ing. They demonstrated that ignoring the reaction ofseawater sulfate with reducers occurring in hydrother-mal solution provides better agreement with naturalobservations. This approach appeared to be very prom-ising and was subsequently used by many authors. Inaddition to sulfur compounds, the metastable characterof ore deposition processes is evidently reflected insome other hydrothermal components. In particular,evidence from oceanic smokers suggests that the con-centrations of silica in their solutions correspond toequilibrium with amorphous silica rather than withquartz (Wells and Ghiorso, 1991). The equilibriumdynamic approach provides no criteria for the distin-guishing of metastable constituents of the system, andthis can be done only by comparison with natural pro-totypes.
Figure 14b shows the results of simulation for ametastable case of the sequential mixing scenario. Itdiffers from the equilibrium case in the association ofnewly formed minerals and ranges of their precipita-
Δmi mi 1–
Ti 1– Ti–Ti TSW–---------------------,=
tion. In particular, the order of precipitation changesfrom talc–pyrite–anhydrite in the equilibrium case toanhydrite–talc–pyrite in the metastable case; there is alow-temperature deposition of magnesium carbonates,brucite, etc. The geochemical aspects of this problemare discussed in more detail in Section 4.3.1.
2.4. Conclusions
The scenarios described in this chapter do not obvi-ously include all aspects of the simulation of hydrother-mal processes. Natural situations are certainly muchmore diverse. For instance, scenarios of processesinvolving a gas phase were not considered in this chapter(some examples of such models are given in Chapter 6).As the techniques of geochemical modeling arebeing developed, more sophisticated methods canalso appear for those interactions that were consideredin this chapter.
Significant progress has been made in the develop-ment of combined scenarios for processes includingvarious successive elementary interactions from thoseconsidered above. A spectacular example of such com-bined scenarios is the model of self-mixing developedfor the formation of hydrothermal lode ores (Barsukovand Borisov, 1982a, 1982b; Borisov, 2003). This modelreproduced the generation of pore solutions in a hydro-thermal metasomatic halo at the expense of hydrother-mal solution interaction with rocks (using the MSFRmethod) and ore deposition due to mixing in the areaswhere solutions percolating through the rock arefocused into a fissure conduit. Another example ofcombined scenarios is given in Section 4.3.2 with appli-cation to the simulation of formation of a large sulfideedifice on the seafloor.
Fig. 13. Calculation of the mixing of hydrothermal solution with cold seawater using the Janecky–Seyfried model. Here, after theoccurrence of minor quantities of ore minerals is shown with arrows.
350 300 250 200 150 100 50
Temperature, °CBnSphCcp Cc
0
0.5
1.0
1.5
2.0
2.5
Min
eral
s, g
/kg s
olu
tion
Anh
Tlc
Py Dol
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S183
The MSFR method has found intriguing applicationin the study by Borisov (2003) on vein mineralizationrelated to conductive cooling. He considered two sce-narios referred to as the reaction and layered models.The reaction model describes a prolonged processinvolving metasomatic alteration of previously depos-ited vein material by the solution of subsequent por-tions. The layered model inhibits metasomatic replace-ments, and newly formed sediments joined those ofprevious stages. A comparison with natural prototypes(deposits of the Sadon ore region) showed that the lay-ered model is in better agreement with nature.
One of the most promising directions in the devel-opment of equilibrium dynamic models for hydrother-mal systems is a synthesis of hydrodynamic and ther-modynamic models. It is reasonable to expect that thiswork will give rise to a new class of models allowing acorrelation of chemical interactions with space andtime coordinates, which is of special interest for geol-ogy. Some pilot studies in this direction have alreadybeen published (Steefel and Lasaga, 1994; Tivey, 1995;Tutubalin and Grichuk, 1997a, 1997b).
With the advent of diverse modeling techniques,there is a need to choose the best method for the solu-
tion of particular problems. The analysis performed inthis chapter suggests that errors of two types may asso-ciate with this choice.
Errors of the first type arise in the schematizationstage (Section 1.1), when the logical scheme con-structed on the basis of natural data appears to be inad-equate to the problem formulated. This situation can beexemplified by the choice of a stable scenario for oredeposition during mixing instead of metastable one(Sections 2.3.3 and 4.3.1). First-type errors can berevealed by comparing the results of simulation withthe natural prototype.
Errors of the second type originate in the internalinconsistency of the model that appears during the tran-sition from a logical scheme to a physicochemical andmathematical model. Among such errors is the applica-tion of the method of the degree of reaction progress toflow-through systems. In many cases such defects ofthe method can be unraveled by the logical analysis ofthe model construction (Fritz, 1975). A comparisonwith natural prototypes does not usually permit recog-nition of second-type errors. The only reliable way todistinguish them is a comparison with an independentalternative model that reproduces, at least qualitatively,
350 320 290 230 200 125 50
(a)
Temperature, °C
Min
eral
, g/k
g s
olu
tion
(b)
0.25
0.20
0.15
0.10
0.05
0
Min
eral
, g/k
g s
olu
tion
0.5
0.4
0.3
0.2
0.1
0
Tlc
Anh
Bn
Ccp
Cc
Sp
Py
Brc
Srp
Tlc
Anh
Mgs
Dol
Ccp
Cc
Sp
Py
Anh
Py
Tlc
Anh
Tlc
Py
Ccp Sph Bn Cc
Srp
Mgs
Brc
Dol
CcSphCcp
Fig. 14. Results of calculations for the scenario of sequential mixing. (a) Equilibrium mixing; (b) metastable mixing with “frozen”reactions of sulfate reduction, methane oxidation, and quartz precipitation.
260
S184
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
the logical scheme of the process. An example of suchan analysis is given in Section 2.1, where the qualitativesolution of the problem of formation of infiltrationmetasomatic zoning obtained in the Korzhinskii theorywas used to control quantitative simulation methods.
CHAPTER 3. THERMODYNAMIC MODEL OF A CONVECTIVE HYDROTHERMAL SYSTEM
IN A MID-OCEAN RIDGE
3.1. Geologic Setting and Compositionof High-Temperature Oceanic Hydrothermal Systems
The discovery of high-temperature hydrothermalactivity on the ocean floor is one of the most significantbreakthroughs in geology in recent decades. Advancesin the understanding of oceanic hydrothermal systemsare related to studies by many researchers from variouscountries. The development of innovative techniques inoceanology and marine geology, the advent of mannedsubmersibles operating at great depths, played a crucialrole in this process. The use of submersibles, side-look-ing sonar systems, deep-tow cameras, samplers, andhigh-precision navigation systems enabled detailedinvestigation of such in accessible and small-sizedobjects as submarine hydrothermal vents and relatedsulfide edifices.
The investigation of oceanic hydrothermal systemsinvolves a great diversity of techniques and methodsand rival in this respect any studies of modern geology.This resulted in considerable progress in the under-standing of the nature of these amazing objects. Thehistory of the discovery and investigation of oceanichydrothermal systems was described in the publica-tions of leaders and members of research teams (Rona,1984; Lisitsin et al., 1990; Hydrothermal Sulfide…,1992; Bogdanov, 1997). This section does not attemptto give a chronologically ordered record of ideas andthe results of investigations.
The goal of this section is to provide a synopsis ofmodern knowledge on the geologic structure of hydro-thermal systems in the oceanic crust, the chemical andmineral compositions of products of this activity (oresand solutions), the genetic schemes of this process, andtheir experimental and theoretical substantiation; somecontroversial issues are analyzed in more detail.
3.1.1. Geologic setting
There are a number of compilations on the hydro-thermal activity in the world ocean (Rona, 1984;Hydrothermal Sulfide…, 1992). The results obtained inrecent years were reviewed by Rona and Scott (1993)and Bogdanov (1997). Figure 15 shows the locations ofthe high-temperature objects whose geochemistry wasbest characterized. The majority of currently knownsubmarine hydrothermal systems occur in the modernspreading zones of mid-ocean ridges. High-tempera-ture hydrothermal systems were found at spreading
centers in the back-arc basins of the southwesternPacific (Manus, Woodlark, Lau, and North Fiji basins),island-arc troughs (Mariana and Okinawa troughs),hotspot-related submarine volcanoes (Loihi, Teahitia,and Macdonald). A group of similar objects is repre-sented by submarine hydrothermal vents in the coastalregions of New Zealand, New Guinea, Iceland, theKuril Islands, Italy, and Japan. New discoveries arecontinuous. The geologic setting and properties ofhydrothermal systems have been most comprehen-sively studied in mid-ocean ridges. This group of sys-tems was used as a natural prototype for the construc-tion of a thermodynamic model.
All the known active hydrothermal systems of thisgroup are located in the axial valleys of mid-oceanridges on a very young crust. They are usually foundeither in the axial (neovolcanic) zone or at the base ofvalley-wall faults (Fouquet et al., 1988, 1996; Lisitsinet al., 1990; etc.). Such a correlation with the morpho-logically distinct elements of the geologic structureconsiderably facilitates the search for hydrothermalsystems. Hydrothermal systems are unevenly distrib-uted along mid-oceanic ridges. Within fast-spreadingridge segments bounded by transform faults, activehydrothermal systems are most frequently observed intopographic highs in the central parts of the segmentsand are less common at their ends (Ballard andFrancheteau, 1982; Crane, 1985; Macdonald and Fox,1988). Such a distribution is currently interpreted asreflecting a greater magma chamber thickness and ahotter crust beneath the central part of the segment.
This pattern of hydrothermal system distribution isnot universal. The Logachev hydrothermal field, whichwas found on the slow-spreading Mid-Atlantic Ridge at15° N (Akimtsev et al., 1991), is situated in an upliftedblock of ultrabasic rocks near a transform fault. Severalother systems related to blocks of ultrabasic rocks weresubsequently found on the Mid-Atlantic Ridge. TheAxial Seamount vent field on the Juan de Fuca Ridge isconfined to the caldera of a submarine volcano sittingdirectly on the ridge axis. Evidence for ancient hydro-thermal activity was detected at several off-axis volca-noes at 13° and 21° N on the East Pacific Rise (EPR)(Alt et al., 1987; Fouquet et al., 1988; Hekinian et al.,1989; Hydrothermal Sulfide…, 1992).
A particular group of hydrothermal systems withspecific geologic structures, mineralogy, and chemicalcompositions develops in the segments where mid-ocean ridges are overlain by thick sedimentary covers.Three such systems have been found up to now: in theGuaymas Basin in the Gulf of California, EscanabaTrough, and Middle Valley (Fig. 15). Their peculiaritiesare related to the interaction of hydrothermal solutionswith sediments, and they are not considered in thispaper.
The hydrothermal systems of back-arc basins do notform a compact group. Some of them (North Fiji,Woodlark, and Manus basins) are related to back-arc
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S185
80°
0°
90°
180°
90°
60° 60°
0°
60°
60°
0°
60° 8
0°
0°
90°
180°
90°
60°
21
20
19
18
17
16 15 14
22
29
–33
26,
27
25
28
24
23
34
10
9 11
12
13
2 1 34
5
67
8
Fig
. 15.
Sch
emat
ic d
istr
ibut
ion
of g
eoch
emic
ally
cha
ract
eriz
ed s
ubm
arin
e hy
drot
herm
al s
yste
ms.
Num
eral
s co
rres
pond
to th
ose
in T
able
6.
S186
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
spreading centers in a basaltic crust and are geochemi-cally similar to the systems of mid-ocean ridges. Inter-mediate and silicic igneous rocks occur at the hydro-thermal sites of the Lau and eastern Manus basins (Fou-quet et al., 1993; Binns and Scott, 1993). The JADEhydrothermal field occurring in a submarine caldera inthe Okinawa Trough is considered as a modern ana-logue of Kuroko-type deposits (Halbach et al., 1989,1993; Marumo and Hattori, 1999; Glasby and Notsu,2003), and its geochemical characteristics are signifi-cantly different from those of mid-ocean ridge systems.
The energy of mid-ocean ridge hydrothermal sys-tems is usually supplied by magma bodies. Early stud-ies supposed the existence of linear magma chambers,about 20 km wide, beneath ridge axes. They wereregarded as sources of considerable heat flows and gen-eration of a new oceanic crust. This view evolved grad-ually as new comprehensive geophysical data becameavailable. The existence of wide magma chambers wasnot confirmed, instead, they appeared to be much nar-rower. Moreover, there is often no evidence for thepresence of magma bodies under slow-spreading ridges(i.e., either there is no magma chamber or it is narrowand not detectable by geophysical methods). It is cur-rently thought that magma chambers beneath fast-spreading ridges are no wider than 2–4 km and pinchout toward transform faults (Phinney and Odom, 1983;McClain et al., 1985; Macdonald and Fox, 1988; Wil-cock and Delaney, 1996). The depth of chamber roofsis determined from seismoacoustic data andmicroseism records by bottom seismographs (Riedeselet al., 1982) and is usually 1–3 km. Beneath slow-spreading ridges, magma chambers are even narrowerand probably exist sporadically.
The structure of the oceanic crust was reconstructedon the basis of investigations of ophiolitic complexesand is supported by the results of deep-sea drilling. Theupper layer is made up of pillow lavas produced by sub-marine eruptions. The underlying dike complex isformed through filling of extensional fractures, whichserved as crustal conduits for submarine eruptions. Thedike complex rests on a layer of isotropic gabbro, whichis interpreted as the result of melt freezing on the roofof magma chambers. A layered intrusive complexbeneath it was formed at the bottom of a magma cham-ber during the crystallization differentiation of melt.The intrusive complex is underlain by ultrabasic uppermantle rocks. Hydrothermal alteration occurs in therocks of the upper section including the isotropic gab-bro layer, and the degree of alteration increases down-ward (Coleman, 1977; Gillis and Robinson, 1990;Anderson et al., 1990; Gillis, 1995).
The igneous rocks of the oceanic crust show fairlystable petrochemical characteristics. They are classifiedas tholeiitic basalts and are collectively referred to asmid-ocean ridge basalts (MORB). The differentiationtrend of these basalts shows an enrichment of residualmelts in iron and titanium. Their extensive differentia-
tion in the oceanic crust occasionally produces fer-roandesites and ferrorhyolites, which is exemplified byigneous complexes of the Galapagos Spreading Center(e.g., Perfit et al., 1983; Embley et al., 1988). The anal-ysis of such occurrences suggests that Fe–Ti basalts aswell as ferroandesites and ferrorhyolites can be formedin small residual magma chambers before the completesolidification of magma bodies. The EPR and MARsegments hosting active hydrothermal systems arecomposed of relatively weakly differentiated basalts(Stakes et al., 1984; Davis and Clague, 1987; Hekinianand Walker, 1987).
The compositions of newly formed mineral associa-tions in the country rocks near hydrothermal systemswere studied by many authors, reviewed by Mottl(1983), Silantyev (1984), and Kurnosov (1986). In thebeginning of these studies, Humphris and Thompson(1978) pointed out the existence of two types of basaltalterations characterized by the occurrence of chloriteand epidote, respectively. The mineral association ofthe former type includes chlorite, mixed-layer chlo-rite/smectite, quartz, occasionally hematite and anhy-drite. The association of the second type includes epi-dote, chlorite, albite, actinolite, and minor quartz. Therecharge area of a convective hydrothermal system(downwelling limb of convection) is characterized bythe magnesian and ferromagnesian compositions ofchlorite. It was shown that the mineral association cor-relates with temperature: epidotized rocks are usuallyrelated to higher interaction temperatures, and chlori-tized rocks correspond to moderate temperatures. Forlow-temperature conditions, Kurnosov (1986) distin-guished zeolite and smectite metamorphic facies, butthese processes are characterized by very low rates andare probably related to halmyrolysis rather than tohydrothermal events. Strictly speaking, the chlorite andepidote assemblages of secondary minerals are metaso-matic rather than metamorphic, because the composi-tions of altered rocks are significantly different fromthose of initial basalts, which suggests a substantial lossor gain of major elements (Humphris and Thompson,1978). Both associations are typical of metasomaticallyaltered basic rocks and correspond to the chlorite andpropylite facies, respectively (Plyusnina, 1983).
There is still limited evidence on the character ofmetasomatic alteration in the upwelling limb of a con-vection cell for oceanic hydrothermal systems. The rel-evant data were mainly obtained by dredging extinctsystems where feeder channels had been exposed bytectonic processes. In the TAG hydrothermal field, afootwall zone of metasomatic alteration was drilledduring Leg 158 of the Ocean Drilling Program(Humphris et al., 1995; Herzig et al., 1998). The zoneof metasomatic alteration near the feeder conduit ofoceanic hydrothermal systems is no wider than a fewhundred meters (Jonasson et al., 1986; Ridley et al.,1994). According to available observations, basaltsundergo greenschist alteration and propylitization inthis zone, producing a characteristic assemblage of
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S187
chlorite with epidote and albite (Mottl, 1983; Bideauet al., 1985; Embley et al., 1988). A zone of paragonitedevelopment was documented in the TAG field, in theupper part of a channel immediately below a sulfidebody (Humphris et al., 1998). In this respect, oceanichydrothermal systems are different from ancient mas-sive sulfide deposits, where quartz–sericite metaso-matic rocks passing to chlorite rocks are commonlyformed around the channel (Franklin et al., 1981; Meta-somatism and…, 1998).
3.1.2. Compositions and properties of sulfide edifices and hydrothermal solutions
Mid-ocean ridges are significantly different withrespect to the mode of occurrence of ore-forming pro-cess. The fast-spreading EPR (up to 18 mm/y) com-prises hydrothermal fields with numerous small hydro-thermal vents (up to 130 in the segment between12°41′ N and 12°52′ N; Hydrothermal Sulfide…,1992), most of which are already inactive. The massesof individual bodies are n × 103–n × 104 t. The ridgesspreading at a moderate rate comprise fields with asmaller number of larger edifices. Only several verylarge hydrothermal mounds were found in the TAG andMARK fields on the slow-spreading Mid-AtlanticRidge (spreading rate <2 cm/y) and the GalapagosSpreading Center. Their masses are estimated from theapparent sizes as several million tons (Bogdanov, 1997;Herzig et al., 1998). Isotopic investigations suggest thatthe characteristic growth time of a small sulfide bodyon the EPR is n × 102–n × 103 y, whereas the large struc-tures of the TAG and MARK fields experienced a longand complex evolutionary history (Lalou et al., 1993,1995).
The morphology of sulfide edifices was comprehen-sively described by Lisitsin et al. (1990) and Bogdanov(1997). Small edifices are usually steep-sided cones,several meters high (occasionally, up to 20 m), with thechimney of an active hydrothermal vent on top. Combi-nations of several conical edifices are common. Thebasements of the edifices are built up of clastic materialcomposed of cemented products of destruction and oxi-dation of pipelike chimneys. Large hydrothermal edi-fices are gently sloping conical mounds, up to 60 mhigh and up to 250 m and even more in diameter. Thetops of such edifices are often truncated and host groupsof active hydrothermal vents (Lisitsin et al., 1990,Figs. 21–23, 25). Small hydrothermal edifices oftenshow rather complicated shapes.
Of special interest are the sizes of sulfide bodiesoccurring in the sedimentary cover. Deep-sea drilling inthe Middle Valley region (Davis et al., 1992) confirmedthe suggestion by Lisitsin et al. (1990) on the extensionof such bodies into sedimentary layers. During Leg 139of the Ocean Drilling Program, an ore body was pene-trated to a depth of 94 m.
The mineral composition of oceanic hydrothermaledifices was discussed in a number of publications (seeHydrothermal Sulfide…, 1992 for a review). There area limited number of major minerals accounting formore than 10% of the rock volume: the ore constituentsare pyrite, marcasite, chalcopyrite, sphalerite, andwurtzite; and the gangue minerals are opal, anhydrite,and barite. Isocubanite and talc were occasionallyfound in significant amounts. Sulfide edifices develop-ing on a sedimentary sequence (Guaymas Basin,Escanaba Trough, and Middle Valley) show peculiarmineral compositions with considerable amounts ofpyrrhotite; the Middle Valley assemblage also includesmagnetite (Krasnov et al., 1994; Krasnov andStepanova, 1996). Some hydrothermal systems ofback-arc basins and island arcs are enriched in chalco-phile trace elements (Pb, Ag, As, Sb, etc.), which resultsin the formation of a number of silver (Mozgova et al.,1993) and arsenic minerals (Halbach et al., 1989,1993). Table 3 lists all the minerals detected in oceanichydrothermal edifices.
The mineralogy of hydrothermal sulfide bodies israther diverse, and various authors distinguished differ-ent types of sulfide ores on the basis of mineral compo-sition (up to 10 types or even more; Fouquet et al.,1988). The simplest classification was proposed inHydrothermal Sulfide… (1992), where the two mostwidespread mineral assemblages were distinguishedamong the objects studied: (a) pyrite–marcasite–sphalerite ores with opal, characteristic of small sulfidebodies and outer parts of large edifices, which are oftenhighly porous; and (b) massive pyrite and pyrite–chal-copyrite ores with quartz, which are probably morecommon in the inner parts of large sulfide mounds.Their mineralogical characteristics have been studiedin detail only in the outer parts of hydrothermal sulfideedifices, whereas the interior structure and mineralogi-cal zoning of large bodies remain poorly known.7 Prob-able schemes of mineralogical zoning in the sulfidebodies and their evolution in time are discussed in Sec-tion 3.2.
The chemical composition of oceanic sulfide ores ingeneral mirrors their mineralogy. Based on the propor-tions of the major ore elements Fe, Cu, and Zn, Krasnov(Hydrothermal Sulfide…, 1992) distinguished fourgeochemical types of ores: (1) massive pyrite (Cu <4.5% and Zn < 1%), (2) copper sulfide (Cu > 4.5% andZn < 1%), (3) zinc sulfide (Cu < 4.5% and Zn > 1%),
7 Up to now, drilling operations have been performed at fourhydrothermal locations. Drilling in the Snake Pit hydrothermalmound at 23° N on the MAR (Detrick et al., 1986) resulted invery low recovery. In 1994 drilling operations were carried out onan active edifice of the TAG field (Humphris et al., 1995; Herziget al., 1998), and its geochemical structure was characterized indetail. A sulfide body at Middle Valley (Davis et al., 1992; Kras-nov and Stepanova, 1996) showed a peculiar composition, whichis related to its occurrence in the sedimentary sequence. Drillingduring ODP Leg 193 was carried out at the Pacmanus hydrother-mal vent field, which developed in dacitic rocks, but only somepreliminary results have been reported (Paulick et al., 2004).
S188
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
and (4) mixed copper–zinc sulfide (Cu > 4.5% and Zn >1%). The data set studied by Krasnov (262 samples)included 46% of type 3 samples, 24% of type 2 sam-ples, 15% of type 1 samples, and 15% of type 4 sam-ples. The chemical compositions of these ore types areshown in Table 4. Copper and zinc are negatively cor-related in most objects and show strong local variationswithin individual hydrothermal edifices and even theirsegments.
The relative proportions of metals in the ores arecorrelated with the size and tectonic setting of bodies.Small edifices in the axial zones of ridges with high andmedium spreading rates are strongly enriched in zinc.Larger bodies from the same setting show higher ironcontents. The largest sulfide structures from the MAR,axial volcanoes of the EPR, and the Galapagos rift
show the maximum iron content and copper enrichmentrelative to zinc, mainly at the expense of zinc depletion(Krasnov, 1990; Hydrothermal Sulfide…, 1992). Newdata by Hannington et al. (1998) supported these rela-tionships.
The minor elements of the ores show ubiquitouscorrelations with major components. The pyrite andcopper sulfide types are relatively rich in Co, Se, andMo (high-temperature association), whereas zinc sul-fide ores have higher Pb, Cd, As, Sb, Ag, and Au con-centrations (medium-temperature association) (Han-nington et al., 1991; Hydrothermal Sulfide…, 1992).The investigations by Krasnov and his colleagues onthe chemical compositions of individual mineralsshowed that zinc sulfides are the main carriers of Cd,Ni, and Ag; pyrite, of As and Co; and chalcopyrite, ofSe. As, Sb, Ag, and Au form their own mineral phasesin oceanic deposits (Table 3).
High-temperature venting into bottom oceanicwater is the most remarkable manifestation of hydro-thermal processes in the oceanic crust. Mixing of thedischarged solution with cold seawater almost immedi-ately produces a suspension of mineral particles(smoke). That is why the hydrothermal vents werenamed smokers (RISE Project…, 1980). Smoke col-umns rise a few hundred meters above the seafloor,merging with each other and dispersing gradually.Based on the color of the smoke, which depends on itsmineral composition, black smokers (pyrrhotite andother sulfides) and white smokers (amorphous silicaand anhydrite) were recognized. Gray and even trans-parent smokers were also described (Craig et al., 1987;Kastner et al., 1987; Lisitsin et al., 1992). The characterof a smoker appeared to be closely related to the tem-perature of the hydrothermal system. Black smokersare the hottest and show an issue temperature of upto 350–360°C, whereas the temperature of whitesmokers is no higher than 330°C and usually 200–300°C. In addition to focused discharge through chim-neys, large ore edifices may show diffuse venting,which was observed during submersible dives as watershimmering.
The maximum measured temperatures of hydrother-mal systems are remarkably uniform. Most of thehydrothermal systems of the EPR and the Juan de FucaRidge have temperatures no higher than 350–355°C.Temperatures up to 360–365°C were determined inthree most extensively studied systems of the Mid-Atlantic Ridge. Noteworthy is a correlation betweenthese values and ocean depth. Hydrothermal systemsare discharged at depths of 3600–3100 m in the AtlanticOcean and 2500–2600 m in the EPR. Note that the rel-atively shallow hydrothermal system of Axial Sea-mount on the Juan de Fuca Ridge (1900 m) shows amaximum temperature of 328°C (Butterfield et al.,1990).
The existence of hydrothermal solutions with tem-peratures higher than 350–360°C has been discussed
Table 3. Minerals found in the sulfide hydrothermal edifices onthe ocean floor, mainly after Hydrothermal Sulfide… (1992)with some additional data from other sources (Shadlun et al.,1992; Bogdanov, 1997; Mozgova et al., 1998; Lein et al., 2003)
* Minerals identified in ores from the hydrothermal fields of the Oki-nawa Trough (Halbach et al., 1989, 1993; Glasby and Notsu, 2003).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S189
extensively in the literature (Delaney et al., 1989; Tiveyet al., 1990) in connection with the problem of boilingin oceanic hydrothermal systems (Chapter 6). By 1989four reliable temperature measurements between 375and 405°C had been documented, including one vent ofthe Endeavor hydrothermal field, where a temperaturemonitoring experiment was carried out over 46 days(Tivey et al., 1990). Similar experiments of longerduration have been conducted at 9°50′ N on the EPR(Fornari et al., 1998), in the Cleft segment of the Juande Fuca Ridge (Tivey et al., 2002), and in the activeTAG edifice after drilling (Goto et al., 2002). Thesemeasurements demonstrated that oceanic hydrothermalvents are characterized by prolonged periods of steadytemperatures of 350–365°C. They may be disturbed byvolcanic and seismic events resulting in short-termtemperature variations. The maximum temperature of405°C was measured at 21°33′ S on the EPR (VonDamm et al., 2003) (see also Chapter 6).
The nature of the temperature uniformity of hydro-thermal systems remains elusive. The reason for theconstancy of temperature could be related to the prop-erties of water rather than to the internal structure ofhydrothermal systems. According to the hypothesis ofCann et al. (1985/1986), this phenomenon is caused bythe heat removal efficiency, which passes through a
maximum near the aforementioned temperatures owingto concurrent changes in the heat capacity and viscosityof water.
Many researchers attempted to estimate dischargerates and heat fluxes for individual hydrothermal ventsand whole hydrothermal fields. The discharge rates ofhot springs were estimated from the visual determina-tions of flow velocity and chimney diameter, but theaccuracy of such estimates is not high. There is a signif-icant discrepancy between the data of various authors,which is probably related to the methods of determina-tion and the considerable natural variability of dis-charge rate values (Table 5). In large hydrothermalstructures, solutions are usually emitted throughnumerous vents forming groups of smokers (Lisitsinet al., 1990), and the total discharge rate of such groupsmay be 1–2 orders of magnitude higher than that ofindividual smokers. Another complex problem is theassessment of relative discharge rates of localized anddiffuse venting in large edifices. According to Ronaet al. (1993), the total heat flux for an active edificefrom the TAG field appeared to be 5–10 times higherthan that through the group of smokers on top of theedifice. This suggests a prevalence of diffuse dischargein large sulfide bodies.
Table 4. Average chemical compositions of oceanic sulfide ores (Hydrothermal Sulfide…, 1992)
Element, % Massive ironsulfide type
Massive coppersulfide type
Massive zincsulfide type Copper–zinc type Average for all ores
Fe 37.40 34.28 22.90 24.44 27.60
Cu 0.26 10.42 0.45 6.08 1.36
Zn 0.26 0.21 12.33 4.32 2.23
S 38.59 37.98 31.84 29.87 32.0
Mg – 0.10 0.12 0.02 0.09
Al 0.26 0.34 0.08 0.14 0.15
Si 3.72 1.97 3.61 2.82 2.83
Ca 2.39 2.38 0.32 0.45 0.30
Ba 0.35 0.047 0.23 0.30 0.21
Mn 0.11 0.014 0.035 0.016 0.07
Pb 0.0608 0.0168 0.1188 0.0495 0.0659
ppm
Co 325 448 20 49 66
Ni 54 28 19 47 31
As 782 102 525 252 348
Se 226 118 13 25 29
Mo 90 251 37 98 62
Cd 8.6 17.4 287 250 101
Sb 81 3.5 80 21 37
Ag 30 15 100 73 57
Au 0.093 0.49 1.44 1.8 1.0
S190
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
The chemical composition of oceanic hydrothermalsolutions has been a subject of many studies. Early pub-lications were reviewed by Von Damm (1990). Takinginto account recent data, it is now possible to character-ize to a varying extent the geochemistry of more than30 hydrothermal systems (Fig. 15). Tables 6–9 summa-rize data on the major and trace component composi-tions of solutions, their isotopic parameters, and dis-solved gases.
The main technical problem in the geochemicalinvestigation of hydrothermal solutions is sampling ofalmost point-sized targets at great depths. This problemwas solved only by the introduction of submersible div-ing and special samplers. However, even these sam-pling methods could not exclude bottom water. Evenunder the most favorable conditions, the fraction of sea-water in a sample is several percent and, more fre-quently, several tens of percent. Edmond et al. (1982)proposed a method to circumvent this difficulty: aseries of samples is collected in a hydrothermal vent,and a line of mixing with seawater is constructed foreach component using solution analyses. Edmond pro-posed use of the magnesium concentration as a mixingparameter, because its concentration is very low in thehydrothermal vents of mid-ocean ridges. Compositionsextrapolated to the concentration Mg = 0 using the mix-ing diagrams are interpreted as pure hydrothermal solu-tions (end-member). The validity of the assumption onzero magnesium content was confirmed using othercomponents (SO4, U, and Mo). Small negative valuesfor the SO4 concentration occasionally obtained bysuch an extrapolation are probably related to the precip-itation of anhydrite and barite during the periodbetween sampling and analysis.
Unfortunately, the method proposed by Edmondet al. (1982) cannot be used to process the data obtainedfrom vents in ultrabasic blocks (Lein et al., 2000). Theconcentration of Mg in water from these ventsdecreases less rapidly than in fluids emanating from thebasaltic crust. There is still no reliable method ofextrapolation for such solutions.
The analysis of chalcophile elements poses a specialproblem, because they may precipitate owing to cool-ing during sampling. These losses must be taken intoaccount (Von Damm, 1990; Trefry et al., 1994). Fur-thermore, rapidly moving hydrothermal solutions cantransport mechanical suspensions from the feeder chan-nel and the sulfide body, and it is not possible to reliablydiscriminate them from the suspension formed inresponse to sample cooling.
Despite considerable variations in numerical values,the oceanic solutions display stable compositionaltrends relative to seawater. These trends were estab-lished in the earliest studies (Edmond et al., 1982;Michard et al., 1984; Von Damm et al., 1985), and thesubsequent work only provided further detail. Hydro-thermal solutions differ from seawater in having muchlower concentrations of Mg, SO4, U, and Mo. They aresignificantly enriched in the major components K, Ca,and Si and the lithophile trace elements Li, Rb, Cs, Be,and Al. Their Sr content may be both higher and lowerthan that of seawater. Hydrothermal solutions are prob-ably enriched in Ba and Ra (which is suggested by thedata on 226Ra and its daughter isotopes 222Rn and 210Pb;Kadko and Moore, 1988), but these elements are exten-sively precipitated during mixing with seawater in thecourse of sampling. The hydrothermal systems occur-ring in areas with thick sedimentary covers producesolutions enriched in NH4 and I, and, to a smallerextent, in Br and B.
Hot springs from oceanic hydrothermal systems arecharacterized by high concentrations of ore elements(Fe, Mn, and Zn) and a number of chalcophile trace ele-ments (Cu, Pb, Cd, Co, Ag, As, Sb, and Se), which areenriched by 3–7 orders of magnitude relative to seawa-ter. These elements (especially trace elements) show aconsiderable scatter, and the observed variations arecorrelated mainly with vent temperatures. The concen-trations shown in Tables 6 and 7 represent the highesttemperature vents of hydrothermal fields and probablymost adequately characterize the solutions ascending tothe seafloor surface.
Table 5. Discharge rates and heat flows of hydrothermal systems
Hydrothermal system No. in Fig. 15
Discharge rateof solution, kg/s Heat flow, MW
Referencesinglesource
wholesystem single source whole system
21° N EPR 1 – – 6 ± 2 × 107 cal/s – MacDonald et al., 1980
MAR, TAG field 16 – – 225 ± 25 (active mound) Rona et al., 1993
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S191
Tabl
e 6.
Che
mic
al c
ompo
sitio
ns o
f hi
gh-t
empe
ratu
re s
olut
ions
fro
m o
cean
ic h
ydro
ther
mal
sys
tem
s ex
trap
olat
ed to
the
end-
mem
ber
with
Mg
= 0
Reg
ion
Seaw
ater
Eas
t Pac
ific
Ris
e
Hyd
roth
erm
al s
yste
m21
° N
Gua
ymas
Bas
in13
° N11
° N9°
17′ N
, Ven
t “F”
No.
in F
ig. 1
51
23
45
Ven
tN
GS
OB
SSW
HG
34
13
419
9119
94
No.
12
34
56
78
910
1112
com
pone
nt,
para
met
erun
it
t°C
227
335
035
535
128
531
531
7–
(380
)34
738
835
1
pH(2
5°C
)7.
83.
83.
43.
63.
35.
95.
93.
23.
23.
33.
12.
82.
6
Cl
mm
ol/k
g54
157
948
949
649
663
759
974
071
876
056
346
.584
6
Na
"46
451
043
243
944
351
348
556
058
759
647
238
.468
3
K"
9.8
25.8
23.2
23.2
23.9
37.1
40.1
29.6
29.8
28.8
32.0
1.16
41.5
Ca
"10
.220
.815
.616
.611
.741
.534
.055
.044
.654
.822
.51.
8345
.6
SiO
2"
0.16
19.5
19.5
17.3
15.6
13.5
13.8
22.0
21.9
17.9
18.8
5.93
20
Alk
mg-
eq/k
g2.
3–
0.19
–0.4
0–0
.30
–0.5
06.
508.
10–
–0.
638
–0.
398
–1.0
21–1
.4–1
.8
CO
2m
mol
/kg
2.3
5.72
––
––
–10
.8–1
6.7
––
–15
.511
.4
H2S
"0
6.6
7.3
7.5
8.4
5.2
4.8
–2.
94.
58.
041
8.71
SO4
"27
.90.
000.
500.
600.
40–0
.34
0.06
––
––
–1.0
40.
86
Fe"
<10
–60.
871
1.66
40.
750
2.42
90.
180
0.07
71.
450
3.98
10.7
66.
471.
4912
.1
Mn
"<
10–6
1.00
20.
960
0.69
90.
878
0.23
60.
139
1.00
01.
689
2.03
50.
766
0.17
43.
28
Zn
μmol
/kg
0.01
4010
689
104
4019
–10
22
105
––
Cu
"0.
007
<0.
0235
9.7
44<
0.02
1.10
––
––
––
Pb"
1 ×
10–5
0.18
30.
308
0.19
40.
359
0.65
20.
230
–0.
135
0.01
40.
050
––
Ref
eren
ceV
on D
amm
, 199
0M
icha
rdet
al.,
198
4B
ower
s et
al.,
198
8V
on D
amm
et a
l., 1
997
S192
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 6
. (C
ontd
.)
Reg
ion
Eas
t Pac
ific
Ris
eJu
an d
e Fu
ca R
idge
Hyd
roth
erm
al s
yste
m17
°25′
S18
°15′
SR
apa
Nui
Ven
t Fie
ld,
21°3
4′ S
, Bra
ndon
ven
tE
xplo
rer
Rid
geE
ndea
vour
seg
men
t, 47
°57′
N
No.
in F
ig. 1
56
78
910
Ven
tN
adir
Ako
rta
brin
eva
por
Ven
t 12d
Hul
kL
obo
Pean
utN
orth
No.
1314
1516
1718
1920
21co
mpo
nent
,pa
ram
eter
unit
t°C
340
>30
537
640
127
635
334
635
035
6
pH(2
5°C
)3
3.25
3.26
3.1
4.62
4.5
4.25
–4.
5
Cl
mm
ol/k
g19
084
855
729
755
950
542
825
347
7
Na
"12
568
644
124
2–
391
336
216
378
K"
6.7
20.3
13.8
6.87
31.1
327
.624
.513
.529
.1
Ca
"5.
247
32.8
15.8
22.3
742
.933
.013
.834
.3
SiO
2"
10.6
16.8
12.5
8.69
4.40
17.0
15.9
11.0
16.8
Alk
mg-
eq/k
g–
––
0.68
5–
0.81
50.
68–
0.06
–0.
06–1
.10.
192
CO
2m
mol
/kg
13.1
8.5
––
–8.
411
.722
.022
.0
H2S
"8.
68.
46.
867.
930.
964
2.9
5.0
8.1
2.1
SO4
"–
–(–
0.5)
2.63
12.7
9–
––
–
Fe"
0.59
12.2
12.3
6.97
5.50
00.
533
0.64
6–
0.17
7
Mn
"0.
251.
731.
30.
622
0.15
50.
194
0.26
30.
325
0.27
1
Zn
μmol
/kg
9031
412
110
00.
077
3228
––
Cu
"10
8810
545
0.01
59
14–
–
Pb"
––
––
––
0.14
5–
–
Ref
eren
ceC
harl
ou e
t al.,
199
6V
on D
amm
et a
l., 2
003
Tun
nicl
iffe
et a
l., 1
986*
1B
utte
rfie
ld e
t al.,
199
4
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S193
Tab
le 6
. (C
ontd
.)
Reg
ion
Juan
de
Fuca
Rid
ge
Hyd
roth
erm
al s
yste
mA
xial
Vol
cano
, ASH
ES
fiel
dSo
uthe
rn J
uan
de F
uca
Rid
ge, C
left
seg
men
t, 45
° NE
scan
aba
Tro
ugh,
41°
N
No.
in F
ig. 1
511
1213
Ven
tIn
fern
oH
ell
Cra
ckV
irgi
nM
ound
Plum
eV
ent 1
Ven
t 3Pi
pe O
rgan
Mon
olith
No.
2223
2425
2627
2829
3031
com
pone
nt,
para
met
erun
it
t°C
328
301
217
299
224
285
–26
232
721
7
pH(2
5°C
)3.
53.
5–
4.4
––
–2.
82.
85.
4
Cl
mm
ol/k
g62
455
025
817
610
8789
695
112
4590
866
8
Na
"49
944
620
914
879
666
178
469
569
556
0
K"
26.8
24.6
11.5
6.98
51.6
37.3
45.6
58.7
41.0
40.4
Ca
"46
.838
.918
.910
.296
.484
.777
.310
979
.033
.4
SiO
2"
15.1
––
13.5
23.3
22.8
22.7
24.0
21.0
6.91
Alk
mg-
eq/k
g–
0.48
–0.
510.
40.
66–
––
–2.0
5–
0.82
3.1
CO
2m
mol
/kg
50(9
0)17
928
54.
464.
34–
––
–
H2S
"7.
1–
–(1
8)3.
52.
694.
41.
993.
71.
1
SO4
"–
––
–0.
50–1
.30
–1.7
0–
4.14
–0.
68–
Fe"
1.06
50.
868
0.01
330.
012
18.7
3910
.349
17.7
7016
.44.
90.
01
Mn
"1.
150
1.13
60.
287
0.14
23.
585
2.61
14.
480
4.25
1.94
0.01
–0.
021
Zn
μmol
/kg
111
134
2.6
2.2
780
374
469
520
580
<0.
1
Cu
"9.
91.
20.
10.
41.
515
.52
1.4
7.8
<0.
1
Pb"
0.30
2–
–0.
101
1.63
0.73
10.
981.
090
–
Ref
eren
ceB
utte
rfie
ld e
t al.,
199
0V
on D
amm
, 199
0; E
vans
et a
l., 1
988;
Tre
fry
et a
l., 1
994;
Met
z an
d T
refr
y, 2
000
But
terf
ield
and
Mas
soth
, 19
94; T
refr
y et
al.,
199
4Z
iere
nber
get
al.,
199
3
S194
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 6
. (C
ontd
.)
Reg
ion
Mid
–Atla
ntic
Rid
ge
Hyd
roth
erm
al s
yste
mL
ogat
chev
,14
°45′
NM
AR
K, 2
3° N
TA
G, 2
6° N
Bro
ken
Spur
, 29°
NL
ost C
ity,
30° N
Rai
nbow
,36
°14′
NL
ucky
Str
ike,
37°
17′ N
Men
ez
Gw
en,
37°5
0′ N
No.
in F
ig. 1
514
1516
1718
1920
21
Ven
t1
Snak
e Pi
tB
lack
smok
ers
Whi
tesm
oker
s4
2608
Ven
tY
3E
iffe
lM
enez
flan
k
No.
3233
3435
3637
3839
4041
4243
com
pone
nt,
para
met
erun
it
t°C
353
350
345
366
273–
320
364
45–7
036
532
832
432
428
4
pH(2
5°C
)3.
33.
93.
83.
353
–9–
9.8
2.8
3.78
3.7
3.7
4.2
Cl
mm
ol/k
g51
555
955
963
6–
469
544–
551
750
526
472
417
313
Na
"43
851
054
655
7–
419
465–
512
553
406
402
347
381
K"
2223
.624
17.1
17.1
18.8
–20
27.4
24.8
21.6
23.0
Ca
"28
9.9
1030
.827
12.8
19.4
–24.
167
42.1
36.7
32.3
33.1
SiO
2"
8.2
18.2
1820
.75
19.1
––
6.9
17.5
15.4
13.3
10.3
Alk
mg-
eq/k
g–
–0.
064
–0.
563
–0.
45–1
.1–
––
–0.
58–
––
CO
2m
mol
/kg
10.1
5.2–
6.7
–2.
9–3.
4–
––
16–
28.4
23.0
20.1
H2S
"0.
85.
92.
72.
5–6.
70.
58.
5–11
.00.
064
1.0–
1.2
4.64
3.0
2.1
1.7
SO4
"–
––
––
6.0–
7.1
5.9–
12.9
–0.
327
13–2
8–
–
Fe"
2.5
2.18
02.
121
5.59
03.
830
2.15
60.
0008
24.0
0.74
90.
863
0.62
40.
0282
Mn
"0.
330.
491
0.44
30.
680
0.75
00.
260
2.25
0.38
90.
446
0.28
90.
068
Zn
μmol
/kg
25–3
050
4746
300–
400
880.
2–6.
916
0–
3726
4.2
Cu
"15
–50
1712
120–
150
368
.60.
2–0.
714
0–
1516
.52.
7
Pb"
0.08
60.
265
–0.
110
–0.
376
0.07
–0.
120.
148
–0.
130
–0.
056
Ref
eren
ceD
ouvi
lleet
al.,
200
2;C
harl
ouet
al.,
200
2
Cam
pbel
let
al.,
198
8;D
ouvi
lleet
al.,
200
2
Jean
–Bap
-tis
te e
t al.,
19
91
Edm
ond
et a
l., 1
995;
Dou
ville
et a
l., 2
002;
Cha
rlou
et a
l., 2
002
Jam
eset
al.,
199
5;C
harl
ouet
al.,
200
2
Kel
ley
et a
l.,20
01*2 ;
Lei
n et
al.,
2004
*3
Dou
ville
et a
l., 2
002;
Cha
rlou
et a
l., 2
002
Von
Dam
met
al.,
199
8C
harl
ou e
t al.,
200
0, 2
002;
Dou
ville
et a
l., 2
002
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S195
Tab
le 6
. (C
ontd
.)
Reg
ion
Indi
an O
cean
Bac
k–ar
c ba
sins
of
the
Paci
fic
Hyd
roth
erm
al s
yste
mR
odri
gues
Tri
ple
Junc
tion,
Kai
rei F
ield
Lau
Bas
in,
Vai
Lili
fie
ldN
. Fiji
Bas
inM
anus
bas
inE
ast M
anus
bas
in,
Des
mos
fie
ldE
ast M
anus
bas
in,
Pacm
anus
fie
ldW
oodl
ark
Bas
in,
Fran
klin
Sea
mou
nt
No.
in F
ig. 1
522
2324
2526
2728
Ven
tV
L–3
Whi
te L
ady
Vie
nna
Woo
ds
No.
4445
4647
4849
50co
mpo
nent
,pa
ram
eter
unit
t°C
360
334
285
275
88–1
2026
8(2
70–3
50)
pH(2
5°C
)3.
52
4.7
4.5
2.08
2.6
–
Cl
mm
ol/k
g64
279
025
071
248
657
262
6–73
7
Na
"56
059
021
053
442
044
5–
K"
14.3
79–
2412
.186
–
Ca
"30
41.3
6.55
827.
615
.154
SiO
2"
15.8
14.5
1515
5.8
16.2
10.6
–15.
1
Alk
mg-
eq/k
g–
0.46
––
–0.
11–
9.2
––
CO
2m
mol
/kg
–(7
.6–1
5.6)
–6
1532
–
H2S
"4.
0(<
0.2)
21.
810
6.8
–
SO4
"0.
3–
––
32.8
––
Fe"
5.4
2.50
~12
0.10
90.
012.
404
1.9–
2.8
Mn
"0.
847.
1–
0.34
80.
113.
116
0.5–
0.7
Zn
μmol
/kg
–30
00–
1012
115
19–2
7
Cu
"–
34–
<2
<2
35.6
39–5
5
Pb"
–3.
9–
<0.
004
0.02
27.
0858
–82
Ref
eren
ceG
amo
et a
l., 2
001
Fouq
uet
et a
l., 1
993
Gri
mau
det
al.,
198
9L
isits
in e
t al.,
199
2;B
ach
et a
l., 2
003
Gam
o et
al.,
199
7;B
ach
et a
l., 2
003
*4B
ach
et a
l., 2
003
Lis
itsin
et a
l., 1
991
S196
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 6
. (C
ontd
.)
Reg
ion
Isla
nd a
rcs
and
volc
anoe
s of
the
Paci
fic
Hyd
roth
erm
al s
yste
mO
kina
wa
Tro
ugh,
Min
ami–
Ens
ei K
noll
Oki
naw
a T
roug
h,C
LA
M f
ield
Oki
naw
a T
roug
h,JA
DE
fie
ldO
kina
wa
Tro
ugh,
Ihey
a N
orth
Kno
llO
kina
wa
Tro
ugh,
Suiy
o Se
amou
ntH
awai
ian
Isla
nds,
Loi
hi V
olca
no
No.
in F
ig. 1
529
3031
3233
34
Ven
tPe
le’s
ven
ts
No.
5152
5354
5556
com
pone
nt,
para
met
erun
it
t°C
265–
278
100–
216
320
238
296–
311
31
pH(2
5°C
)4.
9–5.
15.
2–5.
64.
724.
63.
7–3.
95.
6
Cl
mm
ol/k
g50
1–52
546
2–51
455
044
1–45
840
1–10
0052
2–53
8
Na
"41
0–43
1–
425
377–
385
445
451–
463
K"
49–5
1–
7249
–50
3011
.5–1
3.2
Ca
"21
–22
–22
14–1
589
10.9
–11.
7
SiO
2"
10–1
1–
138
131.
77
Alk
mg-
eq/k
g3.
0–3.
52.
5–10
.31.
90.
5–3.
6–
0.20
8.12
–11.
1
CO
2m
mol
/kg
64–9
674
–86
26–2
00–
3930
0
H2S
"1.
6–2.
49.
2713
.7–
1.3
2.1
SO4
"–
16.4
––
–26
.0–2
6.5
Fe"
––
––
–0.
603–
1.46
Mn
"0.
088–
0.09
4–
0.11
0.44
–0.
447
0.58
60.
0206
–0.
0484
Zn
μmol
/kg
––
––
––
Cu
"–
––
––
–
Pb"
––
––
––
Ref
eren
ceG
lasb
y, N
otsu
, 200
3G
amo
et a
l., 1
991*
5 ;G
lasb
y an
d N
otsu
, 200
3G
amo
et a
l., 1
991;
Gla
sby
and
Not
su, 2
003
Gla
sby
and
Not
su, 2
003
Sedw
ick
et a
l., 1
992*
6
Dat
a no
t ext
rapo
late
d to
the
end-
mem
ber:
*1 M
g =
22.
37; *
2 Mg
= 9
–19;
*3 M
g =
11–
18.8
; *4 M
g =
49;
*5 M
g =
34.
3; a
nd *
6 Mg
= 4
9.5–
51.1
mm
ol/k
g.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S197
Tabl
e 7.
Tra
ce-e
lem
ent c
ompo
sitio
ns o
f hi
gh-t
empe
ratu
re s
olut
ions
fro
m o
cean
ic h
ydro
ther
mal
sys
tem
s ex
trap
olat
ed to
the
end-
mem
ber
with
Mg
= 0
Reg
ion
Seaw
ater
Eas
t Pac
ific
Ris
e
Hyd
roth
erm
al s
yste
m21
° N
Gua
ymas
Bas
in13
° N11
° N9°
17′ N
, Ven
t “F”
No.
in F
ig. 1
51
23
45
Ven
tN
GS
OB
SSW
HG
34
13
419
9119
94
No.
in T
able
61
23
45
67
89
1011
12co
mpo
nent
unit
Li
μmol
/kg
2610
3389
189
913
2272
087
368
861
416
884
1816
20
Rb
"1.
331
.028
.027
.033
.057
.066
.014
.118
–24
––
Cs
nmol
/kg
223
020
223
024
6–
––
115
591
195
––
Be
"0.
0237
1510
1342
29–
–20
––
–
Srμm
ol/k
g87
9781
8365
253
226
175
171
131
802.
3216
7
Ba
"0.
14>
16>
8>
10>
11>
15>
54–
––
––
–
As
nmol
/kg
2730
247
214
452
1071
1074
––
––
––
Se"
2.5
<0.
672
7061
3810
3–
––
––
–
Br
μmol
/kg
840
929
802
877
855
–10
6311
63–
––
74.8
1350
B"
416
507
505
500
548
–15
70–
––
––
–
Al
"0.
024.
05.
24.
74.
56.
73.
7–
––
––
–
Co
nmol
/kg
0.03
2221
366
227
<5
<5
––
––
––
Cd
"1
1715
514
418
046
27–
55–
30–
–
Ag
"0.
02<
138
2637
242
––
––
––
Iμm
ol/k
g–
––
––
–58
––
65–
––
NH
4"
<10
––
––
1030
012
900
––
168
––
–
Ref
eren
ceV
on D
amm
, 19
90V
on D
amm
, 199
0;
Palm
er a
nd E
dmon
d,19
89V
on D
amm
, 199
0;C
ampb
ell a
nd
Edm
ond,
198
9
Mic
hard
et a
l., 1
984
Bow
ers
et a
l., 1
988;
Palm
er a
nd E
dmon
d, 1
989
Von
Dam
met
al.,
199
7
S198
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 7
. (C
ontd
.)
Reg
ion
Eas
t Pac
ific
Ris
eJu
an d
e Fu
ca R
idge
Hyd
roth
erm
al s
yste
m17
°25′
S18
°15′
SR
apa
Nui
Ven
t Fie
ld,
21°3
4′ S
, Bra
ndon
ven
tE
xplo
rer
Rid
geE
ndea
vour
seg
men
t, 47
°57′
N
No.
in F
ig. 1
56
78
910
Ven
tN
adir
Ako
rta
Bri
neV
apor
Ven
t 12d
Hul
kL
obo
Pean
utN
orth
No.
in T
able
613
1415
1617
1819
2021
com
pone
ntun
it
Li
μmol
/kg
183
690
488
270
472
439
379
160
438
Rb
"2
6.8
––
32.9
38.0
33.6
––
Cs
nmol
/kg
––
––
–36
433
1–
–
Be
"–
––
––
––
––
Srμm
ol/k
g13
.318
793
.547
.212
015
313
560
149
Ba
"6.
117
.4–
––
––
––
As
nmol
/kg
––
––
––
––
–
Se"
––
––
––
––
–
Br
μmol
/kg
318
1320
890
490
––
––
–
B"
––
465
429
–69
875
572
372
5
Al
"–
––
––
––
––
Co
nmol
/kg
––
––
––
––
–
Cd
"–
––
––
––
––
Ag
"–
––
––
––
––
NH
4μm
ol/k
g–
––
–18
.250
056
357
7–
Ref
eren
ceC
harl
ou e
t al.,
199
6V
on D
amm
et a
l., 2
003
Tun
nicl
iffe
et a
l., 1
986*
1B
utte
rfie
ld e
t al.,
199
4
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S199
Tab
le 7
. (C
ontd
.)
Reg
ion
Juan
de
Fuca
Rid
ge
Hyd
roth
erm
al s
yste
mA
xial
Vol
cano
, ASH
ES
fiel
dSo
uthe
rn J
uan
de F
uca
Rid
ge, C
left
seg
men
t, 45
° NE
scan
aba
Tro
ugh,
41°
N
No.
in F
ig. 1
511
1213
Ven
tIn
fern
oH
ell
Cra
ckV
irgi
nM
ound
Plum
eV
ent1
Ven
t3Pi
pe O
rgan
Mon
olith
No.
in T
able
622
2324
2526
2728
2930
31co
mpo
nent
unit
Li
μmol
/kg
636
548
286
184
1718
1108
1808
2350
1570
1268
Rb
"–
––
–37
.028
.032
.038
.527
.695
–105
Cs
nmol
/kg
––
––
––
–24
821
771
00–7
700
Be
"–
––
–95
150
150
––
–
Srμm
ol/k
g19
2–
8146
312
230
267
348
236
209
Ba
"26
––
6–
––
890
–
As
nmol
/kg
––
––
(650
)(2
60)
–33
0(3
70)
–
Se"
––
––
––
––
––
Br
μmol
/kg
956
856
401
250
1832
1580
1422
1880
1295
1179
B"
590
––
450
496
––
524
482
1900
–216
0
Al
"–
––
––
––
––
–
Co
nmol
/kg
––
––
450
200
(180
)14
3057
–
Cd
"–
––
–47
091
0(3
75)
550
700
–
Ag
"–
––
–48
120
–40
88–
Mo
"–
––
–31
6–
–30
–
Sb"
––
––
1118
–7
15–
Iμm
ol/k
g–
––
––
––
––
99
NH
4"
––
––
––
––
–56
00
Tl
nmol
/kg
––
––
(92)
(45)
(60)
74(4
2)11
0
Ref
eren
ceB
utte
rfie
ld e
t al.,
199
0V
on D
amm
,199
0; T
refr
y et
al.,
1994
; Met
z an
d T
refr
y, 2
000
But
terf
ield
, Mas
soth
,19
94; T
refr
y et
al.,
199
4;M
etz
and
Tre
fry,
200
0
Zie
renb
erg
et a
l., 1
993
S200
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 7
. (C
ontd
.)
Reg
ion
Mid
–Atla
ntic
Rid
ge
Hyd
roth
erm
al s
yste
mL
ogat
chev
, 14
°45′
NM
AR
K, 2
3° N
TA
G, 2
6° N
Bro
ken
Spur
, 29°
NL
ost C
ity,
30° N
Rai
nbow
,36
°14′
NL
ucky
Str
ike,
37°1
7′ N
Men
ezG
wen
,37
°50'
′ N
No.
in F
ig. 1
514
1516
1718
1920
21
Ven
t1
Snak
e Pi
tB
lack
smok
ers
Whi
tesm
oker
s4
Y3
Eif
fel
Men
ez
flan
k
No.
in T
able
632
3334
3536
3738
3941
4243
com
pone
ntun
it
Li
μmol
/kg
245
843
1030
411
–10
35–
340
357
286
274
Rb
"28
10.0
11.9
9.1
9.4
13.6
–37
39.1
22.7
25.6
Cs
nmol
/kg
385
177
–10
811
314
8–
333
235
167
330
Be
"–
38.5
––
––
––
––
–
Srμm
ol/k
g13
850
5010
391
42.9
–20
011
976
111
Ba
">
4.5
>4.
3–
>19
–>
14.9
–>
679.
442
.422
.6
As
nmol
/kg
––
(70)
––
––
Br
μmol
/kg
818
847
–88
0–10
45–
749
–11
7892
273
571
0
B"
518
–35
638
847
3–
––
––
Al
"4
5.3
––
––
–2
11–
6
Co
nmol
/kg
––
–(2
570)
–42
214
0–16
013
000
––
Cd
"63
––
66–
145
–13
079
–12
Ag
"11
––
51–
––
4725
–17
Mo
"1
––
5–
––
284
––
Sb"
––
–3.
9–
––
3.1
6.3
–4.
7
Tl
nmol
/kg
7–
–13
––
–9
16–
11
Ni
"–
––
––
–64
–126
030
00–
––
Ref
eren
ceD
ouvi
lleet
al.,
200
2;
Cha
rlou
et a
l., 2
002
Cam
pbel
let
al.,
198
8;
Cha
rlou
et a
l., 2
002
Jean
–Bap
-tis
te e
t al.,
1991
Von
Dam
m, 1
990;
Edm
ond
et a
l., 1
995;
Met
z an
d T
refr
y, 2
000;
Dou
ville
et a
l., 2
002;
Cha
rlou
et a
l., 2
002
Jam
es e
t al.,
1995
Lei
n et
al.,
2004
*2D
ouvi
lleet
al.,
200
2;C
harl
ouet
al.,
200
2
Cha
rlou
et a
l., 2
000;
Dou
ville
et a
l., 2
002
Cha
rlou
et a
l., 2
000
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S201
Tab
le 7
. (C
ontd
.)
Reg
ion
Bac
k-ar
c ba
sins
of
the
Paci
ficIs
land
arc
s an
d vo
lcan
oes
of th
e Pa
cific
Hyd
roth
erm
al s
yste
mL
au B
asin
,V
ai L
ili f
ield
Man
us b
asin
Eas
tM
anus
bas
in,
Des
mos
fie
ld
Eas
tM
anus
bas
in,
Pacm
anus
fi
eld
Oki
naw
aT
roug
h,M
inam
i–E
nsei
Kno
ll
Oki
naw
aT
roug
h,JA
DE
fie
ld
Oki
naw
aT
roug
h, I
heya
Nor
th K
noll
Oki
naw
aT
roug
h,Su
iyo
Seam
ount
Haw
aiia
nIs
land
s,L
oihi
Vol
cano
No.
in F
ig. 1
523
2526
2729
3132
3334
Ven
tV
L–3
Vie
nna
Woo
dsPe
le’s
ven
ts
No.
in T
able
645
4748
4951
5354
5556
com
pone
ntun
it
Li
μmol
/kg
623
1010
3672
454
00–5
800
2500
800
–46
.1–5
7.6
Rb
"68
113.
575
––
––
3.19
–4.8
1
Cs
nmol
/kg
1560
––
––
––
––
Srμm
ol/k
g12
029
061
9721
5–22
794
56–6
230
489
–92.
4
Ba
">
3917
0.9
353
–56
5910
–12
100
0.61
–0.8
8
As
nmol
/kg
6–
––
––
––
–
Br
μmol
/kg
1140
1070
750
911
––
––
–
B"
830
––
–37
00–4
000
3400
––
–
Al
"–
1410
9.4
––
––
–
Co
nmol
/kg
–5.
8–
––
––
––
Cd
"15
008.
9–
––
––
––
Ag
"–
6.5
––
––
––
–
Au
"–
7.1
––
––
––
–
Bi
"–
0.19
––
––
––
–
Fμm
ol/k
g–
3113
533
7–
––
––
NH
4"
––
––
4600
–470
020
0–49
0019
00–2
600
<10
02.
88–7
.01
Ni
nmol
/kg
––
––
––
––
39–8
0
Ref
eren
ceFo
uque
tet
al.,
199
3L
isits
in e
t al.,
1992
; Bac
het
al.,
200
3
Bac
h et
al.,
2003
*3B
ach
et a
l.,20
03G
lasb
y, N
otsu
, 200
3–Se
dwic
ket
al.,
199
2*4
Not
e: D
ata
not e
xtra
pola
ted
to th
e en
d-m
embe
r: *
1 Mg
= 2
2.37
; *2 M
g =
11.
0–18
.8; *
3 Mg
= 4
9; a
nd *
4 Mg
= 4
9.5–
51.1
mm
ol/k
g.
S202
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
The majority of hydrothermal solutions are weaklyacidic (pH varies within 3–5 at 25°C). More acidicsolutions with pH lower than 2.5 were found in hydro-thermal sites in the Lau and East Manus back-arc basins(Table 6). Hydrothermal systems developing in the sed-imentary cover usually produce solutions with pHhigher than 5. Owing to the acidic pH values of hydro-thermal solutions, the measurement of their alkalinity(Alk) usually yields negative values. All the high-tem-perature hydrothermal solutions of the ocean containconsiderable amounts of hydrogen sulfide, from 1 to12 mmol/kg (except for the Vai Lili vent field in the Lauback-arc basin, where H2S was not reliably detected;Fouquet et al., 1993).
A special problem is the behavior of Cl and Br in anumber of hydrothermal systems. The chloride ion isone of the most conservative components of seawater,and no process in the oceanic water mass can signifi-cantly change its concentration. Elevated chlorinitywas first discovered in a hydrothermal system at 13° Non the EPR (Michard et al., 1984) and was activelydebated in the literature. Michard et al. (1984) analyzedpossible reasons for Cl accumulation including(a) water binding during hydration of crustal rocks,(b) extraction from basalts, (c) input of concentratedmagma-derived fluids, and (d) boiling of hydrothermalsolution in crustal levels. These authors concluded thatsolution boiling was the most plausible mechanism.Subsequent studies (Von Damm and Bischoff, 1987;Edmonds and Edmond, 1995; Edmond et al., 1995)strengthened this conclusion by additional evidence,and this view is now universally accepted (this problemis considered in more detail in Chapter 6). The mostconvincing argument for the occurrence of fluid phaseseparation in the interiors of hydrothermal systems isthe finding of vents with both high and low mineraliza-tion. The latter are products of vapor condensation.Such vents were found in hydrothermal systems atAxial Seamount and the Endeavor segment of the Juande Fuca Ridge; in the North Fiji back-arc system; at9°17′ N and 21°33′ S on the EPR; and Broken Spur,Lucky Strike, and Menez Gwen fields along the MAR(Table 6).
The concentrations of Br and Cl are strongly corre-lated and the Br/Cl ratio of hydrothermal solutions var-ies within a narrow range of (1.44–1.68) × 10–3 inde-pendent of chlorinity (Campbell and Edmond, 1989),which is similar to that of seawater (1.55 × 10–3). Thesefeatures are in agreement with the boiling mechanism.Br/Cl values as low as (0.6–0.9) × 10–3 were found dur-ing the investigation of very low-salinity high-tempera-ture vents at 9°49′ N on the EPR (Oosting and VonDamm, 1996). The experimental test of existinghypotheses demonstrated that these variations are dueto halite precipitation at almost complete boiling off ofwater and its subsequent dissolution under changing P–Tconditions (Berndt and Seyfried, 1997).
The hydrogen and oxygen isotope composition ofwater was determined in several hydrothermal systems(Table 8). In all the objects studied, it was very similarto the isotopic composition of seawater, which is indic-ative of a marine water source in the oceanic hydrother-mal systems (Welhan and Craig, 1983). According tocalculations, a small positive shift of 1–2‰ relative tothe local bottom water could be related to interactionwith basalts (Bowers and Taylor, 1985). Strontium iso-tope ratios were determined in many hydrothermal sys-tem (Table 8) and lie between the isotopic characteris-tics of modern basalts (0.70280) and seawater(0.70917), which reflects strontium input from thesetwo sources. The Pb isotope ratios of hydrothermal sys-tems from 21° N EPR and the southern Juan de FucaRidge are identical to those of the host basalts, whichare therefore regarded as a source of this element. Amore radiogenic lead composition in solutions from theGuaymas Basin suggests Pb input from a sedimentarysequence (Chen et al., 1986; Chen, 1987). The sulfurisotope composition of hydrogen sulfide was deter-mined in solutions from more than a dozen hydrother-mal systems (Table 8). The values obtained range from+0.5 to +7.8‰ reflecting a mixing of sulfur from mag-matic (leached from basalts) and marine (reduction ofsulfate sulfur of seawater by iron from basalts) sources(Bluth and Ohmoto, 1988). This topic is addressed inmore detail in Chapter 5.
The data on dissolved gases show that the oceanichydrothermal solutions are strongly enriched in hydro-gen and methane, whose concentrations are very low innormal seawater (Table 9). The accumulation of 3Hewas detected in hydrothermal solutions (Table 8): the3He/4He ratios are higher than the atmospheric value(RA) by a factor of 7–8 and approach the isotopic com-position of helium from modern mid-ocean ridgebasalts. According to the scarce measurements, the δDof dissolved hydrogen and methane corresponds to thetemperature conditions within hydrothermal systems(Table 8). The abundances of dissolved gases arestrongly affected by phase separation associated withboiling. The boiled off solutions show a high mineral-ization and low dissolved gas content (Evans et al.,1988), whereas low-concentration hydrothermal solu-tions are usually enriched in gases, especially CO2 andH2S (Table 6).
Summarizing the available data on the chemicalcompositions of ore-forming solutions of oceanichydrothermal systems, it should be pointed out that S,Fe, and Mn usually prevail over other ore components,and that most of the systems show elevated concentra-tions of Zn and minor Cu and Pb. With respect to S, Fe,and Zn content, these solutions correspond to recon-structions of ore-forming solutions of massive sulfidedeposits (Kappel and Franklin, 1988). The data pre-sented in Tables 6 and 7 indicate that the concentrationsof ore elements in solutions from different systems aregenerally similar but not identical and may differ bymore than an order of magnitude. Since ore element
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S203
Tabl
e 8.
Iso
topi
c ch
arac
teri
stic
s of
oce
anic
hyd
roth
erm
al s
olut
ions
Reg
ion
Seaw
ater
Eas
t Pac
ific
Ris
e
Hyd
roth
erm
al s
yste
m21
° N
Gua
ymas
Bas
in13
° N11
° N
No.
in F
ig. 1
51
23
4
Ven
tN
BS
OB
SSW
HG
34
13
4
No.
in T
able
61
23
45
67
89
10pa
ram
eter
unit
Srμm
ol/k
g87
9781
8365
253
226
175
188
189
86
87Sr
/86Sr
0.70
910.
7030
0.70
310.
7033
0.70
300.
7052
0.70
520.
7041
0.70
410
0.70
405
0.70
424
δ34S(
H2S
)‰
(C
D)
–3.
41.
3–1.
52.
7–5.
52.
3–3.
2–
––
––
–
δD(H
2O)
‰ (
SMO
W)
02.
5–
––
––
0.62
…1.
49–
––
δ18O
(H2O
)‰
(SM
OW
)0
1.6–
2.0
––
––
–0.
39…
0.66
––
–
δ11B
‰
39.5
32.7
32.2
31.5
30.0
–23
.2–
––
–
δ13C
(CH
4)‰
(PD
B)
––1
5.0
–17.
4–1
7.2
––
––1
9.5…
–16.
6–
––
δD(C
H4)
‰ (
SMO
W)
––1
02; –
126
––
––
––
––
–
δD(H
2)‰
(SM
OW
)–
–383
; –39
6–
401
–373
––
––
––
–
δ3 He/
4 He
RA
7.8
––
––
–7.
537.
86–
8.34
δ13C
(CO
2)‰
(PD
B)
0–7
.0–
––
––
–5.5
…–
4.1
––
–
Ref
eren
ceV
on
Dam
m,
1990
Von
Dam
m, 1
990;
Wel
han
and
Cra
ig, 1
983
Von
Dam
m, 1
990
Mic
hard
et a
l.,19
84; M
erliv
atet
al.,
198
7
Palm
er a
nd E
dmon
d, 1
989;
Kim
et a
l., 1
984
S204
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 8
. (C
ontd
.)
Reg
ion
Eas
t Pac
ific
Ris
eJu
an d
e Fu
ca R
idge
Hyd
roth
erm
al s
yste
m9°
17′ N
,V
ent “
F”17
°25′
S18
°15′
SR
apa
Nui
Ven
tFi
eld,
21°
34′ S
,B
rand
on v
ent
Mid
dle
Val
ley
Axi
al V
olca
no,
ASH
ES
fiel
dC
left
seg
men
t, 45
° NE
scan
aba
Tro
ugh,
41°
N
No.
in F
ig. 1
55
67
811
1213
Ven
t19
9119
94N
adir
Ako
rta
Bri
neV
apor
Infe
rno
Vir
gin
Mou
ndPl
ume
Ven
t1V
ent3
No.
in T
able
611
1213
1415
1622
2526
2728
31pa
ram
eter
unit
Srμm
ol/k
g–
–13
.342
.893
.547
.2–
––
336
–27
420
9
87Sr
/86Sr
––
––
––
–0.
7037
0–
0.70
358
0.70
988
δ34S(
H2S
)‰
(C
D)
––
––
4.2
4.8
––
–4.
2… 7.3
4.0… 6.4
7.2… 7.4
7.8
δD(H
2O)
‰ (S
MO
W)
1.4
–0.
1–
–2.
91.
7–
––
–2.5
…0.
5–2
…0.
5–1
.0…
0.5
–0.
7
δ18O
(H2O
)‰
(SM
OW
)1.
461.
540.
530.
5–0.
60.
820.
77–
0.8… 1.1
0.7… 0.9
0.65
0.6
0.8
0.43
δ11B
‰
––
––
––
–(3
4.7)
–34
.2–
–10
.1
δ13C
(CH
4)‰
(PD
B)
––
–23.
9–2
2.0
––
–61
.5…
–52.
2–
––2
0.8
–18.
2–
–
δD(C
H4)
‰ (S
MO
W)
––
––
––
–153
…–1
17.6
––
––
––
δD(H
2)‰
(SM
OW
)–
––
––
––
––
––
––
δ3 He/
4 He
RA
––
8.34
8.3
––
––
––
––
–
δ13C
(CO
2)‰
(PD
B)
––
––7
.9–
––3
3.0…
–10.
6–
––
4.4
–4.
7–
–
Ref
eren
ceV
on D
amm
et a
l., 1
997
Cha
rlou
et a
l., 1
996;
Jean
–Bap
tiste
et a
l.,19
97, 1
997a
Von
Dam
met
al.,
200
3T
aylo
r,19
90M
asso
th e
t al.,
1989
Von
Dam
m,1
990;
Pal
mer
and
E
dmon
d, 1
989;
Sha
nks
and
Seyf
ried
, 198
7; E
vans
et a
l., 1
988
Zie
renb
erg
et a
l., 1
993;
Pa
lmer
and
Ed-
mon
d, 1
989
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S205
Tab
le 8
. (C
ontd
.)
Reg
ion
Mid
–Atla
ntic
Rid
ge
Hyd
roth
erm
al s
yste
mL
ogat
chev
, 14
°45′
NM
AR
K, 2
3° N
TA
G, 2
6° N
Bro
ken
Spur
,29
° NL
ost C
ity,
30° N
Rai
nbow
,36
°14′
NL
ucky
Str
ike,
37°1
7′ N
Men
ezG
wen
,37
°50′
N
No.
in F
ig. 1
514
1516
1718
1920
21
Ven
t1
Snak
e Pi
tB
lack
smok
ers
Whi
tesm
oker
s4
Y3
Eif
fel
Font
aine
No.
in T
able
632
3334
3536
3738
3941
42pa
ram
eter
unit
Srμm
ol/k
g13
850
–10
391
42.9
–42
.9–4
8.0
119
7611
187
Sr/86
Sr–
0.70
28–
0.70
380.
7046
0.70
347
––
––
–
δ34S(
H2S
)‰
(C
D)
2.2–
2.8
4.9
4.9–
5.0
6.6–
7.5;
8.6
–0.
5–1.
0–
2.4–
3.1
––
–
δD(H
2O)
‰ (
SMO
W)
–3.
1–
3.0
––
––
––
–
δ18O
(H2O
)‰
(SM
OW
)–
1.9;
2.3
72.
301.
57–
––
––
––
δ11B
‰
–26
.8–
30.9
3525
––
––
–
δ13C
(CH
4)‰
(PD
B)
–14.
6…–1
3.8
––
–8.0
…–
9.5
––1
8… –19
–17…
–17.
6–1
3.4…
–13.
0–1
3.0
–12.
7–1
9.6
–13.
6–1
4.6
–15.
8
δD(C
H4)
‰ (
SMO
W)
––
––
––
––
––
–
δD(H
2)‰
(SM
OW
)–
––
––
––
––
––
δ3 He/
4 He
RA
–7.
9–8.
68.
47.
5–
––
–8.
138.
138.
65
δ13C
(CO
2)‰
(PD
B)
–4.
3… +1
––
9.0
–8.4
…–1
0.0
––5
.6–
–4… +1
–7.2
–10.
6–
9.1
–4.
3–3
.15
Ref
eren
ceL
ein
et a
l.,20
00;
Cha
rlou
et a
l., 2
002
Cam
pbel
let
al.,
198
8;G
amo
et a
l., 2
001
Jean
–Bap
-tis
te e
t al.,
19
91,
1997
; Lei
n et
al.,
200
4
Edm
ond
et a
l., 1
995;
Jea
n–B
aptis
te e
t al.,
199
7; K
nott
et a
l., 1
998;
Lei
n, 2
001;
G
amo
et a
l., 2
001;
Cha
rlou
et
al.,
200
2
Jam
es e
t al.,
19
95; C
har-
lou
et a
l.,
2002
; Lei
net
al.,
200
4
Lei
n et
al.,
20
04L
ein
et a
l.,
2000
; Cha
r-lo
u et
al.,
20
02
Jean
–Bap
tiste
et a
l., 1
998;
Cha
rlou
et a
l., 2
000
S206
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 8
. (C
ontd
.)
Reg
ion
Indi
an O
cean
Bac
k-ar
c ba
sins
of
the
Paci
ficIs
land
arc
s an
d vo
lcan
oes
of th
e Pa
cific
Hyd
roth
erm
al s
yste
mR
odri
gues
Tri
ple
Junc
tion,
Kai
rei F
ield
Lau
Bas
in,
Vai
Lili
fiel
d
Mar
iana
Tro
ugh
Eas
t Man
usba
sin,
Des
mos
fie
ld
Oki
naw
a T
roug
h,M
inam
i–E
nsei
Kno
ll
Oki
naw
aT
roug
h,C
LA
M f
ield
Oki
naw
aT
roug
h,JA
DE
fie
ld
Oki
naw
aT
roug
h,Su
iyo
Seam
ount
Haw
aiia
nIs
land
s, L
oihi
Vol
cano
No.
in F
ig. 1
522
2326
2930
3133
34
Ven
tV
L–3
No.
in T
able
644
4548
5152
5355
56
para
met
erun
it
Srμm
ol/k
g77
120
78–
215–
227
–94
310
100
87Sr
/86Sr
0.70
410.
7044
0.70
354
–0.
7100
–0.
7089
–0.
703
δ34S(
H2S
)‰
(C
D)
6.9
––
–5.6
–1.
47.
4–
–
δD(H
2O)
‰ (
SMO
W)
4.8
––
–8.1
––
––
–
δ18O
(H2O
)‰
(SM
OW
)1.
90.
85–
0.3
––
––
–
δ11B
‰
–26
––
––
––
–
δ13C
(CH
4)‰
(PD
B)
–8.7
––
––
–41
.2–
40.7
…–3
6.1
–8.5
–
δD(C
H4)
‰ (
SMO
W)
––
––
––
––
–
δD(H
2)‰
(SM
OW
)–
––
––
––
––
δ3 He/
4 He
RA
7.9
–7.
7–8.
5–
7.0–
7.5
3.7–
3.8
6.1–
6.9
8.1
–
δ13C
(CO
2)‰
(PD
B)
–6.
2–
–4.
33–
–4.
0…–5
.3–3
.8…
–3.6
–5.0
…–
4.7
–1–
Ref
eren
ceG
amo
et a
l.,
2001
Fouq
uet e
t al.,
19
93; J
ean–
Bap
tiste
et a
l.,
1997
Tsu
noga
iet
al.,
199
4;
Palm
er a
nd E
d-m
ond,
198
9
Gam
o et
al.,
19
97G
amo
et a
l., 1
991;
Ish
ibas
hi e
t al.,
199
5;G
lasb
y an
d N
otsu
, 200
3T
suno
gai
et a
l., 1
994
Palm
er a
ndE
dmon
d,
1989
Not
e:R
A is
the
atm
osph
eric
hel
ium
isot
ope
ratio
, 1.4
× 1
0–6.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S207
Tabl
e 9.
Con
cent
ratio
ns o
f di
ssol
ved
gase
s in
oce
anic
hyd
roth
erm
al s
olut
ions
Reg
ion
Eas
t Pac
ific
Ris
e
Hyd
roth
erm
al s
yste
m21
° N13
° N11
° N9°
17′ N
17°2
5′ S
18°1
5′ S
No.
in F
ig. 1
51
34
56
7
Ven
tN
GS
OB
SSW
Ven
t “F”
,19
91V
ent “
F”,
1994
Nad
irA
kort
a
No.
in T
able
61
23
78–
910
1112
1314
para
met
erun
it
He
10–6
cm
3 /kg
17.3
15.8
9.10
30–4
527
–50
20–5
0–
–24
18
CH
4cm
3 /kg
1.24
0.54
80.
411
0.62
–1.2
41.
151.
5–2.
6–
–7.
5 μm
ol/k
g13
3 μm
ol/k
g
H2
cm3 /k
g30
.537
.59.
9–
3.2
10–1
1.2
1.8
0.30
150
μmol
/kg
40 μ
mol
/kg
Ref
eren
ceW
elha
n an
d C
raig
, 198
3M
erliv
atet
al.,
198
7K
im e
t al.,
198
4V
on D
amm
et a
l.,19
97C
harl
ou e
t al.,
199
6;Je
an–B
aptis
te e
t al.,
199
7a
Reg
ion
Juan
de
Fuca
Rid
ge
Hyd
roth
erm
al s
yste
mE
ndea
vour
seg
men
t, 47
°57′
NA
xial
Vol
cano
, ASH
ES
fiel
dC
left
seg
men
t, 45
° N
No.
in F
ig. 1
510
1112
Ven
tL
obo
Infe
rno
Cra
ckV
irgi
n M
ound
Plum
eV
ent1
No.
in T
able
619
2224
2526
27pa
ram
eter
unit
He
μmol
/kg
–2.
458.
4411
.2–
–
CH
4"
–25
––
8411
8
H2
"16
6–
––
351
527
C2H
6"
––
––
0.2
–
C3H
8"
––
––
0.04
9–
Ref
eren
ceT
ivey
, 199
5B
utte
rfie
ld e
t al.,
199
0E
vans
et a
l., 1
988
S208
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Tab
le 9
. (C
ontd
.)
Reg
ion
Mid
–Atla
ntic
Rid
ge
Hyd
roth
erm
al s
yste
mL
ogat
chev
,14
°45′
NM
AR
K,
23° N
TA
G,
26° N
Bro
ken
Spur
, 29°
NL
ost C
ity,
30° N
Rai
nbow
,36
°14′
NL
ucky
Str
ike,
37°
17′ N
Men
ez G
wen
,37
°50′
N
No.
in F
ig. 1
514
1516
1718
1920
21
Ven
tSn
ake
Pit
Y3
Eif
fel
Font
aine
No.
in T
able
632
3435
3738
3941
42pa
ram
eter
unit
He
μmol
/kg
–2.
0–
––
–12
–26
(10–6
cm
3 /kg)
–
CH
4"
2100
6212
4–14
765
–130
130–
280
2200
400
680
2150
300–
1200
2500
H2
"12
000
190–
480
150–
370
430–
1030
250–
430
1300
076
.712
244
.2
930
1600
0
C2H
6"
––
––
–1.
097
0.06
50.
159
808
C3H
8"
––
––
–0.
048
–0.
009
0.01
8
Ref
eren
ceC
harl
ouet
al.,
200
2Je
an–B
aptis
te
et a
l., 1
991;
Cha
r-lo
u et
al.,
200
2
Lei
n et
al,
2000
;C
harl
ou e
t al.,
200
2K
elle
y et
al.,
2001
*1 ; Lei
net
al.,
200
4*2
Lei
n et
al,
2000
;C
harl
ou e
t al.,
2002
Jean
–Bap
tiste
et a
l., 1
998;
Cha
rlou
et a
l., 2
000
Reg
ion
Indi
an O
cean
Bac
k-ar
c ba
sins
, isl
and
arcs
, and
vol
cano
es o
f th
e Pa
cific
Hyd
roth
erm
al s
yste
mR
odri
gues
Tri
ple
Junc
tion,
Kai
rei F
ield
Lau
Bas
in,
Vai
Lili
fie
ld
Oki
naw
a T
roug
h,M
inam
i–E
nsei
Kno
ll
Oki
naw
a T
roug
h,C
LA
M f
ield
Oki
naw
a T
roug
h,JA
DE
fie
ld
Oki
naw
a T
roug
h,Su
iyo
Seam
ount
Haw
aiia
n Is
land
s,L
oihi
Vol
cano
No.
in F
ig. 1
522
2329
3028
3334
Ven
tV
L–3
Pele
’s v
ents
No.
in T
able
644
4551
5253
5556
para
met
erun
it
He
μmol
/kg
––
0.1–
0.7
0.1
0.3–
0.9
0.3;
0.6
–
CH
4"
823–
526
00–7
000
3100
–490
025
00–7
100
108;
140
7.25
4
H2
"–
25–3
412
–13
110–
150
10–5
025
–
Ref
eren
ceG
amo
et a
l., 2
001
Fouq
uet e
t al.,
1993
Gla
sby
and
Not
su,
2003
Gla
sby
and
Not
su,
2003
*3Is
hiba
shi e
t al.,
1995
Tsu
noga
i et a
l., 1
994;
Gla
sby
and
Not
su, 2
003
Kar
l et a
l., 1
988*
4
Not
e:D
ata
not e
xtra
pola
ted
to th
e en
d-m
embe
r: *
1 Mg
= 9
–19;
*2 M
g =
11.
0–18
.8; *
3 Mg
= 3
3–53
; and
*4 M
g =
50.
3 m
mol
/kg.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S209
abundances depend on many factors, the empiricaldescription of such dependencies is not straightfor-ward. The nature of many important chemical featuresof solutions remains controversial. Among them are theprevalence of Zn over Cu in the hydrothermal systemsof 21° N EPR and the Juan de Fuca Ridge; Cu over Znin the solutions of the TAG field; S over Fe in the 21° NEPR and some other systems; and Fe over S in solutionsfrom TAG, 45° N on the Juan de Fuca Ridge, and theExplorer Ridge. These relationships are very important,because they create a mineralogical and geochemicalportrait of the ores.
A comprehensive review of all the characteristics ofoceanic hydrothermal processes is beyond the scope ofthis paper. Such very important geochemical problemsas the role of hydrothermal processes in the mass fluxesand balances in the ocean, the activity of specific bio-logical populations of hydrothermal systems, and someother problems were not considered. These complexproblems, as well as the strategy and methods of pros-pecting for hydrothermal sulfide ores on the oceanfloor, are not addressed in this work. A detailed discus-sion of these issues can be found in some general stud-ies (e.g., Lisitsin et al., 1990; Hydrothermal Sulfide…,1992).
3.2. Logical Scheme of the Process and Its Experimental Substantiation
A geological and geochemical model for the hydro-thermal process in the oceanic crust was developed inthe 1970s and early 1980s. Its early variants were basedon the results of investigations of ophiolitic complexesand metamorphic rocks recovered by deep-sea dredg-ing and drilling, geophysical data, experimental simu-
lation of seawater–basalt interactions, numerical mod-els of convection systems, etc. (Spooner et al., 1977;Humphris and Thompson, 1978; Wolery and Sleep,1976; Lister, 1980; Rosenbauer and Bischoff, 1983).After the discovery and reconnaissance investigation ofblack smokers, the geochemical scheme of the processwas revised and constrained by Edmond et al. (1982).Their results were repeatedly reproduced in the subse-quent literature. Taking into account the most recentdata, geochemical processes in an oceanic hydrother-mal system are shown in a simplified form in Fig. 16.
Hydrothermal systems at mid-ocean ridges are con-vective systems with an endogenous heat source(magma body) and a fluid phase strongly dominated bythe exogenous component (seawater). Seawater perco-lates through fractures in the tholeiitic basalts of theoceanic crust, is gradually heated, and reacts with thecountry rocks. The basalts are replaced by secondarymineral assemblages of smectite, chlorite, and propy-lite facies with increasing temperature. Some seawatercomponents, including Mg, SO4, K, U, and dissolvedoxygen, are fixed in the solid phase. Simultaneouslymany components, including the ore elements Fe, SII,Cu, Zn, and Pb, are extracted from the basalts. The mostextensive water–rock interaction occurs in the hottestpart (focus) of the hydrothermal system near contactwith a magma chamber.
Magma-derived hydrothermal solutions releasedfrom the crystallizing melt may be introduced into theflow of hydrothermal solution in this zone. An area ofboiling and separation of water and vapor phases maybe formed near the magma body and in the lower partof the upwelling limb.
The movement of a hydrothermal solution in theupwelling limb of a convection system is rapid and
H2S, Fe, CH4, Mn, 3He
CO2, H2O, CH4,H2, H2S, 3He
K, Ca, Fe, Si, SII,Cu, Zn, Pb, H2
Na
Mg, SVI
K, P, U,O2, H2O
Massive sulfides
Magma chamber
SeawaterMetal-bearing sedimentsSeawater
Fe oxidesZeolitesSmectites
Anhydrite
ChloriteAlbiteEpidote
ActinoliteFrom basalt:
QuartzFe chlorite
PyriteChalcopyrite
Sphalerite
Magmatic fluid
?Zoneof boiling?
Fig. 16. Generalized relationships of geochemical processes in the hydrothermal system of a mid-ocean ridge.
S210
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
nearly adiabatic. Cold seawater can penetrate laterallyinto the upwelling flow, which results in the subsurfacemixing and cooling of the hydrothermal solution. Thehydrothermal solution emanates into the water columneither as a localized jet (smoker) and/or by disperseddischarge typical of the flanks of large sulfide mounds(for instance, in the Mid-Atlantic and Juan de FucaRidge). Solid phases, including ore minerals, are pre-cipitated from the hydrothermal solution owing to anabrupt temperature decrease and a change in composi-tion due to mixing with bottom water. During activedischarge, a considerable portion of ore substances isexpelled into the water mass, forming hydrothermalplumes.
Sulfide edifices growing on the seafloor at hydro-thermal vent sites show zonal internal structures andmay have complex growth histories. The character ofore deposition changes during the growth of sulfide edi-fices (Krasnov et al., 1988), because the role of slowpercolation of hydrothermal solutions through the orebody increases gradually. The ore components of thesolution are almost completely deposited within thelarge sulfide body, and the extent of metasomaticreplacement of previously formed phases increases dra-matically.
Several points of this logical scheme have beenwidely debated in the literature. Some important ques-tions are not yet solved. Since the logical scheme isimportant for the construction of the thermodynamicmodel of the process, these problems must be analyzedin more detail in order to obtain comprehensively sub-stantiated conclusions.
Hydrodynamic structure of circulation and heat andmass transfer in a convection cell. The existence ofconvective heat transfer in the oceanic crust was pre-dicted from the results of geophysical surveys (Lister,1972). Direct evidence for the occurrence of convectionwas obtained by deep-sea drilling. Water movementwas documented in a number of boreholes near modernspreading centers. For instance, hot water inflow and asharp increase in temperature at the borehole bottomwere observed in Hole 482C (DSDP Leg 65) after thecompletion of drilling operations (Duennebier andBlackington, 1983). Seawater uptake at a rate of 6–7 m3/h was revealed by hydrogeological measurementsin comprehensively studied Hole 504B in the CostaRica rift. The pressure deficit relative to the hydrostaticlevel was up to 10 atm (Becker et al., 1983), i.e., thisborehole penetrated the region of the descending flowof a convection cell.
There is little direct evidence on the distribution oftemperature within hydrothermal systems. Only outputtemperatures were measured, and temperature at thebase of the upwelling limb (focus of the system), whichis probably the maximum in the system, can be recon-structed. For a typical measured value of 350–365°C(Campbell et al., 1988a, 1988b; Von Damm, 1990) anda correction of 20–30°C for adiabatic cooling during
ascent (Bischoff and Pitzer, 1985), the maximum tem-perature is 370–395°C.
Methods of mathematical modeling were widelyused for the investigation of the character of water cir-culation in the oceanic crust. Early models were simpli-fied and often considered as only one-dimensionalproblems (Lowell, 1975; Ribando et al., 1976). Theproblem of heat transfer from rocks to solution has beenexplored on the basis of such models. It has long beenestablished by observations of submarine eruptions thatthe duration of direct contact between seawater andmelt must be very short and is of some importance onlyduring magma intrusion and eruption, which can beexemplified by the observations of April 1991 at 9° Non the EPR (Von Damm et al., 1995). Hydrothermalsolutions are heated during contact with hot rocks. Ofspecial importance for the understanding of this pro-cess was the cracking front hypothesis of Lister (1974,1980). This hypothesis is based on the fact that coolingof a hot rock causes volume reduction and formation ofa system of contraction fractures, which serve as path-ways for water penetration into the hot rock body. Theconvective circulation of water enhances cooling,which results in the propagation of the cracking front,i.e., this process is self-supporting. It is limited only bythe plasticity of hot rocks under lithostatic pressure.Under oceanic crust conditions, the position of thecracking front and, correspondingly, water penetrationcorresponds to the 400–450°C isotherm. However, theLister model considered only heat removal from a solidrock and therefore overestimated the rate of crackingfront propagation. This implied a very short-term exist-ence of magma chambers even beneath fast-spreadingridges, which was in conflict with geological observa-tions (Sleep, 1983). This controversy was resolved byCann and Strens (1982). They argued that, in additionto water circulation in solid rocks, the thermal model ofthe process must take into account melt convectionwithin the magma chamber. Owing to convective mix-ing within the chamber accompanying crystallizationdifferentiation, heat is continuously conveyed to theroof and the chamber is solidified mainly at the expenseof the cumulus pile (corresponding to the layered intru-sive series of ophiolites). The roof of the chambergrows mainly by the freezing of melt against it, whichproduces a layer of isotropic gabbro. Cann et al.(1985/1986) demonstrated that the critical parameter ofa thermal model is the thickness of the water-imperme-able layer of hot plastic rocks above the chamber roof,where heat is only transferred by conduction. In orderto explain the observed heat flux values, this layer hasto be a few tens of meters thick. The position of thecracking front in the Cann model is relatively stable aslong as melt is retained within the chamber. After thecomplete solidification of magma, the cracking frontbegins to move rapidly downward and hydrothermalconvection ceases almost immediately. The Cannmodel implies that the rocks directly above the roof ofthe magma chamber must experience prolonged and
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S211
extensive hydrothermal alteration in a stable thermalregime. This is consistent with the data on the characterand distribution of metasomatic alterations in the mod-ern oceanic crust and ancient ophiolitic complexes.
The hydrodynamic structure of convective circula-tion of hydrothermal solutions near a magma chamberwas elucidated by the calculation of two-dimensionalmodels. Various authors examined variations in thestructure of circulation under the influence of such fac-tors as magma chamber geometry (Fehn and Cathles,1979; Brikovski and Norton, 1989), nonuniform per-meability of crustal rocks (Fisher and Narasimhan,1991; Tutubalin et al., 1995; Malkovskii and Pek, 1997;Tutubalin and Grichuk, 1997a, 1997b), formation of thezones of boiling and supercritical fluid states (Brik-ovski and Norton, 1989), etc. The general conclusionfrom this work is that the most intense flow in the con-vection cell is established near the contact with themagma chamber and almost parallel to its roof. Thus,the isotherms are concentrated near the magma cham-ber. At any given time, the zone of high-temperaturewater–rock interactions is limited to a relatively thinlayer along the roof of the magma chamber and the ver-tical channel (Brikovski and Norton, 1989; Tutubalinet al., 1995).
The geophysical data on the depth of magma cham-bers in the axial zones of mid-ocean ridges allow esti-mation of the depth of water circulation. The convec-tion cells of axial hydrothermal systems have a charac-teristic vertical dimension of 1.5–3.0 km. According tothe results of hydrodynamic modeling, their lateralextension is about 6 km (Brikovski and Norton, 1989).
After the discovery of significant variations in thesalinity of oceanic hydrothermal solutions, Bischoffand Rosenbauer (1989) proposed an alternative modelof two-layer convection in order to explain these obser-vations and the accompanying geochemical effects.According to their model, the lower convection cell isfilled with boiling brines, and the upper layer is a tradi-tional cell recharged by seawater. Heat transfer betweenthe cells occurs mainly by conduction, and there isalmost no accompanying mass transfer. The authorsassigned the formation of the lower brine cell to the ini-tial phase of hydrothermal system development, duringthe intrusion of magmatic melt, which was probablyaccompanied by a large-scale vapor ejection. However,the investigation of a hydrothermal system at 9°47′ Non the EPR (Von Damm, 2000), which was affected bya magma body, did not confirm the existence of a long-lived brine reservoir within the oceanic crust. Hydrody-namic simulation demonstrated that a two-layer fluidsystem is unstable (Schoofs and Hansen, 2000).
Water–rock interaction and extent of metal extrac-tion from basalts: natural and experimental data. Thefirst high-temperature experiments on seawater–basaltinteraction were reported several years before the dis-covery of seafloor hydrothermal systems (Bischoff andDickson, 1975; Hajash, 1975; Seyfried and Bischoff,
1977; Mottl and Holland, 1978). The results of theseinvestigations were important for an understanding ofthe geochemistry of processes occurring in the oceaniccrust. It was found that during this interaction waterbecomes depleted in Mg and SO4 and enriched in Si,Ca, K, Ba, Li, B, H2S, Fe, and Mn, and the resultingsolution shows acidic pH values (Seyfried and Bis-choff, 1977; Mottl and Holland, 1978; Mottl et al.,1979; Seyfried et al., 1984). Anhydrite precipitatesfrom the solution. Smectite, mixed-layer smectite/chlo-rite, and high-temperature actinolite, talc, albite, waira-kite, magnetite, pyrite, and chalcopyrite develop afterbasalts. It was shown that the character of interactionstrongly depends on the water/rock ratio (W/R), andthere is a boundary W/R value of 50 ± 5 between therock-dominated and water-dominated conditions ofinteraction (Mottl and Seyfried, 1980).
The experimental data were, in general, supportedby the examination of natural objects, which was animportant argument in favor of the model of seawaterconvection in the oceanic crust. On the other hand, theexperiments exhibited some differences from the natu-ral process, which were discussed in detail by Rosen-bauer and Bischoff (1983). In particular, early experi-ments yielded considerable amounts of sulfide sulfurand low concentrations of chalcophile elements, whichwere not consistent with the natural prototypes. Thiseffect appeared to be related to the experimental proce-dure: the seawater–basalt reaction in an autoclave dif-fers from that in natural flow-through systems, wheresulfate sulfur is fixed in anhydrite at the beginning ofthe downwelling limb of the convection system (seebelow). Therefore, modified seawater, a synthetic com-position differing from seawater in the absence of sul-fate species, has been used since 1980s as a startingsolution in experiments.
Although very long experiments were conducted byMottl et al. (1979), Seyfried and Bischoff (1979) (up to1.5 y), and Janecky and Seyfried (1986) (almost twoyears), the natural assemblages of metasomatic miner-als characteristic of altered basaltoids have not beencompletely reproduced. The most important differenceis the lack of chlorite and epidote, which are the majorminerals of metabasalts, in the experimental products.The reason is probably related to the very sluggishtransformations of metastable smectites and mixed-layer smectite/chlorite that formed rapidly during theearly stages of these experiments (Mottl, 1983). Evenin the longest experiments, the system did not reachequilibrium and the composition of solutions continuedto change.
When the general characteristics of seawater–basaltinteractions were established, the experimentersfocused their attention on the determination of condi-tions providing the observed metal concentrations innatural solutions and the estimation of reactions buffer-ing solution compositions. Of crucial importance forthe solution of these problems was a study by Seyfried
S212
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
and Janecky (1985), who demonstrated a strong pres-sure effect on the concentrations of metals in solutionunder near critical conditions. These authors obtainedmetal concentrations corresponding to those observedin oceanic hydrothermal systems (Table 10). Currentlyall major characteristics of the chemical composition ofoceanic hydrothermal solutions have been experimen-tally reproduced.
In addition to basalts, seawater alteration was exper-imentally simulated with other rock types: peridotites(Seyfried and Dibble, 1980; Hajach and Chandler,1981; Janecky and Seyfried, 1986), rhyolites andandesites (Dickson, 1977; Shiraki et al., 1987), andsedimentary rocks (Bischoff et al., 1981; Thornton andSeyfried, 1985, 1987).
It is important to note before further discussion thatboth the experimental and natural data indicate that allthe rock-forming minerals of tholeiitic basalts(pyroxenes, plagioclases, and olivine phenocrysts)readily react with solution to produce secondaryphases. Samples of metasomatized basalts (i.e., thosewith altered K, Na, Ca, and Mg contents) do not displaya distinct zoning, which is characteristic of metaso-matic rocks associating with ore veins.
Little is known about the mineral forms of ore ele-ment occurrence in fresh and hydrothermally alteredbasalts (Nesterenko and Al’mukhamedov, 1973; Hein-richs et al., 1980; Bideau et al., 1985; Doe, 1994;Hydrothermal Sulfide…, 1992). Both in fresh andaltered basalts, copper usually forms its own mineralphase, chalcopyrite. In fresh basalts, lead usually sub-
stitutes for potassium in potassium feldspar. Duringmetasomatic changes, lead is fixed in galena. The spe-ciation of Zn in fresh basalts is dual: it occurs inapproximately equal amounts in orthosilicates substi-tuting for Mg and magnetite substituting for Fe. Inaltered basalts zinc probably forms its own phase,sphalerite, and can be incorporated in Mg-bearinghydrous silicates (chlorite and actinolite). Although theavailable data are fragmentary, they allow us to make animportant conclusion concerning the model: ore ele-ments do not occur in alteration-resistant mineralphases but do react with metasomatic solutions inapproximately the same way as major components.This greatly simplifies the model and allows us to con-sider the initial rock as a homogeneous reactive mate-rial.
Anhydrite deposition. Experiments with seawatershowed that its heating results in extensive anhydriteprecipitation. This observation raised the question ofthe relation of these experiments and the geochemicalscheme constructed on their basis to the natural pro-cess, because anhydrite was not found in metamor-phosed basalts dredged from the ocean floor or recov-ered from shallow boreholes during the early stages ofdeep-sea drilling. Seyfried and Bischoff (1979) andother authors explained this controversy by anhydritedissolution in cold water during the attenuation of con-vection in the hydrothermal system. This explanation isin agreement with the available data on a negative tem-perature dependency of anhydrite solubility in water.As a result, seawater is undersaturated in anhydrite at
Table 10. Concentrations of ore metals and other components in hydrothermal solutions (H2S, Fe, Mn, and SiO2 are in mmol/kg;and Zn and Cu are in μmol/kg) from experiments under near critical conditions (Seyfried and Janecky, 1985)
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S213
temperatures below 150°C. When deep-sea boreholespenetrated the hot zones of active hydrothermal sys-tems, vein anhydrite was systematically detected indrill cores (e.g., Alt et al., 1983, 1989). Drilling in theTAG hydrothermal field revealed a surprising amountof anhydrite in the interior part of the edifice and theupper part of the stockwork zone. Mills et al. (1998)explained this fact by anhydrite formation through sub-surface mixing with cold seawater.
Secondary mineral associations. Based on the anal-ysis of experimental data and results of studies of natu-ral objects, Mottl (1983) proposed a general scheme ofsecondary mineral assemblages formed by seawater–basalt interaction. The main parameter of this scheme isthe W/R ratio, which is a measure of the metasomaticeffect produced by seawater (mainly Mg introduction).High ratios (typical of the low-temperature part of theconvection system) give rise to the quartz + smectites +chlorite ± talc assemblage, while Ca, Na, K, Mn, Cu,Zn, and Ba are completely removed from the rock, andthe solution is weakly acidic to neutral. The excess ofrock produces the chlorite + albite + epidote + actino-lite + sphene ± quartz assemblage, which controls themajor element composition of the solution. Conditionsfavorable for such an excess exist in the high-tempera-ture part of the system.
Rock/water ratio. A fundamental parameter of themodel of a flow-through hydrothermal system is theproportion of the rates of solution–rock interaction andmass transfer. In steady-state experiments, this factor ismonitored by prescribing the proportion of basalt andsolution. The influence of this parameter on the experi-mental results was described above. In the MSFRmethod for the model of an extended nonisothermalsystem with fissure percolation (Section 2.2), the simi-larity criterion is the ratio of effectively reacting massesof fresh water and solution summed over the flow lineof solution, ΣR/W. This parameter was estimated by twomethods in active oceanic hydrothermal systems: fromthe concentrations of mobile elements and from the Srisotope ratios of emanating hydrothermal solutions.
The former method was proposed by Von Dammet al. (1985) and is based on the fact that some elementsare almost completely extracted into solution duringseawater–basalt interaction. Given their initial contentin water and basalt and enrichment in the black-smokersolution, the amount of fresh basalt that reacted withthe given solution portion can be calculated. If the ele-ment considered is not completely transported intosolution, the value may be underestimated. Conse-quently, when ΣR/W is determined from several ele-ments, the maximum value should be used. Table 11shows the data obtained by Von Damm et al. (1985) forthe hydrothermal systems of 21° N EPR. It can beclearly seen that the alkali elements Li, K, and Rb, aswell as H2S, which are highly mobile in high-tempera-ture hydrothermal systems, yield identical ΣR/W esti-mates.
The second method makes use of the fact that sea-water and mid-ocean ridge basalts are significantly dif-ferent in 87Sr/86Sr (0.70906 and 0.70280, respectively,according to Faure, 1986), and oceanic hydrothermalsolutions show transitional characteristics. As there isno Sr isotope fractionation, and the concentrations of Srin seawater and fresh tholeiitic basalts are relatively sta-ble, the amount of rock reacting with the given solutionportion can be calculated from the isotopic mass bal-ance. This approach yields an approximate estimate,because it ignores the effect of Rayleigh exhaustionrelated to Sr partitioning from the moving solution intosecondary minerals. Nonetheless, strontium isotopesystematics and alkali element relationships yield sim-ilar estimates. Table 12 summarizes ΣR/W estimates foroceanic hydrothermal systems. It can be seen that theoceanic hydrothermal systems show similar ΣR/W val-ues ranging within 0.5–2.0. The values determined onthe basis of solution composition are the effectiverock/water ratios integrated along the flow line of solu-tion that are needed for construction of equilibriumdynamic models.
It is obvious from general considerations that theΣR/W values of hydrothermal systems may vary withtime, but the period of observation of real oceanicobjects is still very short, and no changes in solutioncomposition have been yet reported (Campbell et al.,1988a; Edmond et al., 1995) (except for the cases ofvolcanic events discussed in chapters 6 and 7).
The R/W values of ancient analogues were estimatedusing the strontium and oxygen isotopic compositionsof metasomatized rocks, and the balance relationshipsfor their interaction with seawater. Such an estimate isintegrated over time and corresponds to a given point.Hence, the values obtained for ancient and modern sys-tems from the analyses of different materials cannot bedirectly correlated. As can be seen in Table 12, the R/Wvalues of ancient massive sulfide deposits are systemat-ically lower than those of oceanic hydrothermal sys-tems. This difference probably reflects the longer dura-tion of hydrothermal processes in the ancient objects ascompared with the modern oceanic systems.
Table 11. Water/rock ratios in modern hydrothermal sys-tems obtained from the investigation of the composition ofsmoker solutions (Von Damm et al., 1985)
Indicatorcomponent Water/rock Rock/water
K 1.7–2.0 0.50–0.59
Li 1.1–1.8 0.56–0.9
Rb 0.27–1.9 0.53–3.7
H2S 1.7–2.1 0.48–0.59
Note: The water/rock ratio is commonly used in the geochemicalliterature, and the inverse ratio is employed in thermody-namic models. In order to facilitate comparisons, the tablepresents both parameters.
S214
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Existence of a boiling zone within the system. Vent-ing of heterogeneous water–vapor fluids was neverobserved during submersible dives at deep oceanhydrothermal sites. The existence of boiling zones inthe interiors of some convective systems can be sup-posed on the basis of variations in solution chemistrycorrelated with the concentrations of dissolved gases.This question is considered in more detail in Chapter 6.
Subsurface mixing. Starting from the investigationsof thermal springs in the Galapagos Spreading Center(GSC) (Edmond et al., 1979), there is a growing bodyof evidence that some solutions from oceanic hydro-thermal systems have interacted with rocks at tempera-tures higher than those directly measured. To explainthis discrepancy, it was suggested that the high-temper-ature solutions were mixed during their ascent to theocean floor with seawater that did not pass through thehot part of the system. This phenomenon was referredto as subsurface mixing. If it occurs, the decrease intemperature due to mixing must promote precipitationof metals from solutions in the crustal layers, and for-mation of disseminated and vein mineralization (stock-work). Such mineralization patterns were observed, forinstance, in an ancient inactive system in the GSC(Embley et al., 1988) and in drill cores from Hole 504B(Alt et al., 1989). The investigation of the TAG systemsuggested that subsurface mixing was probably respon-sible for the formation of considerable amounts ofanhydrite within this structure (Mills et al., 1998).
The commonly used method of processing solutioncompositions includes extrapolation to Mg = 0 anddoes not allow one to discriminate between the additionof seawater to hydrothermal solution during subsurfacemixing and seawater contamination during sampling.The compositions of solutions determined recently inseveral systems show very low Mg concentrations indi-
cating a total contribution from both sources of 4–5%.This allows us to ignore the effect of subsurface mixingfor the construction of thermodynamic models.
The questions on the role of magmatic fluids andinternal zoning of ore bodies are considered below inmore detail because of their fundamental importance tothe thermodynamic model.
Influx of magmatic fluids. The assessment of thecontribution of magmatic fluids to the hydrothermalprocess is one of the most important problems and hasa long history of research. As to the hydrothermal sys-tems of mid-ocean ridges, when the first data on thechemical and isotopic compositions of hydrothermalsolutions were obtained, it became clear that theirmajor (possibly, the only) source is seawater (Edmondet al., 1982; Wehlan and Craig, 1983). Oxygen andhydrogen isotope systematics are the main indicatorsused in modern geochemistry for the identification ofwater sources. Figure 17 displays the available data onthe compositions of oceanic hydrothermal solutions(Table 8) on the Taylor diagram. It indicates that theocean is the main water source of the hydrothermal sys-tems. According to the calculations of Bowers and Tay-lor (1985), the observed small isotopic deviations canbe adequately explained by seawater–basalt interactionduring convective circulation in fissures. The contribu-tion of the magmatic component is smaller than the res-olution of the isotopic method (<5%). As a result, theproblem of the participation of magmatic fluids in oce-anic hydrothermal systems received little attention inthe international literature over the given time period.The views of Russian researchers on this problem weremore contradictory (Butuzova, 1986a, 1986b; Grichukand Krasnov, 1989; Krasnov, 1994), which was proba-bly related to the long tradition of plutonism in Russiangeology.
Table 12. Reconstruction of the water/rock ratios of hydrothermal systems
Region Object Method Water/rock Reference
Modern and recent systems in the ocean
Costa Rica rift Hole 504B Sr isotope balance in rock 1.6 Kawahata and Shikazono, 1988
" " S isotope balance in rock 2–7 "
21° N EPR Sr isotope ratio in solution 0.7 ± 0.2 Albarede et al., 1981
Galapagos Spreading Center Rose Garden From 3He content 0.34 Dymond et al., 1983
" " From 226Ra content 0.19–0.53 "
Juan de Fuca Ridge,southern segment
Monolith From Li and B contents 0.56–0.94 Butterfield and Massoth, 1994
" Pipe Organ " 0.37 "
" Diffuse " 4.6 "
Ancient deposits
Iberian Pyrite Belt Huelva O isotope balance in rock 0.3–0.7 Fouillac and Javoy, 1988
Troodos ophiolites, Cyprus Agrokipia Sr isotope balance in rock 12 Kawahata and Scott, 1990
Semail ophiolites, Oman O isotope balance in rock 15–40 McCulloch et al., 1980
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S215
Since 1990s this problem has received renewedinterest owing to the discovery of a strong influence ofvolcanic and seismic events on hydrothermal activity(Von Damm et al., 1995; Von Damm, 2000; etc.). Vari-ous pieces of evidence suggested that magmatic fluidsmay be introduced into oceanic hydrothermal systems,which has a bearing on the behavior of metals in solu-tion: (a) a high concentrations of gases (CO2, H2S, andH2) in high-temperature hydrothermal springs (Ishiba-shi et al., 1995; Charlou et al., 2000; Seewald et al.,2003), (b) indications for basalt degassing (see Rubin,1997 for a review), and (c) discovery of ore elementminerals in gas–liquid inclusions from the igneousrocks of oceanic basins (e.g., Kamenetsky et al., 2002).It can be seen that these data have an indirect bearingon ore formation. However, they are indispensable forthe solution of the fundamental question in model con-struction: whether magmatic fluids must be taken intoaccount in the model.
A contribution from endogenous fluid to the hydro-thermal systems of ore deposits is usually identifiedfrom Pb, S, Sr, and He isotopic parameters. However,attempts to use these indicators for oceanic hydrother-mal systems were hampered by the following circum-stances. The aforementioned elements can be intro-duced into a hydrothermal system in two ways:(a) directly with magma-derived fluid or (b) in twostages, through capturing during basalt crystallization(as minerals, admixtures, gas inclusions, etc.) and sub-sequent leaching with hot seawater. In both cases thecomponents are transferred into solution without anychanges in isotopic ratios, which thus cannot be used todiscriminate between these two mechanisms. They alsocannot be used for the identification of a direct influx ofmagmatic fluids into oceanic hydrothermal systems.
The problem can be approached by means of indi-rect methods. They are certainly less accurate but canbe used to constrain natural values.
Estimation from heat and water fluxes. The fractionof magmatic fluid can be assessed from the amount ofmantle-derived water relative to the water flux in thehydrothermal systems of mid-ocean ridges. The flux ofmantle-derived water into the ocean was independentlyestimated by various authors. According to the theoret-ical model by O.G. Sorokhtin (Geophysics of…, 1979),it equals 1014 g/y. The calculation of water balance inthe hydrosphere (Timofeev et al., 1986) yielded a valueof 5 × 1014 g/y. A value of (1.1 ± 0.3) × 1014 g/y wasobtained from the concentration of H2O in magmaticmelts of mid-ocean ridges (Ito et al., 1983).
Water flux through the hydrothermal systems ofmid-ocean ridges was estimated by a number of authorson the basis of heat transfer (Q) calculations for hydro-thermal systems. This value was determined througheither thermal models or measurements of the 3He fluxand the 3He/Q ratio (e.g., Rona, 1984; Merlivat et al.,1987). The estimates obtained show a considerablescatter related primarily to the imprecise determinationof the world’s average temperature of effluent hydro-thermal solutions. Nonetheless, all the estimates arehigher than 1017 g/y. These values suggest that the con-tribution of magmatic water to the total flow of hydro-thermal solutions is no higher than 0.1–0.5%. Thisvalue is perhaps overestimated, because it is based onthe assumption that all the water transported from themantle by basaltic melts is released through the hydro-thermal systems of mid-ocean ridges. However, thisestimate has a global character and does not excludelocal variations.
Estimation from the heat balance of a hydrothermalsystem. A local estimate can be obtained through thecalculation of the amount of seawater necessary for the
–10–15 –5 0 5δ18O
–100
–80
–60
–40
–20
0
20δD
Taylor diagram Meteoric water;
Ocean
Mantle
–2 –1 1 20
–10
–5
5
10δD
Water-rock interactiontrend
“Magmatic”trend
Isotopic composition of water from oceanic hydrothermal systems
Fig. 17. Isotopic composition of water from oceanic hydrothermal systems in the Taylor diagram.
δ18O
S216
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
cooling of one mass unit of basaltic melt and its com-parison with the quantity of volatile components dis-solved in this melt. Heat production associated with theformation of igneous rocks includes the latent heat ofcrystallization and heat release due to basalt cooling.The former constituent is 0.4 kJ/kg (Cann et al.,1985/1986; Brikovski and Norton, 1989), and the latterdepends on the temperature drop. If the temperaturedecreases from 1150°C (approximate solidus tempera-ture) to 370°C (temperature in the center of a hydro-thermal system), heat release can be calculated usingthe average heat capacity of basalts (Physical Proper-ties…, 1984; Brikovski and Norton, 1989). The result-ing value is 0.69–0.82 kJ/g (0.75 on average).
The amount of water needed to cool one gram ofbasalt can be calculated from the amount of heatremoved, the change in water temperature (from 0 to350°C), and the average heat capacity of water in thistemperature interval. This estimate gives =
= 0.71 g water per gram of basalt.
The concentrations of water in primary MORBmelts were estimated by many authors (Table 13).According to their data, the average water content in themelt can be estimated as 0.2% or 0.002 g per one gram ofbasalt. If all this water were released during crystallization(which is certainly an overestimation), its fraction in theflow of cooling seawater would be only 0.28%.
This estimate is local (related to an individual cham-ber) but averaged over time. It must be taken intoaccount that the release of volatiles may vary during thelifetime of a magma chamber. Volatiles with a low sol-ubility in melt, primarily CO2, escape during the initialstages of magma chamber evolution, which is indirectlysuggested by the existence of popping rocks with a totalCO2 content of up to 0.85% (Sarda and Graham, 1990),whereas water behaves as an incompatible componentand is accumulated in the residual melts up to 1–2%
(Table 13). However, such strongly evolved melts canbe formed only immediately before the complete solid-ification of small isolated reservoirs producing Fe–Tibasalts and Fe-andesites (for instance, in the GalapagosSpreading Center; Beyers et al., 1983, 1984). These resid-ual magma bodies have low heat resources, and largehydrothermal systems can hardly be generated by them.
Estimation from CO2 flux. This method utilizes themain geochemical feature of magma-derived gases ofmid-ocean ridges, a strong predominance of CO2 overall other components (Kyser and O’Neil, 1984; Som-mer and Gibson, 1985; Gerlach, 1986; Gurenko et al.,1990). This topic was discussed in detail by Grichuk et al.(1988). Assuming that all the carbon dioxide dissolved inhydrothermal solution came directly from a magmaticsource, its concentration can be used to estimate the frac-tion of the magmatic component in the mixture.
The data of Table 14 show that the maximum frac-tion of magma-derived fluid estimated by this methodis only 0.02–0.3%. A higher CO2 content was detectedin some hydrothermal systems (Table 6). Solutionsfrom the Guaymas Basin contain up to 1 g/kg, but thisCO2 was derived mainly from the metamorphism of asedimentary sequence. A high CO2 content was mea-sured in the freshened fluids from the Juan de FucaRidge (Table 14 shows the maximum values for the Vir-gin Mound and Cantilever vents, 12.54 and 3.12 g/kg,respectively). The water of these vents is a condensateof the vapor that formed in the zone of phase separa-tion. Such vents usually show low metal contents andunstable regimes.
Elevated CO2 levels were reported in hydrothermalsystems from other geodynamic settings including theJADE and CLAM sites in the Okinawa trough (islandarc) and Loihi Seamount (hot spot): 4.0, 8.5, and13 g/kg, respectively. The above-described method ofestimation is not appropriate for such objects, becausethe proportions of H2O and CO2 may be different intheir magmatic fluid.
Table 13. Concentrations of water and other volatile components in glasses from mid-ocean ridge basalts
Region Object Mg#Concentration, %
ReferenceH2O CO2 CO CH4 Cl F
MAR basalts (average) 63 0.21 0.13 – – 0.01 0.02 Byers et al., 1983" basalts (average) 62 0.22 0.015 – 0.002 – – Grichuk et al., 1988
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S217
When using the constraints presented in Table 14,one should keep in mind that some CO2 comes intohydrothermal systems with seawater (decreasing theestimate by 0.013%) and is extracted from crystallizedbasalts (decreasing the estimate by approximately0.01%). The magmatic addition consists mainly ofCO2, and the fraction of magmatic H2O is several timeslower (Grichuk et al., 1988).
Thus, three independent methods indicate a verysmall contribution from magmatic fluids to the hydro-thermal systems of mid-ocean ridges, no higher than0.0n–0.n%.
On the other hand, the existence of a small influx ofmagmatic fluids is demonstrated by the data on CO2content in glasses and solidified basalts (Bottinga andJavoy, 1989; Gerlach, 1989) and U/4He values (Zinglerand Hart, 1986). These data are suggestive of a preerup-tive loss of gases with low solubilities in melts from themagma chambers of mid-ocean ridges.
The second question related to the assessment of therole of magmatic fluids in the metal content of oceanichydrothermal systems is the possible extent of oremetal transportation by these fluids.
The metal content of juvenile fluids can be esti-mated from the direct analyses of metalliferous volca-nic exhalations. Within the available worldwide dataset, the Great Tolbachik Fissure Eruption (GTFE) ispetrochemically most similar to oceanic basalts. Con-densates collected during the GFTE were negligiblyaltered by interactions with the country rocks and mete-oric water and reflected the compositions of magmaticvolatiles (Menyailov et al., 1984). The raw data andresults of calculations are given in Table 15.
It should be noted that the analogy between the flu-ids of mid-ocean ridges and GTFE is not quite ade-quate, because the solid products of the GTFE are pet-rologically more similar to ocean island basalts. TheGTFE gases are strongly dominated by water (83.5%H2O and 5.55% CO2), whereas gases from mid-oceanridge basalts display the opposite relation. The GTFEfluids are more efficient with respect to ore elementtransportation than the magmatic fluids of mid-oceanridges. However, even assuming an analogy betweenthem, Table 15 suggests that the observed metal con-tents in the oceanic hydrothermal systems require a fewtens of percent of Tolbachik-type fluids.
Rubin (1997) analyzed a large body of diverse infor-mation and distinguished a series of elements that mustbe extensively removed from magma chambers withthe fluid phase: As, Sb, Pb, Mo, Tl, and Po. A compar-ison with the data on the compositions of hydrothermalsolutions and ores from mid-ocean ridges clearly showsthat there is no enrichment in these elements. It shouldbe noted that a strong enrichment was observed in theores of the JADE island-arc hydrothermal system (seeGlesby and Notsu, 2003 for a review).
The role of magmatic mass transfer can be qualita-tively estimated from the Zn and Cd relationship. These
Table 14. Calculation of the fraction of magma-related fluidsfrom the concentration of CO2 in hydrothermal solutions
Object
Concentration ofCO2 in solution,
mg/kg (calculatedfrom the dataof Table 6)
Maximum fraction
of magmaticfluid, %*
EPR, 21° N 252 0.03
EPR, 13° N 475–735 0.06–0.09
EPR, 9°40′ N 502–682 0.06–0.09
EPR, 17° S 580 0.07
Juan de Fuca Ridge,Endeavour segment
127–970 0.015–0.12
Juan de Fuca Ridge, Axial Seamount, Inferno
2200 0.28
Juan de Fuca Ridge, Axial Seamount,Virgin Mound
12540 1.6
Juan de Fuca Ridge,Cleft segment
172–196 0.02–0.025
Juan de Fuca Ridge,Cleft segment,Cantilever (1999)**
3120 0.39
MAR, MARK, 23° N 230–295 0.03–0.04
MAR, TAG, 26° N 128–150 0.016–0.019
MAR, Lucky Strike 1250 0.016
Lau Basin, Vai Lili 334–686 0.04–0.08
Manus Basin 260 0.03
East Manus Basin,Pacmanus
1410 0.18
Seawater 101 –
* For a magmatic gas containing ≥80% CO2.** After Seewald et al. (2003).
Table 15. Estimation of the fraction of magmatic fluid fromthe composition of condensates from the Great Fissure Tol-bachik Eruption
elements are geochemically very similar and behavesympathetically in most natural systems. Therefore, theZn/Cd value is rather stable in nature. However, it wasfound that Cd is much more mobile in magmatic gasesthan Zn (e.g., Symonds et al., 1992). The emanation coef-ficient of Cd is two orders of magnitude higher than that ofZn (Rubin, 1997). (The reason for this difference is prob-ably that Cd, in contrast to Zn, can occur as Cd0 in mag-matic fluid.) Table 16 shows Zn/Cd values in someobjects, which suggest that there is no significant contribu-tion from magmatic fluid to hydrothermal systems.
Considering all these data, it can be concluded thatthe transportation of metals by magmatic fluids is not asignificant control on the composition and properties ofhydrothermal ores of mid-ocean ridges. This allows usto ignore this factor in the construction of a thermody-namic model for the hydrothermal system.
It should be noted, however, that the relationships ofelement sources are probably different in island-archydrothermal systems (Chapter 7).
Internal zoning of ore bodies and its evolution intime. Given the diversity of mineral assemblages andchemical compositions observed in oceanic sulfideores, there is a need to reproduce the spatial and/or tem-poral sequence of their formation.
Among the samples and observations obtained fromthe ocean floor, the most distinct zoning was revealed insmoker chimneys. Mechanisms of chimney develop-ment were proposed by Haymon (1983) and Grahamet al. (1988). According to their models, edifice growthbegins with the deposition of anhydrite, which is subse-quently replaced by sulfides starting from the innerchannel of the chimney. The scenario proposed by Gra-ham et al. (1988) includes a successive change in theidentity of predominant sulfides: marcasite + wurtzite–pyrite–bornite–chalcopyrite. A sulfide pipe formed bysuch concentric zones corresponds to the mature stage.
The subsequent evolution of the pipe includes oxida-tion and destruction, and its inner channel is overgrownwith the formation of the reverse zoning pyrite–marca-site–sphalerite–opal.
Such idealized patterns are not always observed,and other variants of vent development were reported(Hydrothermal Sulfide…, 1992). In particular, someelements of the ideal sequence may be missing, and dif-ferent mineral assemblages may occur (e.g., pyrrhotite+ isocubanite). In edifices with a solution dischargetemperature of lower than 350°C, porous cavernouspyrite–marcasite–sphalerite ores are often deposited,whereas copper mineralization is absent. Low- andmedium-temperature vents form small opal and bariteedifices (such objects were found in back-arc spreadingbasins; Lisitsin et al., 1991). The lack of copper miner-alization associating with low-temperature solutions isunanimously explained by a decrease in the transportcapacity of the solution. It is usually noted that Cu and Fesulfides can be deposited in the channelway. The samelogic is applied to barite edifices, because the reconstruc-tions of temperature conditions within such systems oftenyield values identical to those of black smokers.
Compared with chimneys, the situation with smokeredifices as a whole is more complicated, primarilybecause of the paucity of data. The growth of edificesoccurs largely through the formation and subsequentdestruction of ore pipes (Lisitsin et al., 1990; Bogdanovet al., 1997b), and their ore material should consist ofmixed chimney fragments of different compositions,but the available samples do not have such chaotic com-positions. Moreover, large edifices contain ore typesthat are absent in small ones, including massive pyriteand pyrite–chalcopyrite ores confined to the central partsof the edifices (Hydrothermal Sulfide…, 1992). The outerparts of large structures are usually made up of pyrite–marcasite–sphalerite ores and covered by oxidation crusts(Lisitsin et al., 1990).
The interior zoning of sulfide edifices was recon-structed by Hekinian and Fouquet (1985), who pointed outthe difference in the character of hydrothermal dischargebetween small individual chimneys and large edifices: thelatter are characterized by an increasing role of slow solu-tion percolation through the ore body. As a result, pro-cesses of metasomatic replacement become more intensein the inner parts and a large-scale zoning develops withinthe edifice. Proceeding from the data obtained in thehydrothermal field of 13° N EPR, Hekinian and Fouquetproposed a scenario of ore body evolution.8 A generalized
8 It was shown that some elements of this scenario are of onlylocal significance. Furthermore, the suggestion that the processbegins with the appearance of low-temperature springs and isaccompanied by an increase in temperature was not confirmed.Observations at 9° N on the EPR (Von Damm et al., 1995) dem-onstrated that the beginning of hydrothermal activity is very vig-orous and high-temperature. The deposition of anhydrite with asulfide admixture from high-temperature solutions mixed withseawater was observed during an experiment lasting many days atthe vents of the Endeavor hydrothermal field (Tivey et al., 1990).
Table 16. Zn/Cd values for various hydrothermal systemsand related objects after Metz and Trefry (2000), with modi-fications
Object Zn/Cd, molar
MORB 907–1100Juan de Fuca Ridge, Cleft segment,hydrothermal solutions
800–906
Juan de Fuca Ridge, Cleft segment,Zn–Fe sulfides
1220
MAR, TAG, hydrothermal solutions 620MAR, TAG, Zn–Fe sulfides 560–640Zn–Fe sulfides of other fields 630Condensate of volcanic gas, GFTE,Kamchatka*
40
Proportion in the magmaticdegassing flux**
8
* Menyailov et al. (1984).** Rubin et al. (1998).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S219
variant of the Hekinian–Fouquet scenario was devel-oped by Krasnov (Krasnov et al., 1990; HydrothermalSulfide…, 1992) (Fig. 18). It includes three stages:(a) an initial stage, when an embryonic anhydrite–sul-fide body is formed, and the major deposition processis the mixing of the hydrothermal solution with seawa-ter; (b) a stage of high-temperature hydrothermal activ-ity, when a sulfide edifice is formed mainly at theexpense of sulfide precipitation due to the abrupt cool-ing of hydrothermal solutions; and (c) a stage markedby a large hydrothermal edifice with predominantly dif-fuse venting and a leading role of metasomatic rework-ing in the inner parts of the ore body. Similar recon-structions were proposed for large bodies from the TAGfield by Lisitsin et al. (1990) and Rona et al. (1993).They were confirmed by the data obtained by drillingon an active mound of this field (Herzig et al., 1998)(except for the surprising abundance of anhydrite in theinterior of the mound and in the upper parts of thefeeder stockwork).
The change in the character of ore-forming pro-cesses is mirrored in the mineralogy of the ores: thelarge deposits contain abundant massive pyrite andpyrite–chalcopyrite ores, which form the central partsof edifices, whereas such ores either are absent or occurin minor amounts in the small bodies. This causes vari-ations in the bulk chemistry of ore bodies: the size ofore bodies is correlated with the degree of copperenrichment (Krasnov, 1990).
Metasomatic alteration in the central parts of largeedifices has important consequences for the behavior ofminor elements of the ores. They give rise to zone refin-ing phenomena, which produce massive ores stronglydepleted in trace elements (except for Se, Mo, and Co)compared with the outer parts of the edifices (Hekinianand Fouquet, 1985; Fouquet et al., 1988; HydrothermalSulfide…, 1992; Hannington et al., 1998). Krasnovpointed out some delicate differences in trace-element(including Ag and Au) geochemistry between the
sphalerite ores deposited during stage (b) from high-temperature solutions and the sphalerite ores formingthe outer portions of large edifices during stage (c). Thereason for these differences is probably related to theremoval of these elements during metasomatic pro-cesses in the central parts of the bodies.
The zoning developing in sulfide edifices on theocean floor resembles that of ancient massive sulfideore deposits. In contrast to the ocean, the zoning ofmassive sulfide deposits has been extensively studiedand described in a number of publications. Ore bodiesfrom many massive sulfide deposits in the ophioliticassociations of Cyprus and the Mugodzhar Hills(Southern Urals), the contrasting and continuous asso-ciations of the Urals, and the andesite–dacite–rhyoliteassociations of Japan, Australia, and the Caucasus showdistinct zoned structures, which are controlled byfeeder axes. The zoning of ore bodies is expressed by aregular increase in the abundance of main rock-formingminerals in the sequence pyrite–chalcopyrite–sphaler-ite–galena–barite in the direction from the footwall tothe hanging wall and laterally outward from the feederaxis. Siliceous (jasper) deposits with iron and manga-nese oxide phases usually occur above and at the flanksof the massive ores.
The following is a typical vertical zoning of the wellknown massive copper sulfide deposits of the Cyprustype (Zlotnik-Khotkevich and Adriyanova, 1987):(1) maghemite–magnetite ores at the base of the orebody, (2) massive and brecciated pyrite ores with mag-netite relicts, (3) massive chalcopyrite–pyrite ores,(4) brecciated sphalerite–chalcopyrite–pyrite ores, and(5) massive and banded sphalerite–pyrite ores withminor chalcopyrite; banding is due to the presence ofthin siliceous laminae at the roof of the body (Priorskoedeposit in the Mugodzhar Hills).
The nature of zoning in massive sulfide ore bodieswas long a subject of debate (see Franklin et al., 1981
(a) (b) (c)
1 2 3 4 5 6
Fig. 18. Schematic evolution of a black smoker after Hekinian and Fouquet (1975) modified in Hydrothermal Sulfide… (1992).(a) Stage of formation of an anhydrite embryonic structure on the ocean floor. (b) Stage of high-temperature hydrothermal activityand formation of a sulfide edifice. (c) Stage of diffuse hydrothermal activity. (1) Anhydrite, (2) Cu and Fe sulfides, (3) Zn and Fesulfides, (4) low-temperature Fe sulfides, (5) stringer and disseminated mineralization in basalt, and (6) direction of hydrothermalsolution movement.
S220
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
for a review). Many researchers (Smirnov, 1982; Ohm-oto et al., 1983) explained the heterogeneity of ore bod-ies by the sequential formation of ores from solutionswith different compositions, as a result of either the reg-ular evolution of the feeding hydrothermal system orthe repeated resumption of hydrothermal activity withchanging characteristics. Another popular explanationfor ore zoning postulates spatial variations in the condi-tions of ore deposition, which results in sequentialchanges in mineral assemblages (Large, 1977;Solomon and Walshe, 1979; Franklin et al., 1981).Some authors emphasized the role of redeposition andmetasomatic replacement of ore materials during for-mation and diagenetic alteration of the ore body(Skripchenko, 1972). The sequential influx of solutionswith different geochemical characteristics is suggestedby some features systematically observed in massivesulfide deposits, including intersection structures, cor-rosion and cementation phenomena, and replacementof minerals. On the other hand, the regular change inmineral assemblages on the scale of an ore deposit irre-spective of the particular geologic environment may beindicative of ore deposition from a single solution,whose properties and occurrence conditions vary regu-larly in space. Many recent studies noted the coexist-ence of zoning patterns related to different factors(Prospecting of…, 1985).
A comparison of oceanic hydrothermal occurrenceswith ancient massive sulfide deposits reveals an obvi-ous influence of deposition conditions (primarily tem-perature) on the zoning patterns of the ores. The Hekin-ian–Fouquet scenario for the evolution of ore edificesuses the concept of metasomatic replacements. Theidea of the evolution of parent solutions cannot be eval-uated, because the duration of observations in activehydrothermal systems is not yet sufficient. The role ofthis factor cannot be established by direct observationor comparative analysis of the structures of ore bodies,because the relevant drilling data are still scanty and thereconstructions of the interior structure of seafloor sul-fide edifices are oversimplified.
The analogy between modern and ancient ore bod-ies appears to be far from complete. Massive sulfidedeposits do not contain mineral associations corre-sponding to the early stages of oceanic system develop-ment (Fig. 18a). Strong zinc enrichment is typical ofoceanic sulfide bodies but was never reported in mas-sive sulfide deposits.
The considerations presented above were used as abasis for the logical scheme of oceanic hydrothermalore formation, which was used for the construction ofits physicochemical model.
3.3. Modeling of Modern Ore Formation in the Ocean: Previous Work
Numerical thermodynamic modeling of high-tem-perature seawater–basalt interaction was first per-
formed by T. Wolery in 1978, but the results of thatstudy were never published.
A thermodynamic model for the behavior of seawa-ter during heating to a temperature of 300–350°C,which is characteristic of oceanic hydrothermal sys-tems, was calculated by Reed (1982). The results ofmodeling were compared with the experimental data ofBischoff and Seyfried (1978), who studied the proper-ties of seawater between 25–350°C. The goal of thesecalculations was to demonstrate the possibility of anadequate thermodynamic description of the behavior ofseawater in the hydrothermal process. Good agreementwith experimental data allowed the simulation of morecomplex models involving silicate rocks.
The calculations of isothermal seawater–rhyoliteinteractions were performed by Reed (1982), and sea-water–basalt and seawater–peridotite interactions werecalculated by Grichuk et al. (1982) and Reed (1983).These calculations showed adequate agreement withthe available experimental data on the respective sys-tems. The general trends of changes in solution chem-istry were also consistent with natural observations.However, since the method of the degree of reactionprogress was used in the aforementioned studies, theyfailed to adequately reproduce sulfur behavior (the rea-son for this discrepancy was discussed in Section 2.1).Correspondingly, ore elements were also not properlyevaluated.
Further progress in the thermodynamic modeling ofoceanic hydrothermal systems was made in two direc-tions: (a) simulation of ore deposition during cooling ofhydrothermal solutions on the basis of analytical dataon the compositions of natural hydrothermal solutionsand (b) simulation of the formation of hydrothermalsolutions from seawater as a result of interaction withrocks. The former approach makes use of more specificinformation, but it is restricted to the zone of ore depo-sition and does not elucidate the general properties ofthe hydrothermal system. The latter approach encoun-tered considerable difficulties, because it employspoorly constrained model parameters. However, itallows the construction of a comprehensive thermody-namic model for a hydrothermal system, from the areaof hydrothermal solution generation to the zone of oredeposition, and the properties of this model can beexplored. Such a formulation exploits most comprehen-sively and efficiently all the advantages of the methodof thermodynamic modeling.
Ore deposition from smoker solutions was first sim-ulated by Janecky and Seyfried (1984), who calculatedthe mixing of hydrothermal solutions from 21° N EPRwith cold seawater. Their model took into account15 elements including Fe, S, Cu, Zn, Pb, and Ba (thecomposition of the multisystem with respect to miner-als and solution species was not specified). The calcu-lations ignored aluminosilicates and silicates; theformer, because Al was not present in the multisystem,and the latter, because preliminary calculations overes-
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S221
timated (in the authors’ opinion) the amounts of precip-itated quartz, talc, serpentine, and tremolite. The calcu-lations were carried out within a temperature intervalfrom 350 to 25°C and a pressure of 250 bar.
The results of these calculations proved to bestrongly dependent on whether the equilibriumbetween oxidized and reduced sulfur species isattained. Calculations with metastable sulfate reductionreactions yielded the following temperature sequenceof mineral precipitation: chalcopyrite–anhydrite–pyrite(pyrrhotite)–sphalerite–graphite–barite. A scenarioassuming an equilibrium sulfur speciation suggestedearlier pyrite deposition and formation of the followingsequence of copper-bearing minerals: chalcopyrite–bornite–chalcocite–covellite. The assemblages of oreminerals obtained in these models were similar to thoseobserved in nature, and the authors noted that the equi-librium sequence was closer to that observed on theinner walls of smoker chimneys. It is also similar to themineral zoning of ancient massive sulfide ores.
The analysis of the simulation procedure used byJanecky and Seyfried shows that it is a modification ofthe method of the degree of reaction progress by Helge-son (see Chapter 2), which includes a peculiar titrationof the hydrothermal solution with cold seawater with-out the removal of reaction products. The method ofJanecky and Seyfried does not, in principle, reproducethe differentiation of ore components. In the simulationof the equilibrium scenario (Janecky and Seyfried,1984, Fig. 12), the amounts of copper minerals (covel-lite and chalcocite) produced by the end of the processat temperatures below 100°C were identical to theamount of chalcopyrite at the beginning of the process.This is not consistent with the natural process, in whichcopper is removed from solution to form chalcopyriteand does not occur in the subsequent mineral assem-blages. Nonetheless, the results of Janecky and Sey-fried (1984) appeared to be similar in many respects tothe natural prototypes. They demonstrated that themethods of thermodynamic modeling are promising forthe analysis of even such relatively rapid processes ashydrothermal–sedimentary ore formation.
Bowers et al. (1985) used a similar simulation pro-cedure and a more comprehensive data set on the com-position of hydrothermal solutions from the 21° N EPRfield and the Guaymas Basin. Their multisystemincluded some additional elements (Al, Mn, and Ag),and the raw analyses were corrected for a pH shift dueto the cooling of samples. The studies of Tivey andMcDuff (1990) and Tivey (1995) were of considerableimportance for the simulation of oceanic hydrothermalmineralization. These authors explored heat and masstransfer and chemical interaction in the walls of smokerchimneys. The Tivey–McDuff model accounted forheat conduction, diffusion, solution flow through aporous wall, and local thermodynamic equilibriumbetween solution and solid phases. The problem con-
sidered was steady-state with respect to temperatureand concentration profiles.
The calculations yielded several variants of zoningin the chimney wall depending on the parameter F,which is a function of chimney diameter. The typicalzoning is as follows: (chalcopyrite)–(anhydrite +bornite)–(pyrite + bornite)–(silica + pyrite + sphalerite +chalcocite), which is similar to that observed in the nat-ural prototypes. It was shown that, for the processes ofthe chimney-wall scale and low porosity, diffusionplays a significant role in the formation of zoning,whereas the kinetics of precipitation is not important,i.e., equilibrium in the systems was reached rather rap-idly.
Thermodynamic modeling of hydrothermal solutionformation and complete models for oceanic hydrother-mal systems.
All the published models of the type considered usethe concept of a step flow reactor in some form. Thedownwelling limb of a convection cell is usually simu-lated by a chain of steps with increasing temperaturesmaintaining local equilibrium, whereas the upwellinglimb and the discharge zone are approximated by a sin-gle step or several steps with decreasing temperatures.The schemes of these models are in principle identical,but there are many particular differences. The mostsophisticated models were constructed by the author atthe Geochemistry Department of Moscow State Uni-versity (these results are described in detail in the nextchapter) and T.S. Bowers at Massachusetts Institute ofTechnology.
The problems that should be solved during the con-struction of models for convection hydrothermal sys-tems can be divided into two groups: (1) those inherentin any thermodynamic model of hydrothermal pro-cesses and (2) those specific to the scenario of recyclingwith hydrothermal–sedimentary ore deposition (such acombination defines the specific features of models foroceanic hydrothermal systems).
One of the general problems is the creation of aninternally consistent thermodynamic database. Grichuket al. (1985), Grichuk (1988, 1996), and HydrothermalSulfide… (1982) used the UNITHERM data bank(Moscow State University). The thermodynamic prop-erties of solid phases were taken mainly from Robieet al. (1978) and Thermodynamic Properties… (1978–1982) (see also Borisov and Shvarov, 1992). The databank used the model of Helgeson et al. (1981) for sim-ple ions before 1989, and the modified Helgeson–Kirkham–Flowers (HKF) model later on (Shock andHelgeson, 1988; Johnson et al., 1992). The propertiesof complex ions were calculated by a semiempiricalmodification of the Ryzhenko equation (Methods of…,1988; Borisov and Shvarov, 1992). The models byReed (1983), Bowers and Taylor (1985), and Bowers(1989) were based on the thermodynamic constantstaken mainly from Helgeson et al. (1978, 1981) andRobie et al. (1978). Silantyev et al. (1992) used the
S222
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
thermodynamic properties of substances from theDIANIK database (Vernadsky Institute of Geochemis-try and Analytical Chemistry, Russian Academy of Sci-ences), which employs the modified HKF model forsimple ions and the Ryzhenko equation modified byBryzgalin (1985) for complex species. It can be seenthat, although there are differences between the ther-modynamic databases used by various authors, they arenot very significant, because all the databases relymainly on the same fundamental compilations. The dis-crepancies between the results of simulation arerelated, to a greater extent, to different sets of mineralsand dissolved species. Considerable changes in theresults of simulation were only caused by the introduc-tion of the modified HKF model, which significantlyimproved the accuracy of the thermodynamic proper-ties of dissolved species at high temperatures and pres-sures.
With respect to the choice of the compositions ofinitial solutions and rocks, the models of oceanichydrothermal systems have important advantages overthe models of other hydrothermal objects, because thecountry rocks and initial seawater are fairly homoge-neous. The range of conditions is also relatively nar-row: temperatures between 100–150 and 350–400°Cand a pressure of up to 500 bar. The aforementionedstudies are practically identical with respect to theseparameters (except for Silantyev et al., 1992, which isdiscussed below).
There are several specific problems associated withthe simulation of oceanic hydrothermal systems:(a) procedure for prescribing the rock/water ratio con-trolling the intensity of interaction in the steps of thedownwelling flow and (b) scenarios of ore deposition.
A comprehensive model for the hydrothermal sys-tem was developed by Reed (1983). In this model, therecycling process was divided into three sequentialstages: (1) heating of seawater up to 300°C at 500 bar;(2) isothermal water–basalt reaction at 300°C and500 bar; and (3) solution cooling accompanied by oredeposition. Anhydrite precipitates from seawater dur-ing heating, which removes 1/3 of the initial SO4. Reedcalculated water–basalt interaction by the degree ofreaction progress method. The model trends of Mg, Ca,SO4, H2S, Fe, Ag, Cu, and Zn concentrations were inadequate agreement with the available experimentaland natural data. Two scenarios of ore deposition wereexplored by Reed: (a) conductive cooling and (b) mix-ing with seawater at 25°C. The simulation used solu-tions from the previous stage obtained at R/W = 0.045with H2S > (Fe + Zn + Cu) and R/W = 0.062 withH2S < (Fe + Zn + Cu). The cooling of the former solu-tion produces the quartz + pyrite + chalcopyrite + chal-cocite + sphalerite + acanthite assemblage, and excesshydrogen sulfide is retained in the solution. Duringcooling of the latter solution, magnetite is added to theabove assemblage and hydrogen sulfide is completely
fixed. In the scenario with mixing, chalcopyrite doesnot form and magnetite is replaced by hematite.
A detailed analysis of Reed’s study demonstratedthat the use of a two-stage downwelling limb (heatingof solution followed by reaction with basalt) and themethod of the degree of reaction progress significantlydistorted the results. Owing to the possible occurrenceof a back-reaction, Reed’s model exaggerated the sta-bility field of sulfate sulfur, up to the largest R/W valueconsidered, which resulted in the removal of calciumfrom the solution. The concentrations of hydrogen sul-fide appeared to be high at the expense of marine sulfatereduction, which is in conflict with sulfur isotope data.As was demonstrated by subsequent studies, the rela-tionships of hydrogen sulfide and ore metals obtainedin Reed’s model, with a hydrogen sulfide deficit at highR/W, were an artifact of the simulation procedure. Thisaffected the behavior of all chalcophile elements. Themixing scenario of ore deposition did not account forheat balance: all the mixtures were equilibrated at25°C, although their temperatures must depend on themixing proportions (Janecky and Seyfried, 1984).
Despite these shortcomings, Reed’s study played asignificant role in the development of concepts ofchemical processes in oceanic hydrothermal environ-ments (Krivtsov, 1987).
A complete model based on the method of a stepflow reactor was first developed for an oceanic hydro-thermal system by Grichuk and Borisov (1983) andGrichuk et al. (1985). This model considered seawaterinteraction with fresh basalt (i.e., passage of the firstportion of solution). It was assumed that the rock/waterratio is an exponential function of temperature, suchthat the weight ratio ΣR/W is 0.173–1.143 (Fig. 9).Three scenarios were considered for the upwelling limband the zone of ore deposition: (a) slow conductivecooling, which was also simulated by a step flow reac-tor; (b) rapid cooling during ejection on the seafloor butwithout a chemical reaction with seawater; and(c) cooling with concurrent mixing, which is similar tothe model explored by Janecky and Seyfried (1984).Because of the paucity of thermodynamic data avail-able at that time, the model was restricted to a temper-ature interval of 150–350°C at saturated vapor pressureand a 13-component multisystem including Fe, S, andPb as ore elements. The multisystem included 36 min-erals and 48 dissolved species.
The model mineral assemblages of metasomaticrocks and the compositions of model solutions formedin the downwelling limb were generally consistent withnatural observations. The calculation of the rapid cool-ing of model solutions produced typical mineral assem-blages of the suspended particulate matter of smokers.However, these calculations failed to reproduce theobserved concentrations of some components in thesolution. The results obtained for 350°C underesti-mated Ca concentrations and overestimated Fe, Al, andH2 concentrations, and the pHT values appeared to be
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S223
slightly alkalic. These discrepancies reflected errors inthe thermodynamic properties of dissolved species ofthese elements (chloride complexes of Ca and Fe andhydroxide complexes of FeIII).
Bowers and Taylor (1985) simulated the formationof oceanic hydrothermal solutions. Their model alsopresented the downwelling limb of a convection cell asa step flow reactor, and simulated the passage of thefirst portion of solution through it. The thermodynamicmultisystem of this model included 11 elements withonly Fe and S as ore elements, and its complete compo-nent composition was not specified by Bowers and Tay-lor (1985). The bulk model rock/water ratio was takenas 1.787, which corresponds to the estimates obtainedfor the hydrothermal systems of 21° N EPR(2 according to Von Damm et al., 1985). In order to dis-tribute this amount between the steps of the model, theauthors used the maximum likelihood estimates basedon some characteristics of mineral and chemical com-positions (appearance of epidote in the mineral assem-blage and an increase in solution pH), i.e., they solvedthe inverse problem. However, these criteria were notformalized in their study, and the choice of specific R/Wvalues for each step was largely arbitrary. Bowers andTaylor (1985) calculated equilibria using the EQ3/6program. This program utilizes the algorithm proposedby Helgeson, and its main parameter is the degree ofreaction progress (ξ, see Chapter 2). It cannot thereforecalculate equilibrium parameters for the given compo-sition of the rock–water system but constructs the pathof solution titration by the rock. Because of this, Bow-ers and Taylor calculated a closed system for each tem-perature step, with ξ changing from zero to the desiredvalue, after which the resulting solution was transferredto the next step. This method was referred to as thecomplex open-system model, and Bowers and Taylorargued that it is the most realistic simulation method.9
The path of such a process in the R/W–T coordinates isa broken line (Fig. 9b). Ore deposition in the upwellinglimb of the system was not considered by Bowers andTaylor (1985). An important feature of their model isthat it involved H and O isotope equilibria. The isotopicpart of the model was based on the Ohmoto method andutilized the results of chemical equilibrium calculationsas input information (see Chapter 5 for more detail). Inorder to realize isotopic calculations, a special modulewas implemented into the EQ3/6 code.
Bowers and Taylor (1985) reproduced rather realis-tically the mineral associations of oceanic metabasicrocks and obtained a generally adequate description ofrecycling-related changes in seawater composition.However, similar to the results of Grichuk et al. (1985),the calculated concentrations of Fe and Al in the solu-
9 It can be easily seen that, at equal final R/W values in reactorsteps, the result of calculations for such a model will be identicalto those obtained by the step flow reactor method proposed byGrichuk and Borisov (1983) and Grichuk et al. (1985), but thelatter is superior in terms of computational efficiency.
tion appeared to be overestimated, whereas those of Cawere strongly underestimated compared with the natu-ral prototypes, and the pHT values were neutral ratherthan weakly acidic. A fundamentally important resultwas obtained in the isotopic part of this model. The cal-culated values δD = +2.67‰ and δ18O = +2.0‰ forhydrothermal solutions appeared to be consistent withthose observed in natural hydrothermal systems (Table 8).This suggests that the isotopic shift of oceanic hydro-thermal solutions relative to seawater can be fullyexplained by reactions with rocks, and does not requireother hypothetic water sources (including magmaticfluids) in the hydrothermal systems.
Unfortunately, there are some doubts on the accu-racy of the model’s estimate of δD shift. The inputinformation of the Bowers and Taylor model includedextrapolated values for the fractionation coefficients ofhydrogen isotopes between minerals and water. Theirform of the extrapolation function was determined onthe basis of kaolinite–water experiments. For this pairTaylor obtained a sinusoidal dependency with a maxi-mum near 300°C using the data available at that time.He supposed that isotope fractionation with otherhydrous minerals could be approximated by similardependencies (except for epidote). It was later found(Gild and Sheppard, 1996) that the low-temperaturedata for kaolinite that were used by Taylor were errone-ous, because isotopic equilibrium was not reached inthese experiments. Correspondingly, the hydrogen iso-tope shifts obtained in the Bowers–Taylor model areprobably overestimated by a factor of two. However,this error does not affect the aforementioned fundamen-tal conclusion on the role of seawater.
The isotopic direction in the simulation of oceanichydrothermal systems was further developed by Gri-chuk (1988), Grichuk and Lein (1991), and Bowers(1989), who analyzed the behavior of sulfur isotopes,one of the major ore elements, in the descending con-vection flow. These results are discussed in Chapter 5.The chemical compositions of these models wereextended to include Cu and Zn with correspondingsolid phases and dissolved species.
The MSFR method was used to simulate the evolu-tion of an oceanic hydrothermal system by Krasnov etal. (1990), Grichuk and Lein (1991), and HydrothermalSulfide… (1992). These studies used the modified HKFmodel for the calculation of input thermodynamic data,which improved the agreement with natural prototypeswith respect to the concentrations of dissolved Ca, Al,and H2 and pH values.
A novel approach to the description of water–rockinteraction was proposed by Silantyev et al. (1992),who also utilized the concept of a multiwave step flowreactor. In order to simulate metamorphic alterations inthe oceanic crust, these authors considered a heteroge-neous crustal section composed of basalts, gabbroids,and ultrabasic rocks. They did not employ the principleof partial equilibrium and assumed that the whole rock
S224
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
body took part in interactions. The computational pro-cedure was designed accordingly: all water portionsreacted in each step with considerable amounts of rock(10 kg of rock per one kilogram of solution, whichapproximately corresponds to the porosity of basalts)and was subsequently moved to the next step. Themodel was calculated between 150–500°C with a tem-perature increment of 50°C. All the steps up to 350°C,inclusively, contained basalts, and the following stepswere either gabbroic or ultrabasic. The multisystemconsisted of 14 elements (including Fe, S, and Pb)forming 45 solid phases and 39 dissolved species. TheEQUILIBR computer program was used for calcula-tions. The computation was terminated after the pas-sage of 60 water portions, which resulted in an inte-grated rock/water ratio of 0.167 for the first step and0.83 for the last “basaltic” step (350°C).
A comparison with the data presented in Section 2.1shows that this method (step reactor with large rockmasses in steps) adequately reproduces an infiltrationmetasomatic column, but the desired compositions ofthe back zones can be reproduced with a very largenumber of solution portions (R/W must become lowerthan 0.02). Because of this, tremolite and epidote werepredominant in the calculated assemblage of the 150°Cstep, which was interpreted by Silantyev et al. (1992) asan example of the zeolite facies. It seems likely that theassumption on the effective interaction of the wholerock volume with water is not quite realistic for theupper part of the basaltoid section. On the other hand,for the hotter zones, the model showed good agreementwith the typical natural assemblages of metabasicgreenschist-facies rocks.
However, the compositions of model solutionsreported by Silantyev et al. (1992) differ in manyparameters from observations in oceanic hydrothermalsystems. For instance, the compositions obtained forthe 350 and 400°C steps, which most closely match theconditions of oceanic hydrothermal systems, are con-sistent with the natural data only with respect to Na,whereas the concentrations of K, Ca, SiO2, and Pb arelower by 0.5–2.0 orders of magnitude. It is also note-worthy that oxidized sulfur species prevail overreduced ones, which is not observed in oceanic hydro-thermal systems. This is indicative of flaws in themodel related to the thermodynamic description of thebehavior of dissolved species. It is also obvious that theprocedure of model verification must include a compar-ison of solution compositions, and the use of data onlyon mineral assemblages is not sufficient.
It should be noted that the model of Silantiev et al.for the interaction of water with ultrabasic rocks hasacquired special importance in connection with the dis-covery of hydrothermal systems related to ultrabasicrocks.
The results obtained during the past two decadesdemonstrated that quite comprehensive thermody-namic models can be created for the modern hydrother-
mal process in the ocean. On the other hand, the resultsobtained by the simulation of oceanic hydrothermalsystems revealed the most important unresolved prob-lems related to the method of model construction andverification of model results.
3.4. Model Description 3.4.1. Geological model
The geological model (logical scheme) of geochem-ical processes in a mid-ocean ridge hydrothermal sys-tem, which was discussed in detail in Section 3.2, canbe summarized as follows.
The model consists of three parts with distinctly dif-ferent characteristics of geochemical processes(Fig. 8): (a) the downwelling limb of the convectionsystem, where a hydrothermal solution is generated bythe interaction of seawater with hot rocks of the oceaniccrust; (b) the upwelling limb corresponding to thefeeder conduit of ancient deposits; and (c) the zone ofore deposition, the most important part of which is thesulfide body growing on the ocean floor.
Downwelling limb. The convection cells of hydro-thermal systems have characteristic vertical dimensionsof 1.5–3.0 km (extending to the roof of a magma cham-ber) and lateral dimensions of about 6 km. Seawaterpercolates downward through fractures and reacts withthe country rocks under gradually increasing tempera-tures. The high-temperature zone (focus of the system)directly contacts the magma chamber, and its maximumtemperature is 370–395°C. The maximum pressure inthe cell can be estimated from hydrostatic conditions,as the sum of pressures imposed by the seawater col-umn and the column of hydrothermal solution belowthe seafloor. Such estimates for hydrothermal systemsyield 350–400 bar for the EPR and 500–650 bar for theMAR.
The residence time of a hydrothermal solution in thezone of the most intense interactions (focus) and in theupwelling limb was estimated as about ten years(Kadko and Moore, 1988; Kim and McMurtry, 1991).According to the data for hydrothermal systems at13° N and 21° N on the EPR, the lifetime of an individ-ual smoker is several decades, and the activity periodsof systems was estimated as n × 100 – n × 1000 y(Lalou et al., 1993). Hydrothermal processes wererepeatedly reactivated in the TAG field over 140 ka(Lalou et al., 1995).
The qualitative chemical structure of the model (i.e.,the set of elements) is defined by the problemsaddressed. The results of simulation presented belowwere obtained for a 15-element multisystem (H–O–R–Na–Ca–Mg–Fe–Al–Si–C–S–Cl–Cu–Zn–Pb) includingall the major elements and some key ore elements. Thegeochemically interesting Mn and Ba could not beincluded because of the lack of necessary thermody-namic data (Section 3.4.2).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S225
Systems involving sedimentary rocks were not con-sidered in this study, which allowed us to significantlysimplify the model, and assume that the crust is com-posed of a homogeneous tholeiitic basalt sequence. Thecomposition of basalts used in our model (Table 17)corresponds to the most abundant statistical group ofthe oceanic crust (Yaroshevsky and Tsekhonya, 1986).The initial rock was regarded in the model as a homo-geneous reactant. The composition of the initial solu-tion was identical to the standard mean ocean water.The model accounted for changes in the bulk composi-tion of the model rock due to metasomatic phenomena.
The key parameter of equilibrium dynamic modelsis the ratio of the rates of solution–rock interaction andmass transfer. In the MSFR method for extendednonisothermal systems with fracture percolation (Sec-tion 2.2), the similarity criterion of the model is theratio of the masses of fresh rock and solution effectivelyparticipating in reactions integrated over the flow lineof solution, ΣR/W. This parameter was estimated by twomethods for active systems: from the concentrations ofmobile elements and from Sr isotope ratios (Tables 11,12), which yielded similar values of 0.5–2.0.
The products of seawater–basalt interaction aremetasomatic mineral assemblages of the chlorite andpropylite facies. They include, respectively, (a) chlorite,mixed-layer chlorite/smectite, quartz, hematite, andanhydrite; and (b) epidote, chlorite, albite, actinolite,and minor quartz and sulfides. Chlorites from the meta-somatic rocks of the downwelling limb are usuallymagnesian and ferromagnesian. The list of mineralsthat should be taken into account in the model wasdefined by the mineralogy of metasomatic rocks andores (Table 3), the elemental composition of the model,and the availability of thermodynamic data.
The role of magmatic fluids in the hydrothermal sys-tems of mid-ocean ridges was evaluated in Section 3.2,and it was demonstrated that it could not be significantin such systems. This factor was therefore ignored inour model.
The upwelling limb of a convection cell in oceanichydrothermal systems is characterized by rapid solu-tion movement and relatively weak interactions withthe country rocks. The character of processes in theupwelling limb may probably be different between sys-tems. At least three extreme variants can be distin-guished:
(1) solution ascent with adiabatic cooling,(2) solution ascent with significant conductive cool-
ing, and(3) subsurface mixing with cold seawater.The high temperatures of black smoker solutions in
many hydrothermal systems directly indicate the real-ization of the first variant. It is certainly the most effi-cient in terms of ore formation on the seafloor. The adi-abatic ascent of solution (if it is not accompanied byboiling) provides only minor cooling by 20–30°C (Bis-choff and Pitzer, 1985). Conductive cooling and mixing
result in the removal of ore metals from the solution anddevelopment of dispersed and vein sulfide mineraliza-tion.
It is concluded that the effects of solution interactionbetween the walls of the channelways and the ascend-ing fluids can be ignored during the simulation of thehydrothermal system as a whole. Only the simulationof metasomatic zoning in the channel walls is of inter-est.
The zone of ore deposition is certainly the mostinteresting for geologists. This part of oceanic hydro-thermal systems has been explored most extensively.The zone of ore deposition includes the ore body properand the area of mixing with cold seawater on its sur-face.10 Ore deposition in modern oceanic hydrothermalsystems is controlled by two factors:
(a) mixing with seawater, which abruptly changesthe temperature and, correspondingly, the migrationability of ore elements, but also, gives rise to chemicalreactions between the mixed components; and
(b) considerable heat loss from the surface of the orebody, which leads to its conductive cooling and maycause metasomatic reactions within the body betweenthe previously deposited material and the new portionsof hydrothermal solution.
10 Hydrothermal plumes are not considered in this paper.
Table 17. Compositions of initial basalt and seawater usedin the model
Note: For the recalculation of basalt composition to element con-centrations, P and Ti were eliminated as apatite and TiO2,respectively; Mn was summed with Mg; and the FeIII/ΣFeratio was taken as 0.15. The composition of seawaterincluded 0.0025 mol/kg of dissolved O2.
S226
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
As was shown above, the relationships of these fac-tors change with time owing to the growth of the edi-fice, which results in the zonal structure of the ore bodyevolving with time. The ore edifice is made up mainlyof anhydrite, amorphous silica, and sulfides (the com-plete list of minerals is given in Table 3).
The temperature of ore deposition varies from 350to 0°C, and most of the material is precipitated at temper-atures above 200°C. Pressure corresponds to the hydro-static value for the given ocean depth, 250–350 bar.According to available measurements, the dischargerates of individual black smokers are 0.n–n.0 kg/s(Table 5). The total discharge of large mounds compris-ing groups of active vents and significant diffuse vent-ing may be an order of magnitude higher. The mixingproportions of hydrothermal solution and seawater varyfrom 0 to 100%, and the temperature of the mixture alsodepends on the mixing proportion. Mixing is a very fastprocess, which can result in the formation of metastablecompounds (for instance, pyrrhotite instead of pyrite insmoke) and nonequilibrium states in some chemicalreactions, especially redox reactions (Janecky and Sey-fried, 1984).
There is an important aspect of the geological modelthat must be clarified: boiling in the focus and theascending channel of the hydrothermal system, and itsinfluence on ore formation. The thermodynamic mod-eling of boiling hydrothermal systems poses a specialproblem, which is examined in Chapter 6.
Thus, the input information for thermodynamicmodeling includes the following characteristics of thegeological model:
—distribution of temperature and pressure along theflow line of the solution;
—initial compositions of seawater and basalt;—the general set of possible metasomatic and ore
minerals; and—bulk rock/water ratios, ΣR/W.The following observations can be used for model
verification:
—mineral assemblages of metasomatized rocks,—compositions of hydrothermal solutions pro-
duced by the hydrothermal system and their depen-dence on the parameters of the system, and
—mineral assemblages of the newly formed ores.The isotopic compositions of ore minerals and solu-
tions can be used as an additional criterion, which isdiscussed in Chapter 5.
3.4.2. Physicochemical model
The physicochemical model that was utilized in thisstudy to describe processes in oceanic hydrothermalsystems belongs to the class of equilibrium dynamicmodels. It makes use of the MSFR method described inSection 2.2. According to this method, the hydrother-mal system is represented by a series of reactors (steps),where an equilibrium thermodynamic state is achievedin local chemical systems in each step and in each stagecomputation. The bulk composition of the system iscontrolled by dynamic relations: transportation of solu-tion between reactors in each stage of computation andinteraction of solution with new portions of fresh rock.The mineral products formed during the previousstages of calculation remain in the reaction zone.11
Thus, the passage of a solution portion (wave) througha step reactor simulates the process of water–rock inter-action along the flow line of the solution. The sequen-tial movement of many portions (waves) of solutionthrough the downwelling and upwelling limbs of con-vection cells and the zone of ore deposition simulatesthe development of the hydrothermal process in time.
Calculations were carried out in the temperaturerange 150–400°C. Lower temperature conditions werenot considered because of the weak influence on the orepotential of the system. The division of the convectionsystem into isothermal steps is shown in Fig. 19. Thetemperature increment was taken as 25°C at 150–200°C and 10°C at 210–400°C. The temperature incre-ment was reduced to 5°C in the upwelling limb and inthe zone of ore deposition till 300°C in order to charac-terize ore deposition occurring in these zones in moredetail. The reactor chain consisted of up to 50 sequen-tial steps in some model variants. The correlationbetween temperature and pressure in each step (Fig. 19)approximated the conditions of solution flow throughthe contact zone of a magma chamber (Fig. 16). Insome model variants, the solution compositionobtained in the 350, 380, and 400°C steps were used asinitial solutions for the upwelling limb of the system.This allowed us, in particular, to estimate the influenceof temperature in the focus of the system on metaltransportation by solutions.
In each stage of calculation, the solution was com-pletely transferred from step to step. The rate of solu-
11 A more complex dynamic system was used for the zone of oredeposition, see Section 4.3.2.
100 200 300 400Temperature, °C
500
400
300
200
Pre
ssure
, bar
Fig. 19. Scheme of a chain of reactors in the T–P coordi-nates.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S227
tion–rock interaction was defined in the model in thesame manner as described in Section 2.2 [Eq. (9)] usingthe similarity criterion (ΣR/W)1. This parameter variedover a wide range from 0.03 to 10, which overlaps itsnatural dispersion (Table 12). The propagation of thezone of metasomatized rocks with time was governedby the diffusion model of Korzhinskii [Eq. (11)]. Thelimiting proportion for the complete basalt alteration[right-hand side of inequality (12)] was taken as 80 kgof rock per one kilogram of solution, which corre-sponds to basalt and hot water densities of 2.9 and≈0.7 g/cm3, respectively, and a fracture porosity of ≈5%(Hyndman and Drury, 1976).
Thermodynamic data. Thermodynamic equilibria inthe system were described within the 15-componentmultisystem H–O–K–Na–Ca–Mg–Fe–Al–Si–C–S–Cl–Cu–Zn–Pb. The model did not include Mn and Ba,
because they do not form their own phases in the down-welling limb but occur as isomorphous components inmetasomatic minerals. There are no reliable thermody-namic data for such compounds, which prevents a reli-able calculation of the behavior of Mn and Ba duringhydrothermal solution formation. The concentration ofMn in the initial chemical analysis of fresh basalt(Table 17) was added to Mg, and that of Ba, to Ca.
The multisystem included 45 fixed-compositionminerals and aqueous solution, which was described by50 dissolved species. The minerals and the solutionspecies are listed in Table 18. Since some rock-formingminerals of metasomatic rocks (actinolite, epidote, andchlorite) are solid solutions, the model included bothend-members and intermediate compositions from thetypical ranges of the objects modeled (Gillis and Rob-inson, 1990). The energies of formation of intermediate
Table 18. Minerals and dissolved species used in the thermodynamic model
Magnetite (Mag) Heulandite (Hul) Pyrite (Py) Na+ FeOH ZnOH+ Na
Albite (Ab) Pumpellyite (Pmp) Troilite (Po) NaCl0 Fe ZnCl+ Ca
Microcline (Mc) Muscovite (Ms) Galena (Gn) NaH Fe Zn Mg
Clinochlore (Cch) Paragonite (Pg) Sphalerite (Sph) Ca++ Al Zn H2S0
Daphnite (Dph) Diaspore (Dsp) Chalcocite (Cc) CaCl+ Al Zn HS–
Tremolite (Tr) Kaolinite (Kln) Chalcopyrite (Ccp) Ca Al CuCl0 Cl–
Fe-tremolite (Fe-Tr) Pyrophyllite (Prl) Bornite (Bn) Mg++ Cu HCl0
Epidote (Ep) Talc (Tlc) Sulfur (Su) Cu
Intermediate minerals
Chlorite–50 (Cch : Dph 1 : 1) (Chl50)
Chlorite–75 (Cch : Dph 1 : 3) (Chl75)
Actinolite–80 (Tr : Fe-Tr 4 : 1) (Act80)
Epidote–75 (Ep : Zo 3 : 1) (Ep75)
Epidote–60 (Ep : Zo 3 : 2) (Ep60)
Sericite (Ms : Pg 1 : 1) (Ser)
Note: Symbols of minerals used in this paper are given in parentheses. Quartz was changed to amorphous silica (ASi) during the simulationof solution cooling.
H20
Cl20 CH4
0
Cl3– O2
0
Cl20 Cl4
–– SO4––
HSO4–
Cl2– SO4
–
OH( )2+ SO4
0
CO30 OH( )3
0 Cl20 SO4
0
OH( )2+ Cl3
–
OH( )30 Cl4
––
Cl20 OH( )4
–
H2CO30 Cl2
–
H2SiO40 HCO3
– Cl3–
S228
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
chlorites, actinolite, and sericite were calculatedassuming ideal multisite mixing, and the model of Birdand Helgeson (1980) was used for epidote. The modelaccounted for the most important dissolved species ofthe elements: free ions and chloride and hydroxo com-plexes. Sulfate and hydrocarbonate complexes of themajor components were also calculated. Preliminarycalculations (Grichuk et al., 1985) showed that underthe model conditions, sulfate, hydrosulfide, andhydroxo–carbonate complexes of ore elements arenever predominant, which allowed us to omit these spe-cies from the model. The set of complex compoundsused in our model is rather typical of such modelingexercises with hydrothermal processes (Rafal’sky,1993).
The set of minerals considered (Table 18) providessome simplifications to the physicochemical model incomparison to the geological model, which must betaken into account during interpretation. The mostimportant among them are the following.
(a) The multisystem does not include smectitesbecause of the lack of reliable thermodynamic data.Hence, the mobility of Al can be overestimated and thepH may be too alkalic in the low-temperature part of thesystem (150–200°C).
(b) Isomorphic substitution of Zn and Pb in alumi-nosilicates is ignored. This results in the overestimatedmobilization of ore elements from metasomatic rocksin the absence of Zn and Pb sulfide minerals.
Thermodynamic data (gT) for equilibrium calculationsin the multisystem were taken from the UNITHERM databank (Geochemistry Department, Moscow State Uni-versity). The gT values of simple (base) ions were cal-culated in the data bank using the modified Helgeson–Kirkham–Flowers model (Johnson et al., 1992), andthose of complex ions were extrapolated by theRyzhenko equation. The extrapolation of the propertiesof complex ions of ore elements was based on the fol-lowing reference experimental data: Plyasunov andIvanov (1990) for Zn, Var’yash (1991) for Cu, andSeward (1984) for Pb. The thermodynamic characteris-tics of H2S and complex Fe species were additionallyrefined. The data used to derive the thermodynamicproperties of substances compiled in the UNITHERMdata bank were described by Borisov (2003).
The activity coefficients of dissolved species werecalculated by the third approximation of the Debye–Hückel equation. The values of å were taken after Gar-rels and Christ (1965), and those for complex ionsabsent in this compilation were taken as 4.5 (Rafal’sky,1993). The temperature dependence of the coefficientof the linear term was taken from Helgeson andKirkham (1974).
Quality control of thermodynamic calculations. Theerrors of thermodynamic calculations related to theaccepted characteristics of substances must be dis-cussed. The determination of these errors is one of themost difficult problems in the modern thermodynamics
of geochemical processes, and it is not yet rigorouslysolved. Various approaches were discussed by Karpovet al. (1976) and Dorogokupetz and Karpov (1984).
The UNITHERM data bank does not include globalprocedures for the derivation of internally consistentdata, which are implemented in the DIANIK andTHERMINEOS data banks (e.g., Dorogokupetz andKarpov, 1984). Thermodynamic parameters fromUNITHERM must be preliminary refined using someparticular systems. It is known that experimental dataand, correspondingly, thermodynamic constantsderived from them involve random and systematicerrors. The confidence limits provided in publicationsusually characterize only random experimental errors.A combined processing of many experimental investi-gations may reveal systematic errors in particular dataseries. If the experimental information is limited, sys-tematic deviations are possible. Their influence on theresults of calculations in multicomponent systems can-not be determined from the confidence intervals.
The overall quality of calculations can be assessedfor particular (with certain sets of elements and condi-tions) types of systems by the simulation of experi-ments in complex systems. A number of reliable exper-imental studies have been reported for multicomponentsystems approaching complex geologically relevantsystems. The results of these experiments cannot com-monly be used to derive fundamental thermodynamicconstants (formation constants of complexes and freeenergies of compounds). However, they can be appliedfor the integrated assessment of thermodynamic datasets obtained in simpler systems. This method was usedin a number of studies (e.g., Rafal’sky, 1993).
For the problems related to hydrothermal ore forma-tion, the experimental study by Hemley et al. (1992) isof special significance in this respect. These authorsstudied the combined solubility of the most importantsulfides of Cu, Zn, and Pb in chloride solutions in equi-librium with the microcline–muscovite–quartz andpyrite–pyrrhotite–magnetite buffer associations. Animportant point is that the experiments of Hemley et al.(1992) reached equilibrium. Experiments in the basalt–seawater system are chemically closer to the subject ofour study. Unfortunately, they are not suitable for theassessment of the quality of equilibrium calculations,because equilibrium states are not reached in such sys-tems even in very long experiments (up to two years).
Using the UNITHERM database, we calculated amodel reproducing Hemley et al.’s (1992) experiments.Figure 20 compares the results of these experimentsand calculations for ore metal concentrations at 300–400°C and a pressure of 500 bar. It can be seen thatthere is general agreement between the calculated andexperimental parameters, although there is a small sys-tematic error for all metals, positive for Fe and Zn andnegative for Cu and Pb. In this case the mean squaredeviation for the total concentrations of ore metals is±0.26 logarithmic units. This value can be regarded as
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S229
a general approximate estimate of the error of equilib-rium calculations for the minor components of the sys-tem. The concentrations of major components aredefined by mass balance equations, and the quality ofcalculation depends on the accuracy of the bulk compo-sitions of initial solutions and rocks. The relative errorsin these values are no higher than 20%.
The uncertainty of this estimate of the accuracy ofthermodynamic calculations is related primarily to thefact that the calculated model assemblages of alumino-silicate minerals are more complex than those in Hem-ley et al.’s (1992) experiments. Unfortunately, there areno reliable experimental data on the composition ofequilibrium solutions in systems containing coexistingepidote, actinolite, chlorite, albite, sulfides, and otheraluminosilicates present in the model multisystem(Table 18). However, the experiments that were usedfor testing were carried out under temperature, pres-sure, pH (pHT = 4.5), and chloride contents similar tothose of oceanic hydrothermal solutions. Hence, thesolubilities of ore metal sulfides in the experimentalsolutions are also close to those observed in the ocean.
The test performed here does not guarantee the reli-ability of calculations for individual species, becauseowing to the competitive character of complexation, thecorresponding errors can be mutually compensated inthe total solubility of elements. This compensation sug-gests that any extreme extrapolation of the estimationof accuracy is dangerous, especially with respect totemperature. However, Hemley et al.’s experimentswere performed in the same range of conditions as ourcalculations (300–400°C and 500 bar as compared with150–400°C and Psat. vap = 500 bar in model calcula-tions). The estimated accuracy of total concentrationswill be little changed if proper allowance for compen-sated errors is made.
Thus, the total errors of the equilibrium calculationare estimated as ±0.1 logarithmic units for the concen-trations of major components and ±0.5 logarithmicunits for trace components.
3.4.3. Software for thermodynamic modeling
Thermodynamic calculations relevant to the prob-lem addressed in this study were performed from 1980–1990s using the GIBBS program developed byYu.V. Shvarov at the Geochemistry Department ofMoscow State University. The results were reported byGrichuk and Borisov (1983), Grichuk et al. (1985), Gri-chuk (1988), and Hydrothermal Sulfide… (1992). TheGIBBS program was designed for equilibrium calcula-tions in chemical systems of an arbitrary phase com-plexity (Methods of…, 1988) and was implemented inhigh-performance computers of the Minsk and ESseries. The GIBBS program was not intended for modelcalculations with dynamic scenarios.
In 1991–1992, together with M.Yu. Korotaev, theauthor developed the GBFLOW program for the calcu-
lation of equilibrium dynamic models by the method ofa step flow reactor on IBM-compatible personal com-puters. All the model calculations presented henceforthwere performed by the author using this program. TheGBFLOW program includes a block for the minimiza-tion of the Gibbs free energy of the system, which waspreviously created by Shvarov as part of a simplifiedversion of the GIBBS program for small computers.12
Subsequently, GBFLOW was optimized by the authorfor the calculation of models using the MSFR method,which increased its performance rate in such problemsby approximately one order of magnitude. TheGBFLOW v.3.1 (1995) used for calculations in thisstudy includes a block for isotopic chemical modeling(Chapter 5), and the format of its output files is orientedfor the subsequent graphical processing using the MSExcel package. On the basis of the GBFLOW program,the author developed the following versions for the cal-culation of models with special dynamic scenarios:
—GRDEP, for the simulation of a growing ore body(Section 4.3.2); since such models require considerablecomputation time, GRDEP is also compatible with theUNIX platform; and
—PENG, for the calculation of boiling systems(Chapter 6).
The precision of computational programs can beestimated by comparing the results obtained by severalprograms for the same problem. Such a test forGBFLOW was performed by Borisov and colleaguesusing the GIBBS program, which, in turn, was com-pared with the SELECTOR program developed at theSiberian Institute of Geochemistry (e.g., Borisov andKhodakovsky, 1989). The programs yielded identicalresults for thermodynamic equilibria within thedeclared calculation accuracies (GBFLOW: 0.001% for
12 This block allows the calculation of systems with partially deter-mined phase compositions, where the presence of an aqueoussolution is mandatory.
300275 325 350 375 400 425Temperature, °C
–4.5
–4.0
–3.5
–3.0
–2.5
–2.0
–1.5
–1.0logMe [mol/kg]
Experiment
Calculation
Fe
Fe
Cu
Cu
Zn
Zn
Pb
Pb
Fig. 20. Comparison of model results with experiments onthe dissolution of buffer associations (Hemley et al., 1992).Experimental conditions: T = 300–400°C, P = 500 bar,1 mol NaCl, and excess Py + Po + Mag + Kfs + Ms + Qtz +Ccp + Sp + Gn.
S230
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
150 220 260 300 340 380
(a)
Temperature, °C
(c)
1
0
–1
–2
–3lo
g(R
/W)
100
(b)80
60
40
20
0
Act80Tr
Dph
Chl75
Chl50
Cch
Tlc
Pmp
Ep60
Ep75
Ep
Anh
Ab
Qtz
Hem
Py
Act80
Chl75
Ep60Ep75
Ep
Ab
Pmp
Chl50
Cch
Anh
Tlc
Qtz
Hem Py
Dph
%m
ol/
kg
(d)
0.06
0.04
0.02
0
(e)
K
Ca
Mg
Fe
Si
S(VI)
S(II)
log
Me
–3
–4
–5
–6
–7
–8
Cu
Zn
Pb
(f)
log(H
2)
pH
7
6
5
4
0–2–4–6–8
–10
Fig. 21. Results of the simulation of seawater–basalt interaction in the downwelling limb of a hydrothermal system for the modelcase with total R/W(400°C) = 1.26 and wave no. 1. (a) Cumulative R/W, (b) mineral assemblages, (c) major components of solution,(d) heavy metals in the solution, (e) pH, and (f) equilibrium hydrogen activity.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S231
the concentration of a substance, if the abundances ofits elements are higher than 10–7 mol/kg H2O).
The thermodynamic data bank UNITHERM wasused in this study. It was created at the GeochemistryDepartment, Moscow State University by Shvarov withthe participation of Borisov and Grichuk. The process-ing of experimental data and derivation of the thermo-dynamic characteristics of dissolved species were car-ried out by the author using the UT-HEL and UT-RYZprograms designed by Shvarov.
CHAPTER 4. SIMULATION OF GEOCHEMICAL PROCESSES IN THE HYDROTHERMAL SYSTEM
OF A MID-OCEAN RIDGE
4.1. Results of Simulation for the Downwelling Limb of a Convection System
4.1.1. Model of a short-lived hydrothermal system
Let us consider first the results of calculation forseawater interaction with fresh basalt, i.e., a model vari-ant (first wave) corresponding to the initial stages ofhydrothermal activity or short-lived systems. This cal-culation is illustrated by Fig. 21.
Metasomatic mineral assemblages are controlled inthe model by temperature and the amount of rock react-ing with solution. At low R/W and low T, basalts arealtered to metasomatic mineral assemblage I: quartz +chlorite + hematite + anhydrite ± kaolinite ± talc. Mg isremoved from the solution to form chlorite, and ironand sulfur are oxidized in the basalts by seawater-dis-solved oxygen. Most of the basaltic components(except for Al, Fe, Si, and Mg) are extracted by thesolution. The migration of Ca from the rock to watercauses anhydrite precipitation.
An increase in T and related intensification of solu-tion–rock interaction (i.e., an increase in R/W, Fig. 21a)change the metasomatic association. In the case withT > 230°C and R/W > 0.027 (Fig. 21b), the basalt isaltered to assemblage II: chlorite + epidote + actino-lite ± quartz ± sulfides. Albite is added to the assem-blage at R/W > 0.06. Pumpellyite is also stable in thisassociation up to temperatures of 260–290°C. In theassemblage II field, the solution is reduced and containssignificant amounts of H2S and H2. The development ofassemblages I and II is a persistent feature of the modeland, as will be shown below, a change from one assem-blage to the other affects the composition of solutionand the behavior of ore elements.
Since the transition from association I to associationII is of particular importance, it should be determinedwhether temperature or R/W value is the major control-ling factor. To this end, we performed a series of calcu-lations, in which varying ΣR/W values were reached(from 0.02 to 12 at a temperature of 400°C). In theΣR/W–T coordinates, the calculated trends are subpar-allel (Fig. 22a) and the points of reactor steps with iden-tical mineral assemblages form separate areas. It can beclearly seen that the boundary between the fields of
assemblages I and II, which is marked by the appear-ance of epidote, is controlled primarily by R/W and lieswithin the interval 0.02–0.03. The position of thisboundary shows only a very weak dependence on tem-perature. It is caused by two reasons: (a) changes in thethermodynamic properties of the reactants (pumpelly-
20–2–4log(ΣR/W)
400
350
300
250
200
150
1
2
Ab + Chl + Ep ± Qtz + sulfides
Anh + Hem +Chl + Qtz
Py
Po
Sph
Bn, Ccp
210–1–2–3–4
(a)
(b)
(c)
log[R/W]
400
350
300
250
200
150
100
Hematite
Tremolite
Anhydrite
210–1–2–3–4
Albite
Wairakite
Pumpellyite
400
350
300
250
200
150
100
Tem
per
ature
, °C
(without epidote)
log[R/W]
Fig. 22. Stability fields of mineral assemblages and someindicator minerals in the downwelling limb of a hydrother-mal system (wave no. 1). (a) Stability fields of assemblagesI (Anh + Hem + Chl + Qtz) and II (Ab + Chl + Ep ± Qtz ±sulfides) and sulfide minerals with (1) solution flow linesand (2) the boundary between associations I and II. (b) Sta-bility fields of hematite, anhydrite, and tremolite (withoutepidote). (c) Stability fields of albite, wairakite, andpumpellyite.
I
II
S232
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
ite and quartz occur in association II at low tempera-ture, whereas pumpellyite is absent at high tempera-tures; Fig. 22c); (b) the previous history of the solution:solutions attaining ΣR/W ≈ 0.03 at high temperatureslose a significant amount of anhydrite during heating.
A detailed analysis of the results of the simulationshowed that assemblage I changes to assemblage II notdirectly but through transitional mineral associations(Fig. 22b), which consist of various combinations oftremolite, chlorite, quartz, hematite, and some alumino-silicates observed only in this transitional zone (sericiteand pyrophyllite). The real existence of these transi-tional assemblages is not quite clear. The fractions ofthese aluminosilicates are not large, and their appear-ance could be related to the fact that the model ignoredaluminous end-members of rock-forming actinolite andchlorite. When albite and pumpellyite appear in associ-ation II, these phases disappear and the R/W rangewhere the transitional assemblages are stable is onlyfrom 0.025 to 0.04 (Figs. 22b, 22c). These transitionalassociations have no particular effect on the calculatedsolution composition.
Within the downwelling limb, sulfide minerals arestable only in the assemblage II zone (Fig. 22a). Sul-fides are dissolved and partly oxidized at low R/W. Thestability conditions of sulfide minerals are different inthe zone of assemblage II. The smallest field is charac-teristic of iron sulfides: pyrite is stable only up to a tem-perature of 250–270°C, and pyrrhotite is present in themineral association at higher temperature and R/W > 0.3.At low temperatures sphalerite and galena are stable atR/W > 0.03, but when temperatures increases above300°C, their fields are reduced and shift to higher R/W.The zone of assemblage II closely corresponds to thestability field of copper sulfides. At high R/W dissemi-nated chalcopyrite forms in the rock, and bornite occursnear the boundary. In the transitional assemblages, cop-per is usually incorporated into chalcocite. These fea-tures suggest that sulfide minerals can occur in metaso-matized basalts only in reduced environments. The dif-ferences between their stability fields are relatedprimarily to the different temperature dependencies ofsulfide solubilities. The small stability field of pyrrho-tite, which is the main repository of sulfur in freshbasalts, indicates rapid mobilization of magmatic sulfurunder the influence of hydrothermal solutions, whichwas in fact observed in ocean floor basalts (Gitlin,1985).
Major components of hydrothermal solutions. Thecompositions of hydrothermal solutions equilibratedwith mineral assemblage I bear some resemblance totheir source, seawater. They show near neutral acid–base and redox characteristics, the concentration of Mg
in them is higher than those of Ca and K, and isthe second most abundant anion after Cl–. In the stabil-ity zone of assemblage II, the chemical compositions ofhydrothermal solutions change dramatically. In thehigh-temperature part of the system, they are essen-
SO42–
tially free of Mg and and enriched in Ca, K, Si,Fe, and H2S (Fig. 21c). Some representative composi-tions of hydrothermal solutions obtained in variousmodel variants are shown in Table 19. These results are,in general, consistent with the previously publishedversions of the model (Grichuk et al., 1985; Hydrother-mal Sulfide…, 1992).
The components of a solution can be divided intofour groups on the basis of their behavior in the system:(1) those occurring mainly in the liquid phase: Cl, Na,K, and H2O; (2) components occurring mainly in thesolid phase and whose behavior is controlled by the sol-ubility of rock-forming metasomatic minerals: Mg, Ca,Fe, Al, Si, H+, and OH–; (3) components whose behav-ior is controlled by redox reactions: C, S, and H2; and(4) chalcophile trace elements, which are controlled bythe sulfur regime: Cu, Zn, and Pb.
Chlorine behaves conservatively in the solutions,and its concentration only slightly rises at high R/Wowing to water fixation in the rock. The main chlorinespecies is Cl–; of lesser importance are the complex
species NaCl0, Ca , and, at high temperatures, HCl0.
During water–basalt interaction under the stabilityconditions of assemblage I, sodium is completelyextracted into solution. In the moderate-temperaturetransitional region, the concentration of Na ceases toincrease when sericite appears in the system and albitecrystallizes after R/W ≈ 0.6 is reached (at any T). Athigher R/W values, it gradually decreases, mainly at theexpense of Ca–Na exchange (Table 19, solution nos. 5–
7). The main dissolved Na species are Na+, NaS , andNaCl0.
Almost within the whole range of conditions, potas-sium is extracted from the solid phase and is rapidly accu-mulated in solution. Its concentration is limited at high Tand R/W > 2 by the formation of microcline (Table 19,solution nos. 6 and 7). Potassium is also partly retained inthe transitional zone in sericite. Its predominant dissolvedspecies is K+ at low temperature, and KCl0 occurs in sim-ilar concentrations at high temperatures.
Water is captured by the secondary mineral products ofsolution–basalt interactions. During the formation of meta-somatic assemblage I, the water content is up to 8–12% ofthe mass of metasomatic rock. The formation of metaso-matic assemblage II is accompanied by much less exten-sive hydration, about 2.5–3.0 wt %, because epidote andactinolite occurring in this assemblage are relatively poorin hydroxyl groups. However, during interaction with alarge amount of rock (about 34 kg per one kilogram ofsolution), the solution can be completely exhausted.13 Thebinding of water is about 3% in oceanic hydrothermalsystems with ordinary R/W values of 0.5–2.0.
13 This variant is not realized in actively discharging hydrothermalsystems.
SO42–
Cl20
O4–
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S233
Magnesium is intensely removed from solutionwithin the whole range of model conditions. In the sta-bility field of assemblage I, Mg is incorporated in chlo-rite (clinochlore) and talc; and into the fields of assem-blage II and transitional assemblages, it is incorporatedinto actinolite. Its residual concentrations established inequilibrium with mineral assemblage II are 10–5–10−7 mol/kg (in contrast to 53.2 mmol/kg in seawater).The decrease in magnesium content in the solution isapproximately proportional to an increase in R/W. The
main dissolved magnesium species is Mg2+, and MgSalso occurs at low R/W.
Calcium behavior in the hydrothermal system ismore complicated, and its concentration forms a rathercomplex pattern in the ΣR/W–T coordinates (Fig. 23a).At low R/W values, calcium is extracted from basaltsand magnesium substituted for it in metasomatic min-erals. The extracted calcium immediately precipitatesfrom the solution as anhydrite, which results in a slightdecrease in the residual concentration of calcium in thesolution (Fig. 21c). This decrease is more pronouncedat high temperatures, because the solubility of anhy-drite is negatively correlated with temperature. Sincethe amount of sulfates in hydrothermal solution is lim-ited (seawater contains about 28.2 mmol/kg of sul-
O40
fates), the concentrations of Ca and becomeequal at R/W > 0.01, and then the concentration of Cabegins to increase. It reaches a local maximum in equi-librium with transitional assemblages (marked by thepoint of anhydrite disappearance), after which itdecreases until the appearance of albite in the mineralassemblage. From that moment the concentration ofcalcium increases gradually again, owing mainly to its
complexation and formation of Ca at high temper-atures. This species is predominant in the high-temper-ature part of the system, whereas the low-temperature
speciation includes major Ca2+ and minor CaS andCaCl+. At temperatures of 370–400°C, the concentra-tion of Ca in the hydrothermal solution is 3–8 timeshigher than that of the initial seawater.14
Almost within the whole range of concentrations,silica is gradually extracted from the basalts andaccumulated in the solution. Its concentrationincreases monotonously with temperature reaching 10–
14 These calculations are in agreement with the analyses of naturalsolutions and are strongly different from the results obtained inthe early versions of the model (Grichuk et al., 1985) and the dataof Bowers and Taylor (1985), where such a resemblance was notattained (see Section 3.3).
SO42–
Cl20
O40
Table 19. Composition of ore-forming solutions calculated in different model cases
Solution component Unit
Conditions and results of model calculations
T = 300°C,P = 500 bar,
R/W = 0.0169,first wave
T = 370°C,P = 500 bar,
R/W = 0.0244,first wave
T = 370°C,P = 500 bar,R/W = 0.244,
first wave
T = 370°C,P = 500 bar,
R/W = 0.0335,wave no. 53
T = 370°C,P = 500 bar,R/W = 0.732,
first wave
T = 370°C,P = 500 bar,R/W = 2.439,
first wave
T = 400°C,P = 400 bar,R/W = 2.967,second wave
Solution no. 1 2 3 4 5 6 7
Na mmol/kg 481.4 482.5 478.4 470.5 441.7 363.2 267.8
K " 10.6 10.8 19.22 7.03 37.9 82.6 72.5
Ca " 18.7 26.5 23.22 33.29 32.1 48.8 84.9
Mg 7.19 0.0005 0 0.0004 0 0 0
Fe " 3.53 2.06 1.88 1.71 2.53 2.69 15.7
Zn μmol/kg 18.6 26.8 115.9 18.5 57.3 58.4 470.0
Cu " 19.6 28.3 4.77 35.1 1.41 1.06 0.76
Pb " 0.30 0.39 0.69 0.30 0.35 0.37 0.185
SiO2 mmol/kg 10.9 15.2 12.76 14.4 14.2 15.1 10.51
Cl " 545.9 545.9 545.9 545.9 545.9 545.9 545.9
H2S " 0 0.298 4.73 0.688 12.4 19.2 22.16
SO4 " 1.73 1.82 0 0 0 0 0
pHT 4.02 5.63 6.02 5.69 5.94 5.79 6.67
pH25°C 6.09 5.60 4.05 4.20 3.51 3.63 3.63
Phase assemblage
Qtz Hem CchChl50 TlcAnh
Ab Ep Chl50Chl75 Act80
Ab Chl50Act80 Ep75Ep60 Sp Bn
Ab Chl50Chl75 Act80Ep Ep75 Bn
Ab Chl75Dph Act80Ep75 Ep60Gn Sp Ccp
Ab Mc Chl75Dph Act80Ep60 Wai PoGn Sp Ccp
Ab Mc Chl75Act80 Ep75Ep60 Po GnSp Ccp
S234
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
15 mmol/kg in the hot part of the system. Note that athigh R/W, quartz is absent in the mineral assemblage,and the hydrothermal solutions are not quartz-satu-rated. Similar conclusions were derived by Wells andGhiorso (1991) on the basis of experimental data andinvestigation of natural prototypes.
Aluminum shows inert behavior in hydrothermalsystems: it is tightly bound in aluminosilicates (chlo-rite, epidote, and albite) and is not extracted by solu-
tion. Its concentration in the low-temperature part of
the system is n × 10–6 mol/kg, and Al(OH is the mostabundant species. In the high-temperature part of thesystem, the concentration increases up to n × 10–5, andAl(OH is prevalent in the solution.
Considering the basalt–seawater system as a whole,iron is characterized by a relatively low mobility, whichincreases with increasing temperature owing to more
)30
)4–
150
logCa(a)
200
250
300
350
400
T, °C
–2.0
150
logSVI (b)
200
250
300
350
400
T, °C
–3
150
–3 –2 –1 0 1
logFe(Ò)
200
250
300
350
400
T, °C
–3.5
log(R/W)
–1.8
–1.6
–1.6
–1.8
–1.8
–1.6
–1.4
–1.2
–4 –
5
–6
–10
–2
–4.0–
5.0
–6.
0
–3.
0
–2.5
–3.0
–3.0
–2.5
–3.5
–3.5
–4.0
–4.0
Fig. 23. Concentrations of components in solution as a function of temperature and rock/water ratio. (a) Ca, (b) SVI, (c) Fe, (d) SII,and (e) Fe/SII.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S235
extensive complexation, primarily with the chlorideion. On the other hand, iron behaves differently in equi-librium with mineral assemblages I and II (Fig. 23b). Atlow R/W, when the main iron repositories are hematite(FeIII) and ferromagnesian chlorite (FeII) (Figs. 21b,22a), the solution is gradually enriched in iron (up to0.3 mmol/kg at 200°C and ΣR/W = 0.015 and3.5 mmol/kg at 300°C and ΣR/W = 0.017; Table 19, no. 1).In the stability field of assemblage II, in equilibriumwith iron-rich chlorite, epidote, and actinolite, iron con-tent in the solution decreases abruptly and iron remainspractically immobile at temperatures up to 280°C. Theconcentration of Fe in the solution increases only attemperatures higher than 350°C in response to moreextensive complexation and reaches 3–12 mmol/kgunder near critical conditions (Table 19, nos. 5–7;Figs. 21c, 23b). Within the whole range of model con-ditions, FeII prevails in solution; the major species are
Fe2+ and FeOH at low temperatures, and FeOH
and FeOHCl0 at high temperatures (Fig. 24a). The con-centrations of FeIII complexes are negligible within thewhole range of model conditions.
Cl2– Cl2
–
The behavior of sulfur in the solution is complicatedby the existence of two valence states with high mobil-ities and strongly different properties, sulfate andhydrogen sulfide. The initial seawater is free of H2S and
contains 28 mmol/kg . The fate of the sulfate ionis controlled in the low-temperature part of the systemby two factors: (a) with increasing temperature, seawa-ter reaches saturation with respect to anhydrite, whosesolubility shows a negative temperature dependency(up to ≈300°C), and (b) Ca extraction from the basaltduring its interaction with the solution enhances anhy-drite deposition. The combined effect of these two fac-tors scavenges sulfate species from the solution: in thecase shown in Fig. 23c for T = 220°C and R/W = 0.02,only 1.07 mmol/kg of sulfate is retained in the solution,which is 4% of its initial concentration. In the cases thatreach R/W = 0.02 at higher temperatures, the concentra-
tions of are even lower (Fig. 23c). The SII of theinitial basalt is completely extracted into solution and
oxidized to in this part of the downwelling limb.The equilibrium concentration of H2S in the solution isno higher than 10–5 mmol/kg.
SO42–
SO42–
SO42–
150
logSII(d)
200
250
300
350
400
T, °C
150
–3 –2 –1 0 1
logFe/SII(e)
200
250
300
350
400
T, °C
10
log(R/W)
–4
–4
–4
–10
–6–8
10
12
12
8
8 6 42
1
0
Fig. 23. (Contd.)
S236
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
A sharp change in the behavior of sulfur occurs atthe transition to assemblage II. When the buffer associ-ation epidote–iron-rich chlorite is produced from basal-
tic material, is reduced, anhydrite disappearsfrom the solid phase assemblage (Fig. 22b), and the
concentration of in the solution falls to almostzero. The concentration of H2S increases (Fig. 23d) atthe expense of both sulfate reduction and enhanced SII
mobilization from the basalt with increasing R/W. Thesolution becomes saturated with respect to initial mag-
SO42–
SO42–
matic sulfides at T > 270°C and R/W > 1, when FeS(model analog of pyrrhotite) appears in the mineralassociation (Fig. 22a). The concentration of H2S in thesolution increases up to 10–20 mmol/kg (Table 19),mainly at the expense of rock sulfur, which is suggestedby the isotopic model (Chapter 5). As a result, in con-trast to Fe, whose concentration in the solution dependson temperature only within the stability field of assem-blage II, the concentration of SII also varies with ΣR/W(Fig. 23d). The relative amounts of sulfide sulfur andiron in the solution change depending on T and ΣR/W
150 220 260 300 340 380
(a) Iron
Temperature, °C
(b) Sulfur
(c) Copper
(d) Zinc
(e) Lead
log(m
)
–2
–4
–6
–8
–10
0
–2
–4
–6
–8
–10
–12
Fe++
FeOH+
FeCl+
FeCl2FeOHCl
FeOHCl2–
SO4– –
HSO4–
NaSO4–
CaSO4MgSO4H2S
HS–
–4
–6
–8
–10
CuCl
CuCl2–
CuCl3– –
–3
–5
–7
–9
Zn++
ZnOH+
ZnCl+
ZnCl2ZnCl3
–
ZnCl4– –
–8
–6
PbCl+
PbCl2PbCl3
–
PbCl4– –
–10
Fig. 24. Speciation of ore elements in solution for the model of the downwelling limb with the cumulative R/W = 1.26 (wave no. 1)corresponding to Fig. 21.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S237
(Fig. 23e): iron prevails over sulfur at low R/W and highT (>370°C). As will be shown below, these features ofsulfur behavior are reflected in the histories of other oreelements.
Within the whole range of model conditions, SII
migrates mainly as H2 . SIV occurs mainly as
Mg in the beginning of the downwelling limb ofthe convection cell, when the solution is similar to sea-water. With increasing Mg bonding in the rock, the
fractions of and NaS rise (Fig. 24), and HSbecomes the major species at high temperatures.
Carbon barely forms mineral phases within therange of model conditions. Only at very high R/W (>1)and low T was calcite obtained in the model mineralassemblage. Within the stability fields of mineralassemblage I and transitional assemblages, the main
carbon species is H2C . Its content is 2 ± 1 mmol/kgwithin the whole range of conditions. The reduction ofthe system results in the appearance of methane in thesolution. Its concentration reaches n mmol/kg at highR/W. The initial model basalt composition included asmall amount of carbon (0.0088 mol/kg, Table 17), andthe bulk C content of the solution increases during itsinteraction with rock. Although carbon was taken asCO2 for the calculation of the redox state of the basalts,its concentration increases mainly at the expense ofmethane. This is not fully consistent with the data onoceanic hydrothermal systems, where CO2 is moreabundant than CH4 (Tables 6, 9).
Acid–base conditions (pH). The downwelling limbof the system generally has near neutral pHT values,within 2 units of the neutral value for the given temper-ature. There is a distinct tendency for pHT to increasewith increasing R/W. Solutions in contact with mineralassemblage I have pHT 4–5, whereas those equilibratedwith assemblage II show higher values from 5 to 7(Table 19). There is an abrupt pH increase by approxi-mately 1.0 when one assemblage changes to the other(Fig. 21e). In the stability region of assemblage I, thepH of the solution is defined by the precipitation of Mgminerals (talc and chlorite) and is therefore controlledmainly by the solution composition, primarily by itsMg content.
At high R/W the acidity of the solution is buffered bythe reactions of mineral assemblage II (epidote +albite + chlorite + actinolite) with Na and Ca in thesolution. A numerical experiment with the addition of astrong acid (HCl) into the 370°C reactor step (Table 20a)showed that the metasomatized rock is an efficient pHbuffer. The addition of 20 mmol of acid to the systemchanged the pH by only 0.1 (addition of the sameamount of acid to seawater shifts pH from 8 to about 2at 25°C). Simultaneously, a significant amount of Ca iswithdrawn into the solution, Na is removed from thesolution, the concentration of the clinozoisite end-
Saq0
SO40
SO42– O4
– O4–
O30
member in epidote increases (Ep60), epidote is partlydissolved, and the amounts of iron-rich chlorite andalbite increase (Table 20a). This reaction can be sim-plified as
(17)
This equation is similar to the reaction derived bySeyfried et al. (1988) from the results of experiments inthe water–basalt system.
In order to compare these results with observationsin the natural analogs, the pHT values need to be recal-culated to pH25. To this end, the following numericalexperiment was performed: the solutions obtained forthe downwelling limb were recalculated to equilibriumat 25°C without the formation of solid phases (exceptfor sulfide and amorphous silica) and carbon dioxidereduction to methane. This experiment closely simu-lates quenching of natural hydrothermal solutions dur-ing sampling.15 The recalculated pH25 values are sys-tematically lower than pHT, and the difference increaseswith increasing temperature and decreasing pressure ofthe hydrothermal solution, reaching 2–3 pH units. As aresult, the pH25 values of solutions obtained at ΣR/W > 0.2vary from 3.5 to 4.05 (Table 19), which is in agreementwith natural observations (2.7–4.5, Table 6). The lowestpH25 values of about 2 were calculated for the solutionsin equilibrium with assemblage II at a temperature of425°C and a pressure of 400 bar; their pHT values wereabout 6.
The change in pH during cooling is related, first, tothe balance of hydrolytic reactions between solutioncomponents: the hydroxo complexes of Al and Fe, on
one hand, and the undissociated acid species H2Cand H2S0, on the other hand; and, second, sulfide pre-cipitation associating with an increase in solution acid-ity (this effect was discussed in detail by Rafal’sky,1993):
Fe + H2S FeS + 2H+ + 2Cl–.
If R/W is low and T is high, the dissociation of
HS , whose concentration may be as high asn mmol/kg under such conditions, also contributes sig-nificantly to the pH change. At high T and R/W andlow P, the aforementioned strong pH changes are due toHCl0 dissociation.
Redox conditions and dissolved hydrogen. Theredox conditions of the downwelling limb are definedby the proportion of valence forms of iron, sulfur, car-bon, and hydrogen. In the multisystem modeled, coppercan also occur in several valence states, but its content
15 During sampling of high-temperature smoker solutions for pHmeasurement, it is not possible to prevent the generation of sul-fide suspension (Trefry et al., 1994).
12Ep 35H+ 6H2 9Na++ + +
Zo 4Dph 9Ab 22Ca2+ 17H2O.+ + + +
O30
Cl20
O4–
S238
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Table 20. Results of a numerical experiment on the estimation of the buffer capacity of mineral assemblage II for the initialcomposition of the system corresponding to the case shown in Table 19 (no. 5): 370°C step, P = 500 bar, ΣR/W = 0.732, andthe first wave(a) Influence of strong acid addition
Parameter Initialcomposition
HCl addition, mmol
1 3 10 20
Solution components
pH 5.936 5.931 5.921 5.884 5.834
Na, mmol 441.7 441.6 441.4 440.0 436.7
Ca, mmol 32.1 32.6 33.7 37.8 44.3
–log(H2) 1.750 1.754 1.761 1.778 1.787
Minerals, mmol
Ab 115.2 115.3 115.4 116.9 120.2
Chl75 19.4 19.2 18.6 18.3 20.3
Dph 5.12 5.46 6.11 6.94 5.93
Act80 56.4 56.5 56.7 56.8 56.1
Ep75 17.2 14.5 9.3 – –
Ep60 57.1 59.4 63.9 71.1 68.5
Addition of acid results in the simultaneous occurrence of several reactions: ion exchange of Ca2+ for H+ and Na+, partialdissolution of epidote, and reduction of FeIII. Dissolved hydrogen and methane served as reducers. The integral reaction is
The coefficients of this equation are noninteger, because it is a sum of several individual reactions (exchange, dissolution, andoxidation) and the contributions of these reactions depend on the composition of the system (for instance, on the ratio of methaneand hydrogen concentrations). In general, if acid is added to the system, it is almost completely neutralized and Ca is extractedinto solution. Epidote is enriched in the zoisite end-member, and chlorite is enriched in the iron-rich (daphnite) end-member;
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S239
is too low to exert a significant influence on the charac-ter of the processes.
Iron is the main variable valence element in themetasomatic rocks after basalts. In assemblage I, ironis incorporated into hematite (FeIII) and ferromagne-sian chlorite (FeII) (Figs. 21b, 22a). In association IIiron is distributed among epidote (FeIII) and iron-richchlorite (FeII), as well as sulfides and actinolite at highR/W. The FeIII/ΣFe ratio is higher than that of the ini-tial basalt (0.15) in all model steps. This is indicativeof the reaction of iron oxidation by oxidants supplied
by seawater, O2 and and H2O at high tempera-ture. Since the solution contains a limited inventory of
O2 and , the formation of epidote results in agradual accumulation of hydrogen (Fig. 21f), whoseconcentration in the solution reaches n × 10–2 mol/kgin some calculated variants. Methane is also accumu-lated in the solution; however, since the amount ofcarbon in the system is low, methane is the third most
SO42–
SO42–
important reduced component of solution, after H2Sand H2.
For the interpretation of model data, it is importantto determine the factors that control the redox state ofthe system. A numerical experiment with the additionof a reducer into the system (Table 20b) showed thataccompanying changes can be described by the reac-tion
(18)
Similar to the addition of acid (Eq. 17), this reactionresults in Fe reduction and migration from epidote toiron-rich chlorite, but its influence on the compositionof solution is the opposite: Ca is fixed and Na isreleased. The reducer has practically no effect on thepH of the system (Table 20b). Since reaction (18)
30Ep 8Cch 16Ab 8Ca2+ 15H2+ + + +
24Zo 5Dph 10Act80 16Na+.+ + +
Table 20. (Contd.)(c) Influence of oxidizer addition
involves dissolved Ca and Na, the redox state of thesystem depends on their proportion:
(19)
Therefore, mineral assemblage II is not a classic redoxbuffer (like NNO or QFM). Exchange reaction (18) onlyimposes some relations but does not control . Thecapacity of the system with respect to reducers is verylow. If only 3 mmol H2 is added, the phase association
is changed, H2C reduction to methane begins, andthe gain of H2 concentration in the solution [Δ(H2) inTable 20b] becomes practically equal to the amount ofhydrogen introduced into the system (i.e., there is nobuffering effect with respect to reduction).
The buffering capacity of the system with respect tothe addition of an oxidant is significantly higher. Theaddition of oxygen to assemblage II shifts reaction (18)to the left. The solution is enriched in Ca2+ substitutingfor Na+, and the iron of the rock is oxidized and trans-ferred from chlorite and actinolite to epidote. However,these changes become significant only with a substan-tial addition of oxygen, more than 30 mmol per onekilogram of solution (Table 20c). A further oxygeninput results in the disappearance of sulfides from thesystem and the appearance of anhydrite, dissolved sul-fates, and hematite. The oxidation of the system isaccompanied by a decrease in pH.
In the assemblage II field, the redox state of theenvironment is controlled by the balance of cations pro-duced in the irreversible solution–basalt reaction. In themodel case shown in Fig. 21 for a reactor step with atemperature of 300°C and ΣR/W = 0.169, this balanceis expressed as (in moles per one kilogram of initialrock):
Basalt + 1.62H2O + 0.123Na+
0.917Ab + 0.198Chl + 0.593Ep + 0.428Act
+ 0.023H4SiO4 + 0.038K+ + 0.042Ca2+ (20)
+ 0.0017Fe2+ + 0.0016OH– + 0.015H2
+ 0.018CH4 + 0.0175H2S.
The same balance for a temperature of 370°C andΣR/W = 0.732 is
Basalt + 1.58H2O + 0.110Na+
0.903Ab + 0.192Chl + 0.582Ep + 0.443Act
+ 0.0016H4SiO4 + 0.038K+ + 0.033Ca2+ (21)
+ 0.0051Fe2+ + 0.00015Zn2+ + 0.0055OH–
+ 0.035H2 + 0.0081CH4 + 0.0183H2S.
With an increase in R/W, the newly formed epidoteis enriched in the clinozoisite end-member, and chloritegains the daphnite end-member (this is reflected in the
f H2K18
aNa+
aCa2+( )1/2
--------------------.≈
f H2
O30
change of calculated transitional phases of variablecomposition: Ep Ep75 Ep60 and Cch Chl50 Chl75; Fig. 21b). Because of this, FeIII/ΣFegradually decreases along the flow line approaching0.30–0.32. In such a way, a balance between the reduc-ing substances of the rock and the accumulation ofreduced products is established in the moving solution.
Trace ore metals. The behavior of ore elements isclosely related to changes of metasomatic assemblages.
In the stability region of assemblage I, copper ismainly removed from the basalt at the expense of chal-copyrite oxidation and dissolution in seawater. In thestability region of assemblage II, solution is in equilib-rium with chalcopyrite and bornite (at high T and lowR/W) or chalcocite (along association boundaries)(Fig. 22a). Owing to the high content of H2S, the equi-librium solubility of copper in the assemblage II regiondecreases abruptly, and its migration ability vanishes. Aredox barrier at the boundary between assemblages Iand II separates the segments of the downwelling limbwith high and low copper mobility (Fig. 21d), and, con-sequently, is a boundary of copper precipitation. The
main mobile copper species is Cu within the wholerange of conditions (Fig. 24c).
The behavior of zinc in the system is simpler: it isremoved from the assemblage I region and shows a lowmobility in the assemblage II region, where it precipi-tates as sphalerite owing to high a H2S content(Fig. 22a). It should be noted that the sphalerite-in linein Fig 22a does not exactly coincide with the boundarybetween the assemblages. It is displaced toward higherR/W at high temperatures. The concentration of Zn insphalerite-saturated solution depends on temperature: itincreases gradually with increasing T owing toenhanced complexation reaching n × 10–4 mol/kg at400°C (Fig. 21d).
The same tendencies are exhibited by the behaviorof lead. Similar to Zn and Cu, it is highly mobile in thestability region of assemblage I and of low mobility athigh R/W in the presence of H2S. Similar to sphalerite,the boundary of galena stability is shifted to high R/W,and both Zn and Pb are mobile within the stabilityregion of assemblage II. The main mobile Zn species
are ZnCl+ and Zn at low temperatures and Zn
at high temperature; and those of lead are Pb and
Pb , respectively (Figs. 24d, 24e).
A comparison of the behavior of Cu, Zn, and Pb inthe downwelling limb of the model is of special inter-est. Almost within the whole stability region of assem-blage I, they are completely removed from the rock,and the proportions of their concentrations in the solu-tion correspond therefore to those of the initial basalt(Zn ≈ Cu > Pb). This relationship holds with an increasein the concentration of the elements in the solutionrelated to increases in R/W and temperature. Within the
Cl2–
Cl20 Cl4
2–
Cl20
Cl3–
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S241
stability region of assemblage II, the concentrations ofthe elements decrease abruptly when the solutionbecomes saturated with respect to the corresponding sul-fides. This change occurs in the sequence Cu Zn, Pb.In the high-temperature part of the system, the model ofthe first wave (Fig. 21d) yields Zn > Cu ≈ Pb in the solu-tion. Thus, according to the model, solutions are signif-icantly enriched in zinc relative to copper and lead dur-ing the initial stage of high-temperature hydrothermalactivity.
The results of calculations imply that the relativemobility (i.e., fraction of dissolved metal in its totalcontent in reactor steps) of these elements decreases inthe sequence Pb (23%) > Zn (12%) > Cu (0.9%).16 Thisarray differs from the sequence of sulfide solubilities,which is Zn > Cu ≈ Pb under the model conditions(Fig. 21d). The relative mobility is affected by the factthat the abundance of Pb in the initial rock is two ordersof magnitude lower than those of Zn and Cu. Becauseof this, despite its relatively low solubility, Pb can becompletely mobilized from the rock by solution,whereas the more soluble Zn is partly retained in themetasomatic rock as sphalerite.
4.1.2. Chemical evolution of the system during the development of the hydrothermal process
(model of a long-lived hydrothermal system)
Long-lived hydrothermal systems differ from short-lived ones in two aspects important for the model con-sidered: (1) the amount of metasomatized rocks gradu-ally increases, and correspondingly, the fraction ofunaltered basalts in the interior of the system and therate of metasomatic zone propagation decrease; and(2) the composition of metasomatic rocks changes,because some elements are introduced with seawaterand others are removed by hydrothermal solution.These phenomena are reproduced in our model by thepassage of many sequential solution portions (waves)through a step flow reactor (Section 2.2). The amountof added fresh basalt decreases gradually in each stepaccording to Eq. (13).
The results of calculations for such a model (corre-sponding to the first wave in Fig. 21) are shown in Figs. 25and 26. The main result of the development of the pro-cess in the downwelling limb of the hydrothermal sys-tem is a gradual expansion of the stability region ofassemblage I under the influence of the addition of oxi-
dants (O2 and ) and Mg and the removal of reduc-ers (H2, H2S, and CH4) by hydrothermal solution. Theboundary between assemblages I and II travels alongthe flow line inside the system to higher temperaturesteps. This can be clearly seen from the expansion of
16These estimates depend on T and R/W; the values presented hereare mean squared values for the zone of assemblage II and thecase shown in Fig. 21. The observed differences between the rela-tive mobilities of Zn and Pb are within the uncertainties of thethermodynamic model.
SO42–
the stability fields of hematite and anhydrite (Fig. 25)and the diminishing of the stability fields of epidote,actinolite, albite, and sulfides. The evolution isimprinted in the solution composition as an increase in
the mobility of Mg and introduced by seawater(Fig. 26).
The system evolves mainly owing to a gradualdecrease in the rate of metasomatic zone propagationperpendicular to fracture walls. Correspondingly, theinvolvement of new portions of fresh basalt slowsdown, and R/W values decrease in the reactor steps.This influences the dissolution of easily mobilizedcomponents, such as SII, K, and Pb. A decrease in R/Wwith each new solution portion (wave) reduces the con-centrations of these components in the hydrothermalsolutions (Fig. 27).17 It is remarkable that iron, whoseconcentration in the solution is controlled by the solu-bility of aluminosilicates, shows a different behavior: inthe beginning of the process, SII > Fe in the solution(Fig. 21c), but this relation later reverses (Fig. 27a). Aninteresting feature of the late stage of development,when the boundary between the assemblages movesinto the high-temperature region, is that the equilibriumconcentration of sulfates increases within the assem-blage II region. As can be seen in Fig. 26b, the residual
concentration of is 2 mmol/kg for wave no. 800.Under the corresponding temperature conditions, compa-rable amounts of sulfate and hydrogen sulfide(0.n mmol/kg) can coexist.
The extensive SII removal results in the dissolutionof sulfides that were formed during the early stages ofsimulation and the reduction of their stability fields(Fig. 25). The fates of the trace ore elements (Zn, Pb,and Cu) are different, and variations in their concentra-tions in the solution are not in tandem. This peculiarityof ore element behavior can be clearly seen from thecomposition of solution in the 370°C reactor step cor-responding to the hottest part of the hydrothermal sys-tem (its focus) (Fig. 27b). Two regimes are possible forthese elements, depending on whether or not sulfidesaturation is reached in the given step.
(1) If the sulfide of an element is not formed at thelocal equilibrium between solution and metasomaticrocks, the element concentration in the solutiondecreases with each successive wave proportionally toR/W, which is exemplified by the behavior of Zn and Pbat large wave numbers (Fig. 27b).
17Periodic variations in hydrogen sulfide content were obtained inthe model within the assemblage II region (Fig. 27a). They areartifacts of the computational procedure related to the discretestructure of a step flow reactor (Section 2.1). In particular, this ismanifested in that within a temperature interval of up to 270°C,when the boundary between assemblages I and II jumps to thenext reactor step, anhydrite and pyrite appear together in the setof minerals and the concentration of SII increases simultaneouslyin solution. The subsequent water portions dissolve pyrite, andthe equilibrium solution is depleted in hydrogen sulfide.
SO42–
SO42–
S242
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
(2) If the sulfide phase is stable, the concentrationof the trace ore element depends via the solubilityproduct on the concentration of sulfur, which is amajor element. Since the concentration of H2Sdecreases during the system evolution, the concen-
tration of the ore element increases, which is exem-
plified by Cu (Fig. 27b).18
18In addition to a gradual increase, Cu concentration shows oscilla-tions opposite to those of the precipitant (SII); the reason for suchoscillation was discussed above.
150 220 260 300 340 380
Wave no. 1
Temperature, °C
Wave no. 50
Wave no. 800
100
80
60
40
20
0
%
100
80
60
40
20
0
%
100
80
60
40
20
0
%
Hem
Mgs
Dsp
Anh
Ep75
Wai
Tlc
Chl50
Dph
Act80
Prl
Qtz
Ab
Kln
Ep
Ep60
Pmp
Cch
Chl75
Tr
Ser
DphChl75
Act80
Ep60Ep75
Ab
Py, Ccp, Sph, Gn
Tlc
Anh
QtzHem
CchPmp
Chl50Tr
Chl50 Act80
Tr
Chl75
Ep60
Ep75
Ep
EpCch
Anh
QtzKln
Hem Hem Bn
AbSer
Prl
Bn
AbQtz
Hem Hem
PrlKln
Dsp
Anh Ep
Cch
Chl50 Act80
Tr
Ep75
Mgs
Fig. 25. Variations in the position of stability fields of metasomatic associations during the development of a long-lived hydrother-mal system. The case corresponds to the initial ΣR/W (400°C) = 1.26. Arrows show the position of the sulfide appearance boundary.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S243
In the case shown in Fig. 27b, zinc and lead show adual behavior. In the initial waves, they have regime (2)(sphalerite and galena are stable), which changes laterto regime (1). As a result, their concentrations in thesolution pass through a maximum at the boundarybetween these regimes (wave no. 17 for Zn and waveno. 27 for Pb). It is obvious that a change from regime (2)to (1) must be a common feature of all the elementsforming their own sulfides in the stability region ofassemblage II. The sequence of ore elements show-ing the change of regimes depends on their relativemobility in equilibrium with mineral assemblage II.According to the calculation of the first-wave model,the sequence of relative mobility is Pb ≈ Zn > Cu,which is in agreement with the results of calculationsfor the multiwave model (Fig. 27b).
The differences in the behavior of Zn and Cu duringthe evolution of the downwelling limb have importantconsequences. The model hydrothermal solution gener-ated in the early waves shows Zn > Cu, the concentra-tions of these elements become equal during further evo-lution, and the relation eventually reverses to Cu > Zn. Inthe simulation case shown in Fig. 27b, initial Zn con-centrations in the solution of 0.1–0.2 mmol/kg H2Odecrease to 0.014. In contrast, the concentration of Cuincreases initially from 0.00n to 0.09 mmol/kg anddecreases abruptly to the Zn level after wave no. 667,when copper sulfides disappear from the mineralassemblage. Similar inversions were obtained for mod-els similar to those shown in Fig. 27b but calculated atdifferent temperatures and R/W values (see also Kras-nov et al., 1990; Abramova and Grichuk, 1994; Gri-chuk, 1996). Thus, an important conclusion derived
from the simulation is that the relationships of ore metalconcentrations must change from Zn > Cu ≈ Pb to Cu >Zn > Pb during the activity period of a hydrothermalsystem.
0.06
0150
Temperature, °C
220 380340260
0.04
0.02
0.06
0m
ol/
kg 0.04
0.02
Wave no. 1Wave no. 10Wave no. 50Wave no. 200
Wave no. 400Wave no. 800
SO4Wave no. 1
H2S
Wave no. 10Wave no. 50Wave no. 200
Wave no. 400Wave no. 800
Wave no. 1Wave no. 800
(a)
(b)
mol/
kg
Fig. 26. Movement of magnesium sulfate metasomatic solutions (heated seawater) during the development of a long-lived hydro-thermal system. The case corresponds to the initial (wave no. 1) ΣR/W(400°C) = 1.26 (Fig. 21c). (a) Mg in solution and (b) sulfateion and hydrogen sulfide in solution.
–3
–71
Wave no.
–4
–5
–6
101 201 301 401 501 701601
CuZnPb
Gn
Sph Bn Cc
–6
log(M
e)
–4
–5
FeS(II)S(VI)
–2
–3
(a)
(b)
Fig. 27. Variations in the concentrations of ore elements inhydrothermal solutions during the development of a long-lived hydrothermal system. A reactor step with T = 370°C,P = 500 bar, and ΣR/W (first wave) = 0.732. The case corre-sponds to that shown in Figs. 21, 25, and 26. Arrows markthe disappearance of sulfide minerals. (a) Iron, sulfate sul-fur, sulfide sulfur, and (b) heavy metals.
300
S244
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
4.2. Metasomatic Alteration of the Feeder Conduit
As was noted in Chapter 3, the oceanic hydrother-mal systems under investigation show a limited devel-opment of metasomatic haloes around feeder conduits.They are different in this respect from ancient massivesulfide deposits on continents, where footwall hydro-thermal alterations are pronounced. The secondarymineral assemblages in the oceanic crust that could beunequivocally interpreted as alteration haloes aroundfeeders are also different from terrestrial ones: thealtered basalts show the development of iron-rich chlo-rite and are free of sericite, which is a typomorphic min-eral in the footwall haloes of massive sulfide deposits.
These data suggest that metasomatic processes inthe upwelling limb of the oceanic hydrothermal systemare different from those of ancient deposits. The reasonfor this difference is obscure. It could be related to solu-tion chemistry (ancient massive sulfide-forming sys-tems could contain magmatic fluids), the character ofinitial material (tholeiitic basalts versus island-arcbasalts and rhyolites), different T–P conditions, orother factors. This topic is discussed in more detail inChapter 7.
The goal of this section is to characterize the meta-somatic alterations developing in the convection–recy-cling model considered.
Model parameters. Metasomatic changes in spread-ing-center basalts, which are fractured and porousrocks, have an inherently coupled infiltration–diffusioncharacter. The infiltration mechanism of metasomaticcolumn formation was utilized in our model. The col-umn was regarded as an isothermal step flow reactorsimilar to that described in Section 2.1.4. The initial
solution was represented by the calculated compositionof downwelling fluid in a convection cell with Tmax =370°C, P = 500 bar, and ΣR/W = 0.732 for wave no. 1(Table 19). This composition is probably typical ofshort-lived hydrothermal systems. The conditions inthe reactor (metasomatic column) are T = 350°C andP = 300 bar. Thus, a solution entering the column wasnot in equilibrium under the given conditions. Themodel reactor consisted of 50 steps, each of which ini-tially contained 10 g of tholeiitic basalt.
The results of simulation after the passage of tensolution portions (waves) are shown in Fig. 28. It can beseen that the basalts are altered under the influence ofhydrothermal solutions to the typical propylite assem-blage albite + epidote + chlorite + actinolite. The chlo-rite shows an iron-rich composition (Chl75 and Dphtotal at ≈80% of iron end-members). There are nopotassium-bearing minerals in the column, because Kis removed by the first solution portion. The composi-tions of mafic minerals in the back zone of the column(reactors 1–3) are even richer in iron, up to the appear-ance of ferroactinolite. This reflects an Fe (also Si andNa) addition with the introduced hydrothermal solu-tions. However, this part of the column is very thin. Themajor elements Al, Ca, and Mg are not moved.
The column displays some redistribution of oreminerals (Fig. 28b). The accumulation of sphalerite (upto 0.45%), galena, and chalcopyrite was observed in thefirst step of the reactor. This process is due to the prob-lem formulation, which implies that the input solutionis generated in the downwelling limb at a temperatureof 370°C and becomes undercooled in the reactor with350°C. Several subsequent steps show a removal of ore
Fig. 28. Simulation of infiltration metasomatic columns around a feeder channel: (a) rock-forming minerals and (b) sulfides. T =350°C, P = 300 bar, and wave no. 1.
0
20
40
60
80
100
Min
eral
s, %
Fe-Tr
Dph
Eh
(a)
(b)
Ab
Chl75
Act80Dph
Ep60
1 11 21 31 41Raector no.
0
0.1
0.2
0.3
0.4
0.5C
onte
nt,
%
Ccp
Sp
Ep75
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S245
metals, after which their concentrations are stabilized ata level corresponding to the prescribed concentration ofbase metals in the initial basalts.
Thus, the results of simulation are consistent withthe observed weak development of metasomatic alter-ations around the channelway of upwelling fluid. Theyare not accompanied by sericitization, which is typicalof continental massive sulfide deposits (Metasomatismand…, 1998). The main metasomatic effect is theenrichment of iron in the altered rock. Ore minerals aredeposited in the feeder area only at the expense of cool-ing, and the metasomatic alteration of rocks is accom-panied by a minor removal of base metals from the backpart of the column.
4.3. Models of Ore Deposition
As was shown in Chapter 2, depending on the char-acter of the hydrothermal solution discharge on theocean floor, the processes of ore deposition can bedescribed by different models: slow cooling, rapidcooling, or cooling with mixing. The main factor of oredeposition is always a decrease in solution temperature,and the models correspond to various dynamic scenar-ios of the process. The simulation results for these oredeposition scenarios are given in Section 4.3.1. Theinvestigation of oceanic hydrothermal systems demon-strated that the style of ore formation varies during thegrowth of a large edifice (Section 3.2). A combined sce-nario describing this evolution and numerical results
for the corresponding thermodynamic models are pre-sented in Section 4.3.2.
4.3.1. Scenarios of ore deposition during cooling
The model of slow cooling is a step flow reactor inwhich temperature and pressure decline gradually fromstep to step (Fig. 10c). This scenario ignores solutioninteractions with the wallrocks in the feeder conduit.
Figures 29a and 29b show some representativeresults obtained for such a model: simulation of theslow cooling of the hydrothermal solution formed inthe downwelling limb at 370°C, 500 bar, and ΣR/W(370°C) = 0.244 (wave no. 10). The deposited mineralsare dominated by quartz and pyrite (Fig. 29a) with anadmixture of base metal sulfides. It can be clearly seenfrom Fig. 29b that the temperature maxima of sulfidedeposition are different for different metals: bornite(350°C); chalcopyrite (310°C); sphalerite (290°C);galena (260°C); pyrite (240°C). Under other modelparameters, the temperature ranges of sulfide precipita-tion may be different, but the deposition sequenceCu Zn Pb is rather stable. The maxima of ironsulfide deposition are more variable, and the modelcases with high R/W produced pyrrhotite at 300–320°C(usually above the temperature of sphalerite forma-tion). Consequently, the model of slow cooling allowsthe fractionation of ore metals at the expense of theirprecipitation at different temperatures.
(a)
(b)
0
0.2
0.4
0.6
0.8
1.0
Min
eral
s, g
Dph
Qtz
Bn
Ccp
Sph
Py
Dph
Qtz
Py
Sph
SphCcpBn
Bn
Ccp
Sp
Py
360 310 260 210Temperature, °C
0
0.04
0.08
0.12
0.16M
iner
als,
g
Fig. 29. Fractionation of metals during the slow cooling of solution. The composition of an ore-forming solution was obtained inthe downwelling limb at T = 370°C, P = 500 bar, and R/W = 0.244 (wave no. 10). The amount of precipitate is given for 1 kg of ore-forming solution. (a) The net result of mineral formation and (b) the deposition of ore minerals.
Py
S246
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
This model is probably valid for hydrothermal sys-tems with low discharge temperatures. As was firstnoted by Rona (1984) and reemphasized in later stud-ies, such systems are characterized by the prevalence ofZn over Pb and other ore components in the depositedproducts. Many authors explained this phenomenon bythe loss of metals (primarily, Cu) in subseafloor envi-ronments. The model of slow cooling confirms the pos-sibility of such a process resulting in the formation ofveins and disseminated mineralization (stockwork). Itsefficiency with respect to the formation of massive sul-fide bodies is probably very low, because the ore min-eralization is strongly diluted by quartz, which isdeposited within a wide temperature range below330°C.
The model of rapid cooling corresponds to the ejec-tion of high-temperature solution into bottom oceanwater, when the chemical reactions between the com-ponents of the resulting mixture do not have enoughtime to occur. This is realized in the model by the trans-portation of solutions from the high-temperature part ofthe reactor directly into one of the low-temperaturesteps, where ore deposition is calculated (Fig. 10b). Insuch a manner, the quenching of ore precipitates is sim-ulated. The main variables of this dynamic process are(a) the parameters controlling the composition of dis-charged solutions (Tmax and ΣR/W for the downwellinglimb and wave number) and (b) the P–T conditions ofthe reactor step where the solution is conveyed. Varia-tions in these parameters produce different model com-positions of ore deposits.
A significant problem with the realization of such athermodynamic model arises because of the possiblenonequilibrium character of reactions and formation ofmetastable compounds in the natural prototype of therapid discharge scenario. The metastable character ofthe natural process is evident for silica: amorphous sil-ica instead of quartz is often observed in sulfide precip-itates. There is no unambiguous evidence for the meta-stable deposition of other identified phases. In order totake this factor into account in our model of rapid cool-ing, gT(SiO2, quartz) was changed by gT(SiO2, amor-phous silica) in the reactor steps corresponding to thezone of ore deposition.
Figure 30 shows the results of calculation for oredeposition during the rapid cooling of hydrothermalsolutions (first wave) to ΣT = 150°C. It can be seen that,at varying Tmax and ΣR/W of the downwelling limb ofthe convection cell, the masses of precipitates and pro-portions of minerals in the zone of ore deposition arestrongly variable. The obvious reason for these distinc-tions is the derivation of hydrothermal solutions in thestability fields of different mineral assemblages.Indeed, solutions formed at low ΣR/W (<0.02) in con-tact with mineral assemblage I yield hematite-bearingprecipitates, whereas those in equilibrium with transi-tional assemblages (ΣR/W = 0.020–0.050) produce pre-cipitates with hematite and pyrite. The solutions thathave reacted with a considerable amount of rock pro-duce sediments dominated by silica and iron sulfide andfree of hematite.
300
350
400
T, °C
–2 –1 0 1 log(R/W)
100mg/kg
ASi
ASiASi
ASi
ASi
ASiASi
ASi
Anh
Py Cc
Hem
Ccp
HemPy
Py
PySph
Py
Py
Po
Sph
Py
Po
Po
Sph
Sph
Fig. 30. Mineralogy of the products of the rapid cooling of solutions as a function of temperature and R/W in the focus of the system.Cooling was calculated to a temperature of 150°C (P = 250 bar); the pie area is proportional to the amount of precipitate; the emptysectors correspond to aluminosilicates (chlorite, actinolite, kaolinite, and pyrophyllite); and ASi is amorphous silica.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S247
The proportions of iron sulfides and silica in the pre-cipitate are strongly dependent on the maximum tem-perature reached in the downwelling limb: the higherTmax, the higher is the fraction of iron in the sediment.Pyrrhotite is predominant in the ore material at a tem-perature of 400°C in the focus of the downwelling limb(Fig. 30). This is evidently related to the significanttemperature dependence of iron accumulation in thedownwelling limb of the system (Fig. 23b). On theother hand, the concentration of silica in the descendingfluid is not particularly temperature-sensitive, and mayeven decrease with increasing T in quartz-free assem-blages. Pyrrhotite-bearing precipitates were obtainedfrom the hottest solutions, whereas the iron sulfides aredominated by pyrite at focus temperatures of 370–300°C. The main reason for this difference lies in the Feand H2S relationships in the mineralizing solutions:pyrite is precipitated from solutions whose sulfur con-tents are much higher than those of iron (Figs. 23b,23d), whereas pyrrhotite forms at approximately equalFe and S concentrations, which is typical of solutions inthe highest temperature steps of the downwelling limb.
For the solutions formed at 370°C, the identity ofiron sulfide depends also on “quench” temperature.Figure 31 shows the dependence of the mineralogy ofthe precipitate on deposition temperature. According toFig. 31a, pyrrhotite precipitates at an ore depositiontemperature of ≥200°C, and pyrite forms at lower tem-peratures. The change of phase is accompanied by achange in solution composition: the transition from themonosulfide to disulfide leads to a decrease in SII con-
tent and an increase in dissolved hydrogen content(Fig. 31b). Thus, the precipitation of FeS2 at theexpense of H2S involves water as an oxidizer, and thechange of mineral forms depends, in such a case, onlyon the temperature dependencies of the thermodynamicproperties of the reactants. Figure 31a shows that therelationships of sulfides and opal are also dependent onquench temperature: opal does not form at T > 200°C,but the relative amount of opal increases progressivelybelow this temperature.
The silica–sulfide assemblages produced at highR/W always contain several percent sphalerite (Fig. 30)and an admixture of chalcopyrite and galena. The frac-tion of sphalerite in the precipitate reflects the accumu-lation of Zn in the mineral-forming solutions, which iscontrolled mainly by temperature. Almost completeand simultaneous deposition of the chalcophile metalsduring the rapid cooling of solutions with excess H2Sprevents fractionation of these elements (in contrast tothe model of slow cooling). Therefore, their propor-tions in the ore material are inherited from the compo-sitions of mineralizing solutions in the hottest steps ofthe downwelling limb of the system. This simple corre-lation is disturbed only in the solutions produced in thestability region of transitional assemblages (ΣR/W =0.02–0.05, Fig. 30), where pyrite and chalcopyrite areselectively deposited owing to the deficiency of sulfidesulfur in the solution.
Anhydrite occurs in the products of rapid coolingonly within a narrow field at low R/W and T = 350°C.The reason for its dependency on R/W is undoubtedly
(a)
(b)
250°C 200°C 175°C 150°C 100°CPy
Po
Sp
Ccp
Asi
Prl
100 mg/kg
Conce
ntr
atio
n,
mm
ol/
kg
20
15
10
5
0275 250 225 200 175 150 125 100 75
Quench temperature, °C
H2
S(II)
Po
Sph
Ccp
Prl PrlSph
Po
Prl
Sph
ASi
Py
PyASi
Po
SphCcp
Prl
Prl
ASi
Py
SphCcp
Fig. 31. Simulation of the effect of quench temperature on the composition of ore precipitates in the model of rapid cooling. Thecomposition of the ore-forming solution was obtained in the downwelling limb at T = 370°C, P = 500 bar, and ΣR/W (370°C) =0.712 (wave no. 1). (a) Mineral proportions in the precipitates and (b) concentrations of selected components in the solution.
S248
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
that solutions in equilibrium with mineral assemblageII are free of sulfates and contain excess reducers. Thetemperature dependency is more ambiguous: it is wellknown that anhydrite solubility in pure water decreaseswith increasing temperature (Ryzhenko, 1981), andcorrespondingly, it is not expected to precipitate inresponse to a T decrease. However, anhydrite solubilityin seawater and seawater-derived solutions is strongly
affected by complexation, which results in that HS
and Ca become the major calcium and sulfate spe-cies. Because of this, the temperature dependency ofanhydrite solubility in such solutions appeared to bemore complicated, which was reflected in the results ofsimulation: anhydrite precipitated from solutions withan initial temperature of 350°C and was absent in othercases.
The suspended matter exiting from smokers is a nat-ural ore material whose deposition must correspond tothis scenario. The calculated precipitates of silica, pyr-rhotite, pyrite, and sphalerite are the main constituentsof smoke produced by high-temperature black smokers(Feeley et al., 1990). It is interesting to note that pyr-rhotite occurs in natural analogs only in smoke and isabsent in chimney walls.19 This was usually interpretedas an outcome of metastable precipitation, because theformation of monosulfides is kinetically preferable tothat of disulfides. Our simulation demonstrated that therapid quenching of hydrothermal solution may give riseto the formation of stable pyrrhotite (Figs. 30, 31).
Grichuk et al. (1985) supposed that the formation ofhematite-bearing siliceous precipitates from the solu-tions of the downwelling limb at low R/W in the modelof rapid cooling could be correlated with white smok-ers. However, this analogy may not be quite adequate.White and black smokers often coexist within singleedifices (Lisitsin et al., 1990). It is difficult to under-stand how hydrothermal solutions with significantly
19 Except for systems in the sediment cover, such as the GuaymasBasin, Escanaba Trough, and Middle Valley fields.
O4–
Cl20
different histories of interaction with rocks could beejected through closely spaced channels without mix-ing en route to the surface. The major-component com-positions of solutions from black and white smokers ofa single system are not very different (Table 6 shows thedata on the TAG hydrothermal field). It is highly prob-able therefore that the solutions of the two smoker typeshave a common source. Hekinian et al. (1983) docu-mented conversion of a white smoker to the state of ablack smoker when an almost closed chimney wasdestroyed during sampling.
It could be suggested that the formation of whitesmokers is related to the partial loss of ore load fromhydrothermal solutions due to a minor cooling eitherbelow the surface of the ocean floor or within the chim-ney. Indeed, the temperatures of white smokers are sys-tematically lower than those of black smokers (Chapter 3).This suggestion can be tested by simulation using acombination of the models of slow and rapid cooling.Figure 32 shows the results of calculations for deposi-tion from a solution formed by wave no. 1 at Tmax =370°C, P = 500 bar, ΣR/W = 1.26 and experiencing(a) rapid cooling or (b) preliminary slow cooling (withan increment of 10°C) to 300°C. These results clearlyshow that preliminary cooling reduces the mass of pre-cipitate without changing the proportions of iron sul-fides, silica, or the general character of the solid mate-rial. The typical sulfate phases of white smokers aremissing in the combined scenario. A characteristic fea-ture of the combined discharge scenario is the slightincrease in the relative amount of sphalerite. It is knownfrom natural observations that the precipitates ofmedium-temperature hydrothermal springs are rela-tively rich in Zn.
Thus, the model of rapid cooling corresponds to theactive discharge of hydrothermal solutions of the blacksmoker type into cold seawater. The correspondencebetween the equilibrium thermodynamic model ofrapid discharge and the natural process can certainly beapproximate, because the process is very rapid, suchthat metastable phases may be formed and crystalliza-tion of some other minerals may be kinetically inhib-
ASi
Prl
Py
SphSph
ASiPy
100 mg/kg
(a) (b)
Fig. 32. Comparison of the results of rapid and stepwise cooling of hydrothermal solutions. The initial ore-forming solution wasobtained in the downwelling limb at T = 370°C, P = 500 bar, ΣR/W (370°C) = 0.712 (wave no. 1). (a) Rapid cooling to 150°C and(b) slow cooling to 300°C followed by rapid cooling to 150°C.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S249
ited. The results of calculations (Figs. 30–32) shouldprobably be regarded as a general sketch rather than atool for detailed interpretations of natural observations.However, the main conclusion that chalcophile metalsdo not significantly fractionate during the rapid dis-charge of hydrothermal solutions seems to be ade-quately substantiated.
The model of cooling coupled with mixing was con-sidered in detail by Janecky and Seyfried (1984) andBowers and Taylor (1985), but their simulation proce-dures allowed the occurrence of back-reactions withpreviously precipitated products. In order to describethe sequence of deposition of mixing products, one hasto construct a model based on the scenario of succes-sive mixing (Fig. 10d).
Figure 33 presents the results of calculation for onevariant of such a model. Mixing of hydrothermal solu-tion with cold seawater precipitates the following min-eral sequence: tremolite (350°C)–talc (350–270°C)–
pyrite (300–250°C)–serpentine (270–250°C)–magnes-ite (250–240°C)–anhydrite (230–175°C) (Fig. 33a).Minor amounts of copper sulfides (bornite and chal-copyrite, 340–280°C), sphalerite (270–230°C), andgalena (250°C) are formed simultaneously. Variationsin ΣR/W and Tmax in the downwelling limb do notchange the general sequence of mineral precipitation.At Tmax = 400°C, pyrrhotite crystallizes before pyrite;and at high R/W, magnesite and dolomite are depositedbefore anhydrite.
This model shows a decrease in the extent of anhy-drite deposition with increasing R/W. This feature sug-gests the nonconservative behavior of the seawater sul-fate ions during mixing with hydrothermal solution.Figure 33b compares SVI and SII concentrations in thesolution with their hypothetical conservative concen-trations calculated under the assumption that there areno reactions. It is clearly seen that the equilibrium con-centrations of sulfate sulfur are much lower and those
(a)
(b)
0.30
0.25
0.20
0.15
0.10
0.05
0350 320 270 220 125
Temperature, °C
Min
eral
, g
mol/
kg
0.025
0.020
0.015
0.010
0.005
0
CH4cons
Tr
Srp
Tlc
Anh
Mgs
PyAnh
Srp
PyTlc
MgsTr
350 300 250 200 150 100 50
S(VI)
S(II)
S(VI)cons
S(II)cons
H2
CH4
H2cons
Fig. 33. Sequence of mineral deposition during the mixing of hydrothermal solutions with seawater. The composition of ore-forming solution was obtained in the downwelling limb at T = 370°C, P = 500 bar, and ΣR/W = 0.732 (wave no. 1). The amountof precipitate is given for 1 kg of ore-forming solution. (a) The mineralogy of the precipitate and (b) the concentrations ofsulfur species and reducers in the solution. The subscript cons denotes the hypothetical conservative behavior of componentsduring mixing.
S250
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
of sulfide sulfur are much higher than the conservativevalues. This indicates that a sulfate reduction occurs inthe model up to a temperature of 210°C. As H2 andCH4 serve as reducers in the system, their concentra-tions are much lower than the conservative values.
With increasing R/W, when the solutions generatedin the downwelling limb become more reduced, thisprocess is more pronounced and continues to tempera-tures of 200–175°C. The reality of sulfate reduction inthe course of mixing is disputable. Experiments suggestthat the sulfate ion can be reduced to hydrogen sulfideunder hydrothermal conditions (Ohmoto and Lasaga,1982), but this reaction is rather sluggish at tempera-tures below 300°C and can be slower than mixing.
In order to account for this factor, we calculated ametastable model for cooling with mixing. In thismodel sulfur was represented by two quasi-elements(Methods of…, 1988), SVI and SII, which prevented thereduction reaction. The results of calculation for such ametastable model are shown in Fig. 34. They indicate thatthe metastable and equilibrium models diverge not only insulfur behavior but also with respect to other components.The deposition of magnesium silicates (except for talc) isdepressed in the metastable model, and the main newlyformed phase is anhydrite (Fig. 34a). The deposition of allsulfides is shifted to lower temperatures: pyrite,220−100°C; sphalerite, 210–75°C; chalcopyrite,
230−220°C; and galena, 150–125°C.20 In cases withlow R/W, considerable amounts of opal appear startingfrom 290°C. The reason for the decrease of the role ofmagnesium silicates in the metastable model is the signif-icantly more acidic solution pH (< 4) as compared with theequilibrium model (Fig. 34b). In the equilibrium model,the pH of solution increases at the expense of the reactions
+ CH4 H2S + H2CO3 + 2OH–
and + H2 H2S + 2H2O + 2OH–.The metastable model is much more realistic than
the equilibrium one and the composition of its mineralproducts corresponds to those of white smokers.
Thus, the difference between white and black smok-ers is probably related not to the compositions of min-eral-forming solutions or temperature parameters butrather to different styles of the discharge process. In thecase of a vigorous discharge, smoke generationapproaches the quenching scenario during rapid cool-ing and black smokers are formed. In the case of acalmer discharge, the role of mixing processes (which
20 Graphite crystallization was observed in the metastable modelbelow 100°C at the expense of methane oxidation by oxygen dis-solved in seawater. The attainment of equilibrium in this reactionis improbable at such temperatures for kinetic reasons. However,small amounts of graphite were found in the smoke (Jedwab andBoulegue, 1984). Graphite formation was excluded in the calcu-lations shown in Fig. 34.
SO42–
SO42–
360 320 270 220 1250
0.05
0.10
0.15
0.20
0.25
0.30
Min
eral
s, g
Tlc
Anh
Py
Tlc
Anh
Py
350 330 290 250 210 125Temperature, °C
3.0
3.5
4.0
4.5
5.0
5.5
6.0
pH
(a)
(b)pH, metastablepH, equilibrium
Fig. 34. Metastable sequence of mineral deposition during the mixing of a hydrothermal solution with seawater with inhibitedreduction of the sulfate species. The ore-forming solution is identical to that shown in Fig. 33. It was obtained in the downwellinglimb at T = 370°C, P = 500 bar, and ΣR/W = 0.732 (wave no. 1). The amount of precipitate is given for 1 kg of ore-forming solution.(a) The mineralogy of the precipitate and (b) the acidity of the solution during mixing.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S251
probably begin already within the edifice) and relatedchemical reactions increases. The temperaturedecreases inevitably, and white smokers are developed.
It should be noted that neither of the cases consid-ered produces ore occurrences significantly enriched inbase metal sulfides. The fraction of sphalerite is nohigher than 4% in the model of rapid cooling (Fig. 31),and 1% in the model of slow cooling with mixing. Theformation of massive sulfide bodies with high Zn andCu concentrations, which are often observed in nature,is not reproduced in such models.
4.3.2. Scenarios of growing edifices
As was mentioned above, the style of ore-formingprocesses in large sulfide mounds is significantly differ-ent from that in small individual smokers. In particular,they are characterized by the predominance of diffusedischarge over the whole surface of the body. Such acharacter of solution discharge is favorable for efficientore matter deposition (Krasnov, 1993). Short-lived sys-tems with rapid discharge lose more than 90% ore met-als with smoke, whereas large bodies retain much moreof the ore material. Large edifices show zonal internalstructures. The simulation of the formation of such oreobjects is a special challenge. The model of this processwas developed by the author together with E.E. Abram-ova and A.V. Tutubalin (Grichuk et al., 1998).
The available data on the internal structure of edi-fices and their development in time are summarized inChapter 3. The most general conclusion from these datais that the large edifices have zonal structures (Fig. 18)and the duration of their formation is n × 102–n × 104 y.During the initial stages, ejection of high-temperaturesolutions on the ocean floor produces a rapidly growingembryonic anhydrite–sulfide edifice, mainly owing tomixing with cold seawater. It grows further at theexpense of both the deposition of ore minerals fromnew solution portions percolating through the body,and the mixing of these partially altered solutions withthe surrounding seawater on the surface of the body.The development of internal zoning in an ore body iscontrolled by two factors (Krasnov, 1993): (a) varia-tions with time in the composition of mineralizing fluidand (b) metasomatic redeposition of material within theedifice during the percolation of hydrothermal fluidthrough it. Such a mechanism for the formation ofzonal ore edifices is not specific to spreading zones andmay operate in all submarine hydrothermal systems.
The development of an adequate model for the for-mation and evolution of a zonal ore body requires spe-cial modeling methods. The changes of ore-formingsolutions over time are reproduced by the model of along-lived hydrothermal system (Section 4.1.2). Thismodel must be supplemented by a block describingprocesses in the zone of ore deposition (within and onthe surface of the ore body).
All the parts of the hydrothermal system, includingthe convection cell and the zone of ore deposition, canbe simulated by step flow reactors. A model for thezone of ore deposition must include two constituents.One of them describes deposition and replacement pro-cesses during solution filtration within the ore body,and the other describes deposition on the surface of thebody during mixing of the emanating solution withambient seawater (Fig. 35). A characteristic feature ofthe reactor simulating the zone of ore deposition is thatthe boundaries of temperature steps move with time inresponse to the growth of the ore body. As a result,materials deposited at some temperature appear later ina higher temperature region, which is one of the mainfactors of metasomatic replacements.
The thermodynamic modeling of a growing orebody requires the knowledge of temperature distribu-tion within the body, its surface temperature, andgrowth rate. The distribution of temperature within thebody controls zoning patterns of ore deposition, and thesurface temperature defines the relative amount andcomposition of matter deposited during mixing withseawater. These characteristics vary along with thegrowth of the body. There is almost no information onthe distribution of temperatures within modern sulfideedifices on the ocean floor. The temperatures of the for-mation of most ancient ore bodies are known only qual-itatively: hot central parts (300–350°C) and coolerouter zones (150–200°C) were reconstructed in suchbodies (Franklin et al., 1981). We therefore used a the-oretical estimate for the distribution of temperature in agrowing ore body obtained through the calculation ofits thermal model.
Thermal model of a sulfide edifice model conditions.The ore body is approximated in the thermal model bya hemisphere with radius R, whose center is situated inthe point of ore-forming solution influx (Fig. 36). Theore-forming solution percolates through the ore bodytoward its surface (tangential component of the perco-
Redeposition
of ore m
atter
Seawater
M I X I N G
T1
T2
T3
T4
T5
T6
Fig. 35. Cartoon illustrating the model of a growing orebody.
S252
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
lation velocity is zero) and emanates into bottom water.Heat is transferred within the ore body by convectivefluid flow and thermal conduction through the solidphase. Heat loss into the environment occurs via theejection of solution filtered through the body and heatremoval from the body surface. The thermal regime ofthe body can be regarded as steady-state, because itsgrowth is much slower than heat transfer.
In order to simplify the model, the thermal conduc-tivity and filtration properties of the body wereassumed to be uniform throughout the whole volume.The heat capacity (C) and density of the solution weretaken to be temperature-invariant, and correspondingly,the density component of solution convection withinthe ore body was ignored. Heat transfer across thelower boundary of the body was also ignored.
The spherical symmetry of the ore body allows us toobtain an analytical solution to this problem (it wasderived by Tutubalin).
Calculation of the surface temperature of the body.For the model considered, heat removal from the sur-face into the environment can be described by the New-ton equation
(22)
where α is the heat transfer coefficient, S is the surfacearea of the body (area of a hemisphere with radius R isS = 2πR2), TR is the surface temperature, and Text is thetemperature of bottom water (it can be taken as 0°Cwithout sacrificing the model quality).
Heat input from the source is
(23)
where q is the discharge rate of the hydrothermal sys-tem, and T0 is the source temperature. Owing to themass balance condition for the solution entering fromthe source and exiting into seawater, the convective heatloss into the environment is
(24)
For a thermal steady-state of the ore body, heat bal-ance must be satisfied. Taking into account Eqs. (22)–(24), it can be expressed as
(25)
The surface temperature of the ore body can bereadily obtained from Eq. (25):
(26)
The coefficient of the T0 term reflects the contribu-tion of convective heat transfer in the heat balance onthe body surface.
∂Q∂t-------⎝ ⎠
⎛ ⎞cond
–
αS T R Text–( ),=
∂Q∂t-------⎝ ⎠
⎛ ⎞+
qCT0,=
∂Q∂t-------⎝ ⎠
⎛ ⎞conv
–
qCT R.=
qCT0 2πR2αT R qCT R.+=
T R T0qC
2πR2α qC+------------------------------.=
T(r)
TR
T0
rR
123
Fig. 36. Thermal model of the formation of an ore body.(1) Solution flow and convective heat transfer, (2) heattransfer by conduction, and (3) heat loss at the surface.
0.2
0
T(r)/T0
Radius, m
0.4
0.6
(Ò)
0.8
1.0
1.2
10 20 30 40 50
Discharge rate ofthe system, kg/s
110
20
0.2
T(r)/T0
0.4
0.6
(b)
0.8
1.0
1.2Outer radius of
the body, m
5
10
20
0
0.2
T(R)/T0
0.4
0.6
(a)
0.8
1.0
1.2 Discharge rate ofthe system, kg/s
20
0
0.5
1
3
10
50
Fig. 37. Results of the calculation of the thermal model ofan ore body. (a) The surface temperature of the ore body asa function of its radius and the discharge rate of the vent;(b) distribution of temperature within bodies of varioussizes at a fixed discharge rate of 10 kg/s; and (c) distributionof temperature within a body with a radius of 50 m and avarying discharge rate.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S253
Figure 37a shows the dependency of surface tem-perature on the radius of the body and the discharge rateof the source. This diagram suggests that the surfacetemperature can be very high only for a very smallbody, whereas the surface of large ore bodies is effi-ciently cooled under any realistic discharge rate.
Temperature distribution within the body. Let usconsider heat transfer across an arbitrary spherical sur-face of radius r within the body (Fig. 36). It consists ofconvective transfer due to solution circulation and con-ductive transfer. The latter component can be written as
(27)
where λ is the heat conductivity of the solid. Since theproblem is steady-state, heat transfer through anyspherical surface is the same and equal to heat supplyfrom the source, it follows from Eqs. (23), (24),and (27):
(28)
The solution to this equation is
(29)
where γ is the constant of integration. It can be deter-mined by setting r = R and equating the T(r) value fromEq. (29) with T(R) from Eq. (26):
(30)
After substituting this value into Eq. (29) and sim-plifying, we obtain
(31)
Figure 37b displays the calculated temperature dis-tribution within bodies of various sizes at a certain dis-charge rate of the source, and Fig. 37c shows the distri-bution of temperature at a certain size of the body andvarying discharge rates of the source. These diagramsindicate that small-sized ore bodies (10–20 m) with dis-charge rates similar to those observed in modern hydro-thermal systems (0.5–5.0 kg/s) (Little et al., 1987) arecharacterized by extensive high-temperature zones andnarrow peripheral zones with high temperature gradi-ents. The hot zone of large bodies with a radius of morethan 50 m occupies a smaller volume fraction and maypropagate into the outer part of the body only at highsource discharge rates. The temperature gradients in theouter parts of large edifices are significantly smaller.
∂Q∂t-------⎝ ⎠
⎛ ⎞r cond
λSr∂T∂r------,–=
2πr2λ∂T∂r------– qCT r( )+ qCT0.=
T r( ) T0 γ qC2πλr------------–⎝ ⎠
⎛ ⎞ ,exp+=
γ T02παR2
2παR2 qC+------------------------------ qC
2πλR--------------⎝ ⎠
⎛ ⎞ .exp–=
T r( ) T0 12παR2
2παR2 qC+------------------------------ qC
2πλ---------- 1
R--- 1
r---–⎝ ⎠
⎛ ⎞exp–⎩ ⎭⎨ ⎬⎧ ⎫
.=
The growth of an ore body can be described by theequation
(32)
where Mn is the mass of the ore body after the passageof the nth solution portion from the downwelling limbof the system, ΔSi is the mass of ore precipitate pro-duced from one kilogram of hydrothermal solution forthe ith wave (calculated in the thermodynamic model ofthe ore deposition zone and dependent on the parame-ters of the downwelling limb), and τ is the time of dis-charge of a single solution portion.
In contrast to the downwelling limb of the system,where the time parameter τ is in fact a scaling factorand is not used in the calculation of R/W by Eq. (11),Eq. (32) requires the numerical value of τ. This param-eter connects the rate of evolution of the downwellinglimb of the convection system with the evolution of theore body.
The observation periods of modern oceanic systemare not long enough for the determination of τ from thechemical analyses of solutions. It can be estimatedfrom indirect evidence as 108–109 s (3–30 y). Givensuch values, the solutions in young hydrothermal sys-tems from the East Pacific Rise must not significantlychange their composition during the observation period(≈10 y), whereas the long-lived TAG system in the Mid-Atlantic Ridge has evolved significantly, which wasrecorded in the relationship Cu > Zn in the solution(Edmond et al., 1995).
Equation (32) implies that the solid phases formedin the ore deposition zone join the ore body. It is knownfrom natural observations that this is not the case. Inyoung systems the major portion of ore components (upto 98%) is removed with the smoke (Hekinian et al.,1983), and this situation changes only with the growthof bodies. The parameter q of Eq. (32) therefore has themeaning of effective discharge, accounting for part ofthe measured total discharge. The discharge rates ofindividual smokers were estimated between 0.5 and5.0 kg/s (Little et al., 1987). Some other pieces of evi-dence for this parameter are shown in Table 5.
Since the objective of this section does not includethe simulation of a particular object, the following esti-mates can be accepted for the reference model: τ = 108 s,q = 10 kg/s; λ = 20 J/s cm2 (corresponding to the mix-ture of 75% quartz and 25% pyrite), α = 100 J/s cm2 °C(Handbook of…, 1987), and the heat capacity of wateris taken to be temperature independent (4184 J/g °C).
Thermodynamic model of a sulfide edifice. The ther-mal model allows the parameterization of the flow reac-tor simulating the process of submarine ore deposition.This reactor consists of two sequential parts. The first ofthem is correlated with the inner part of the body, whereore components are deposited owing to a decrease intemperature and replacement of previously formed pre-
Mn ΔSiqτ( ),i 1=
n
∑=
S254
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
cipitates. The second part describes reactions accompa-nying the mixing of solution exiting from the edificewith the ambient seawater. The boundary between theseparts of the reactor is defined by the surface tempera-ture of the ore body calculated by Eq. (26). Since thesteps of the flow reactor are fixed in our model toboundary temperatures, the number of steps corre-sponding to the interior part of the body will increase atthe expense of steps related to the mixing zone with thegrowth of the ore body and cooling of its surface. Thesetemperature steps gradually expand and move awayfrom the mouth of the feeder during the growth of theore body at the expense of continuous ore deposition onits surface. As a result, the previously deposited mate-rial appears under conditions of progressively increas-ing temperature and can be partially removed and rede-posited by solution closer to the surface of the ore body.Since the temperature distribution in the interior part ofthe ore body is uniform (Figs. 37b, 37c), in order to bet-ter reproduce its structure the increment of reactor stepswas reduced to 5°C within the range 350–330°C. Thepressure in the ore deposition zone was taken to be250 bar, which corresponds to a sea depth of about2.5 km. In accordance with mineralogical data on themodern hydrothermal ores of the ocean, quartz waschanged in the model of ore deposition by more solublemetastable amorphous silica.
The software realization of this model required theaddition of thermal and dynamic blocks to the standardcomputational procedure of a multiwave flow reactorimplemented in the GBFLOW program. After the pas-sage of a portion of hydrothermal solution through thereactor (ore deposition zone), the mass of the ore body,the temperature of its surface, the distribution of tem-perature, and the radii of temperature zones (reactorsteps) were recalculated. Then, the composition of thematter in the reactor steps was matched to the new posi-
tions of temperature boundaries within the body. Forthe outer zone of ore deposition where mixing takesplace, the amount of seawater participating in the mix-ing was calculated from heat balance using the methoddescribed in the previous section, Eq. (14). A new por-tion of seawater was then passed through the down-welling limb of the convection cell, the feeder channel,and the zone of ore deposition, after which the calcula-tion cycle was repeated. These computations were car-ried out using the GRDEP program, which was devel-oped for this purpose.
Results of modeling. Several calculations were per-formed for different Tmax and ΣR/W in the downwellinglimb of the system. The results of calculations for a casewith Tmax = 350°C and (ΣR/W)1 = 1.22 are describedbelow in some detail. This case is presumably mostsimilar to the widespread natural prototypes (referencevariant). The major differences related to variations inΣR/W are characterized. The composition of the parentsolution from the sulfide edifice is shown in Fig. 38. Itsvariations with time are identical to those discussed inSection 4.1.2 (Fig. 27), but the compositional evolutionof solution is slower because of the higher rock/waterratio. In Fig. 23e the initial point of the reference vari-ant is situated in the field of sulfur predominance overiron in the parent solution. The solution becomes rela-tively richer in iron during the evolution of the system(Fig. 38).
The history of the model ore body is distinctly sep-arated into several stages. The first stage corresponds tothe earliest solution portions passing through the down-welling limb, when the main factor of deposition is themixing of hydrothermal solution with cold seawater.This stage produces precipitates consisting mainly ofanhydrite and pyrite, which form an incipient (embry-onic) edifice (Fig. 39a). The percolation of hydrother-
1 501
Me, mol
Wave no.1001 1501 2001 2501 3001
Fe
10–1
10–2
10–3
10–4
10–5
10–6
10–7
S
Cu
Zn
Pb
Fig. 38. Ore elements in the parent hydrothermal solution for the reference variant with T = 325°C at the vent mouth andΣR/W = 1.22.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S255
mal solution through it causes a rapid replacement ofanhydrite by a silica–sulfide material in the inner part ofthe edifice (Fig. 39b). The mixing-related deposition ofanhydrite in the outer zone ceases when the tempera-ture of solution emanating from the surface of the bodyfalls below 150°C (anhydrite solubility is characterizedby a negative temperature dependency). Extensive sil-ica deposition begins simultaneously in the peripheralzone. Zinc sulfide (sphalerite) is concurrently depositedin the zones with temperatures of 300–200°C, and itsfraction in some temperature steps is higher than 50%.This stage is completed by the disappearance of anhy-drite from the ore body (solution wave no. 28 in the ref-erence variant, which corresponds to a growth time ofapproximately 100 y).
The second stage is characterized by the depositionof silica and pyrite, and the maximum of pyrite deposi-tion moves from the hot zone to temperature steps of280–230°C during the growth of the ore body. Simulta-neously, sphalerite is gradually redeposited in the outerzones of the body (200–150°C), and its content may beas high as several weight percent in some steps. In the
reference variant, this stage continues up to solutionwave no. 2600 (≈8 ky) (Fig. 39c).
During the third stage, the progressive removal ofsulfur from the rock and the decrease in its content inthe parent solution give rise to partial pyrite replace-ment by magnetite in the hottest part of the body(Fig. 39d). Sphalerite disappears from the hottest cen-tral zones of the body, and minor amounts of coppersulfides (chalcopyrite and bornite) are deposited (0.1–0.2%) (up to wave no. 3200 in the reference variant,corresponding to ≈10 ky).
The maximum contents of copper incorporated inchalcopyrite and, later, bornite and chalcocite arereached at the late stage of development, when anhy-drite appears in the hot zone, which indicates a penetra-tion of seawater sulfate through the whole downwellinglimb of the system (Figs. 39e, 39f). Because of the defi-ciency of sulfide sulfur in the solution, sphalerite andgalena are rapidly dissolved in the outer zone duringthat period, and lead and zinc are removed outside theore body. This is the longest stage, which was tracked inthe reference variant to solution wave no. 6500 (≈20 ky).
00 10 20 30 40 50 60 70
20
40
60
80
100
Min
eral
, %
(e)
0 10 20 30 40 50 60 70 80
(f)
Radius, m
AL
Anh
Ccp
PySph
ASiASi
Py Mag
Mag
AL
Anh
Ccp
0
20
40
60
80
100M
iner
al, %
0 5 0 5 10
Anh
(d)(c)(b)(a)
Py
Sph
AL
ALAL
Py
Py
Anh
ASi
Sph
ASi
ASi
AL
Sph
Py
Mag
0 10 20 30 40 50 600 10 20 30
Fig. 39. Mineral proportions in the radial cross-section of a growing ore body. (a) Wave no. 3, (b) wave no. 20, (c) wave no. 500,(d) wave no. 3000, (e) wave no. 4000, and (f) wave no. 6500. AL is aluminosilicate and ASi is amorphous silica.
Sph
S256
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
During the final stage of simulation, extensive hematitedevelopment begins in the central part of the bodyunder the influence of the oxidized components of sea-water, pyrite and copper sulfides are gradually oxi-dized, and the hydrothermal structure loses its similar-ity to massive sulfide ores.
During the whole history of the ore body, minoramounts of galena (0.n–0.0n%) are formed in the outercold zones of the body (T < 175°C) and also in its center
at wave nos. 100–300. Throughout the whole evolutionof the ore body, its central part is made up of massiveores (pyrite ± magnetite ± sphalerite ± chalcopyrite).Silica is the most abundant component in the peripheralpart of the body.
The relationships of the size of the body and itsinternal zoning for various growth stages are presentedin Fig. 40. This diagram clearly shows that distinct zon-ing patterns are developed in the ore body under theinfluence of temperature gradients, and this zoningevolves in time owing to changes in the composition ofthe parent solution, i.e., owing to the metasomatic evo-lution of the interiors of the hydrothermal system.
The maximum Zn concentrations (up to 1.8% in thecalculated case) are attained in the model ore body dur-ing the initial stage of its growth (Fig. 41c). Sphaleriteis very unevenly distributed in the volume of the body,and there are temperature zones strongly enriched inZn. Although the influx of zinc from the interiors of thesystem continues during subsequent stages and its totalamount in the body increases (Fig. 41e), the averageconcentration declines owing to dilution with pyriteand silica. Lead shows the same behavior. The concen-tration of Fe in the ore body changes relatively weaklyand is about 10% in the reference variant (Fig. 41b).The Fe/S ratio decreases regularly in accordance withthe changing proportions of these elements in the par-ent solution and the gradual replacement of pyrite bymagnetite in the hottest part of the ore body. In contrastto these elements, the abundance of Cu in the ore bodyis very low during the early growth stages. Copper con-
(a) (b) (c)
(e)(d)
50 m
1 2 3 4 5 6 7 8
Fig. 40. Relationships of the size and structure of a modelore body for various growth stages. (a) Wave no. 20,(b) wave no. 500, (c) wave no. 3000, (d) wave no 4000, and(e) wave no. 6500. (1)–(3) Massive sulfide ores: (1) iron,(2) zinc, and (3) copper; (4) quartzite and disseminated ore;(5) magnetite; (6) anhydrite; (7) pyrite, and (8) sphalerite.
(a)
(b) (d)
(c) (e)
0
1000000
2000000
3000000
4000000
Mas
s, t
05
101520253035
Conce
ntr
atio
ns,
% Febulk S
0
50000
100000
150000
200000
250000 Febulk S
1 2001 4001 6001Wave no.
0
0.5
1.0
1.5
2.0
Conce
ntr
atio
ns,
%
1 2001 4001 6001Wave no.
0
5000
10000
15000
20000CuZn
CuZn
Fig. 41. Reference model for the growth of an ore body. (a) The mass of the ore body, (b) the concentrations of major ore compo-nents, (c) the concentrations of base metals, (d) the amounts of major ore components, and (e) the amounts of base metals.
Am
ount,
tA
mount,
t
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S257
centration begins to increase only after the intensifica-tion of copper removal from the interior parts of thesystem (Fig. 38) and becomes equal to that of Zn at thefinal growth stage, when anhydrite appears within thebody. In contrast to the concentrations of the ore ele-ments, their total amounts in the model body increasein accordance with the growing mass of the body(Figs. 41b, 41c). Only the final stage of development ismarked by zinc and lead removal from the body.
An increase in the total mass of the ore body isshown in Fig. 41a for the reference model variant. Itincreased from 12 kt at the stage of an embryonic anhy-drite edifice (Figs. 39a, 39b, 40a), 2 Mt at the stage of apyrite–sphalerite body (Figs. 39c, 39d, 40b, 40c), to3.5 Mt by the beginning of hematitization (Figs. 39f,40e).
For other values of the R/W parameter, the evolutionof zoning patterns and the dynamics of accumulation ofore elements may be somewhat different. At low R/Wvalues (ΣR/W1 = 0.40), the complete cycle of develop-ment of the model body corresponds to a smaller num-ber of waves, 100–600. Because of this, the masses ofresulting edifices are proportionally smaller than thoseof the reference variant for the same discharge rate val-ues. The evolution of the ore body in these variants ofsimulation differs from that described above because ofthe lower contents of S, Zn, and Pb in the parent solu-tions and different Fe and S relationships (Fe > S,Fig. 23e). This gives rise to the rapid appearance of sul-fide sulfur deficiency in the ore-forming solution owingto its precipitation by excess Fe. Therefore, the removalof Zn and Pb outside the ore body begins already at thestage of a silica–pyrite edifice with a zoning shown inFig. 39d. As a result, by the final stage of development(replacement of sulfides by hematite), the edifice has aniron sulfide composition with a small (≈0.2%) copperadmixture.
The simulation with an initial ΣR/W of 8 is, in gen-eral, identical to that described above except for ratherslow variations in the downwelling convective flow: thepenetration of sulfate species through the hydrothermalsystem is only detected at a wave number of more than10 000, which is equivalent to a time period of n × 104
y if τ is about 108 s. According to the observations inrelatively long-lived hydrothermal systems of the Mid-Atlantic Ridge, such total lifetimes may be attained inthe case of repeated reactivation of hydrothermal sys-tems within a single crustal block (Lalou et al., 1995).Such nonsteady-state systems are not described by ourmodel. In the case of long-lived systems, it is also pos-sible that the sulfide body will be overlain by extrusiverocks, which will give rise to a multilevel deposit.Therefore, the case with ΣR/W = 8 can be regarded as aboundary case for our model.
Discussion. The simulated evolution of the modelore body at the early stages of ore formation is consis-tent with the well-known significant zinc enrichment inmodern oceanic sulfide edifices and with the recon-
structions of their formation sequences (Fig. 24). Theinitial anhydrite body on the seafloor is also replaced bysilica and sulfides with the formation of characteristiczoning patterns: Cu in the middle and Zn in the periph-ery. The zoning patterns obtained by the simulation ofmedium and late stages are, for the most part, similar tothose observed in massive sulfide deposits of theCyprus and Ural types. The model reproduces the spa-tial and temporal sequence magnetite → pyrite + cop-per sulfides → pyrite + zinc and lead sulfides. The evo-lution of the ore body in time with a late stage of copperenrichment and the replacement of sulfides by magne-tite near the mouth of the channel (and later by hema-tite + anhydrite) are in agreement with the availabledata on natural analogs. This suggests that the mecha-nism of ore formation is controlled by two factors, themetasomatic evolution of the feeding hydrothermalsystem and temperature gradients in the ore depositionzone, and corresponds in major aspects to the physico-chemical processes accompanying the formation ofboth modern sulfide bodies on the oceanic floor andmassive sulfide ores of the Cyprus and Ural types.
It is obvious that there are other important factorsinfluencing ore deposition during the formation of mas-sive sulfide deposits. A detailed analysis of the resultsof simulation reveals a number of discrepanciesbetween the model and its natural prototypes, whichsuggests that some factors of the natural process werenot accounted for in the model.
(1) The simulation of a large ore body showed thatconsiderable amounts of matter are deposited in tem-perature steps occurring within the body, i.e., in its porespace. Our model ignores changes in the filtration prop-erties of the ore body and implies that they provide con-ditions for the percolation of hydrothermal solutionwith the given discharge rate. Clogging of pores(mainly by silica) must lead to the cessation of solutionflow. According to the calculation of hydrostatic equi-librium in a hydrothermal system, the clogging of themouth of the feeder channel can generate an excesspressure of about 30–60 bar, which is sufficient for theformation of hydrofractures in the ore body. In recentpublications on the structure of large sulfide mounds inthe Mid-Atlantic Ridge (Bogdanov et al., 1997b), thisphenomenon was referred to as the hydrothermalexplosion. The appearance of fissure channels for dis-charge will result in the penetration of high-tempera-ture solutions into the outer parts and distortion of zon-ing patterns predicted by the model. Indeed, many largesulfide structures on the ocean floor comprise localizedgroups of vents (smokers) (e.g., Lisitsin et al., 1990),which account for 10–20% of the total discharge of thefeeding system (Rona et al., 1993). The collapse ofthese chimneys results in the input of relatively high-temperature products of ore deposition into the outerparts of the sulfide edifice. The occurrence of hydro-fracturing phenomena in nature is also suggested bywidespread ore breccias in ancient massive sulfidedeposits. These breccias show a persistent correlation
S258
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
with a certain domain of zoning at the transition frompyrite–chalcopyrite to pyrite –sphalerite ores (Zlotnik-Khotkevich, 1992).
(2) The simulated cases revealed a significant deficitof sulfide sulfur in the ore body in the stage of coppertransfer activation (Fig. 39f). Under model conditions,combined copper and sulfur transport is thermodynam-ically limited because of the low solubility of coppersulfides. The deficit of solubility is far greater than thepossible uncertainties of thermodynamic calculations.The probable explanation for this discrepancy is thatthe model ignores possible phase separation (boiling)in the interior parts of the system and in the ascendingchannel (see Chapter 6).
(3) The penetration of seawater sulfates through thewhole convective system that was obtained in themodel of the final stages of ore body development isprobably not realized in modern high-temperature oce-anic hydrothermal systems. Anhydrite found in theinterior parts of a sulfide body from the TAG field prob-ably has a different origin (subsurface mixing; Millset al., 1998). However, late hypogenous anhydrite iscommon in the footwall levels of some ancient massivesulfide deposits of the Central and Southern Urals andother regions (see Stolyarov, 1972 for a review). Theorigin of this anhydrite was traditionally explained bythe influence of an endogenous source, and the changefrom sulfides to sulfates was attributed to the hypothet-ical redox evolution of a magma chamber. The modelpresented above readily explains this fact by the pene-tration of seawater sulfates through the whole hydro-thermal system in the case of its prolonged activity.This explanation is consistent with the sulfur isotopesystematics of anhydrite from the footwall of massivesulfide deposits (Vinogradov et al., 1968; Grinenkoet al., 1969), which appeared to be similar to that ofseawater of the corresponding age.21
Thus, the combined mechanism of submarine oredeposition includes ore matter deposition due to themixing of hydrothermal solutions with bottom seawaterand cooling and takes into account the evolution of thefeeding hydrothermal system and metasomatic replace-ments within the ore body. The results of our simulationshow that this mechanism can produce a zonal ore bodywith time-varying chemical and metallogenic charac-teristics. The early stages of this evolution are corre-lated with the well known sulfide edifices on the oceanfloor. The mature stages give rise to massive sulfide orebodies, whose compositions and zoning patterns aresimilar to those of known ancient continental ore bod-ies. The advanced evolution of the feeding systemresults in the hematitization of the sulfide edifice,which loses its similarity to massive sulfide ores. Thissituation is more probable in small near-surface sys-
21 A marine sulfur source for hypogenous anhydrite was proposedby Stolyarov (1965), but his explanation was not supported bysubsequent studies.
tems, which provide conditions for the extensive alter-ation of rocks by hydrothermal solutions.
4.4. Verification of Simulation Results
The first question arising during the analysis ofresults derived from a complex model is the correctnessof this model, i.e., its consistency with the naturalobject. The verification of the model of an oceanichydrothermal system was based on the comparison ofsecondary mineral assemblages developing afterbasalts in the downwelling and upwelling limbs of theconvection system, the mineral and chemical composi-tions of ore precipitates, and the chemical compositionsof solutions produced by the hydrothermal system.22
(1) The model described in the previous sectionsreproduces the discrimination of metasomatic rocksinto two large groups, chlorite and epidote ones. Such adiscrimination was first established in natural oceanicmetabasalts by Humphris and Thompson (1978) andrepeatedly confirmed by subsequent studies (Mottl,1983; Kurnosov, 1986). The model sets of mineralsobtained for these two associations (quartz + chlorite +hematite + talc and albite + epidote + actinolite + chlo-rite + sulfides, respectively) are typical of the green-schist and propylite metasomatic facies. The samesequence of mineral assemblages is characteristic ofmetasomatic haloes around massive sulfide deposits ofthe Ural type (Surin, 1993). The simulation implied anarrow stability field of carbonates (150°C, dolomite atlow R/W and calcite at extremely high R/W). The lowcontent of carbonates in the secondary associations ofthe oceanic crust distinguishes them from the metaso-matic rocks of ophiolitic complexes, where carbonatesare widespread (Coleman, 1979). The presence ofanhydrite in mineral assemblage I and its highest abun-dance within 200–250°C are in line with evidence fromthe sections of the young oceanic crust (data for ODPHole 504B; Alt et al., 1983, 1989).
When rock-forming metasomatic minerals weresimulated as phases of intermediate composition (epi-dote, chlorite, and actinolite), their compositionsappeared to be consistent with their natural prototypes.For instance, according to Gillis and Robinson (1990),epidote with 20–25% pistacite end-member is typicalof the Troodos ophiolites. Such intermediate phases(epidote-60 with 20% and epidote-75 with 25% pistac-ite end-member, respectively) were obtained in assem-blage II of the downwelling limb (Fig. 21b). Amongother characteristic features reproduced by the modelare the increase in the iron content of chlorite with theintensification of metasomatic alteration (Mottl, 1983;Surin, 1993) and the higher Fe/(Fe + Mg) ratio of chlo-rite compared with actinolite, which is typical ofmetabasic rocks (Plyusnina, 1983).
22 Data on sulfur isotopes are discussed in Chapter 5.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S259
However, some results of our simulation cannot beunambiguously correlated with natural observations.The mineral assemblages of the downwelling limboften contain minor amounts of aluminous minerals:kaolinite and pyrophyllite in assemblage I and wairak-ite and sericite in transitional assemblages. These min-erals were identified in deep-sea drill core samples(Kurnosov, 1986), and there is no apparent conflictbetween their presence in the simulated assemblagesand natural observations. However, their appearancecould be related to a flaw in the model, the lack of alu-minous end-members in the compositions of chloriteand tremolite. It is conceivable that, if the multisysteminvolved mixed-layer smectite/chlorites and actinolitecontained several percent of alumina, these phaseswould not appear in the simulated products. The accu-racy of our thermodynamic model is insufficient toresolve these details.
(2) The minerals obtained by the simulation of oredeposition (Section 4.3.1) correspond to those observedin ore bodies (Table 3). The mineralogy of solid phasesfrom the quenching scenario of rapid mixing (Fig. 30)is identical to that observed in the smoke of submarinevents (Feeley et al., 1990). The model reproduces evensuch an indicator feature as the presence of pyrrhotitein the smoke, although it is absent in sulfide edifices.The model of disequilibrium mixing shows a predomi-nance of anhydrite in the precipitate (Fig. 34), which isin good agreement with the results of natural experi-ments reported by Tivey et al. (1990).
The simulation of the growing edifice scenario (Sec-tion 4.3.2) reproduces the development of oceanic sul-
fide edifices (Fig. 24), with respect to both the zoningpatterns of sulfide bodies and stages of their develop-ment. This model yielded zinc enrichment during theinitial stage of sulfide edifice development, which ischaracteristic of oceanic hydrothermal ores, and Zn andPb redeposition in the peripheral part of ore structuresduring their prolonged evolution. The formation ofmassive pyrite ores with relatively high copper contentsin the central part of edifices is consistent with the con-clusions derived by Krasnov (1993) from the data onthe mineral and chemical compositions of oceanic sul-fide ores.
(3) The composition of the hydrothermal solution isthe most sensitive and complex indicator of the correct-ness of the model. The data presented in Section 4.1.1show that the model adequately reproduces all the mainchanges in the composition of natural hydrothermalsolutions relative to the initial seawater: enrichment inK, Ca, Si, SII, Fe, chalcophile elements, dissolvedhydrogen, and methane; depletion in Mg; and adecrease in pH (recalculated to 25°C). A comparison ofdata presented in Table 19 with the chemical character-istics of oceanic hydrothermal solutions (Table 6) indi-cates a qualitative and quantitative correspondencebetween the model and natural solutions.
A detailed inspection (Table 21) reveals that thechemical compositions of solutions from 21° N EPR(and similar solutions from other sites) correspond tothe model solutions obtained in cases with low wavenumbers, i.e., with excesses of fresh basalts. It is inter-esting to note that a number of components (K, Ca, Zn,and S) showed a closer match with wave no. 2 rather
Table 21. Comparison of model results with the data obtained from some natural objects
Solutioncomponent Unit
Object Model calculation Object Model calculation
EPR, 21° N,HG, T = 351°C
T = 370°C,P = 500 bar, waveno. 1, R/W = 0.732
T = 370°C,P = 500 bar, waveno. 2, R/W = 0.517
MAR, 26° N, TAG, T = 366°C
T = 360°C,P = 500 bar, wave
no. 53, R/W = 0.027
No. in Table 6 4 – – 35 –
Cl mmol/kg 496 545.9 545.9 636 545.9
Na " 443 441.7 468.0 557 472.6
K " 23.9 37.8 21.5 17.1 6.92
Ca " 11.7 32.1 27.1 30.8 32.85
Mg " 0 0 0 0 0.001
Fe " 2.429 2.527 2.249 5.59 1.16
Zn μmol/kg 104 57.3 89.2 46 15.3
Cu " 44 1.41 3.82 120–150 74.5
Pb " 0.359 0.349 0.535 0.11 0.267
SiO2 mmol/kg 15.6 14.17 13.46 20.75 15.0
H2S " 8.4 12.39 6.58 3.5 0.661
pH(25°C) 3.3 3.51 3.45 3.35 3.82
S260
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
than with wave no. 1 (Table 21). The smoker solutionsof the TAG field are similar with respect to Zn and Curelationships to those from a long-lived system model(Table 21, simulation for wave no. 53). Thus, the resultsof simulation are in reasonable agreement with natural
observations with respect to a number of indicators.23
(4) The inversion of Zn and Cu proportions in themodel ore-forming solutions is of special importancefor the interpretation of geochemical and metallogenicdata on oceanic ores. It is therefore important to esti-mate whether this result is of general or local signifi-cance. According to the model, the inversion is relatedto two phenomena: (1) faster Zn removal from a younghydrothermal system compared with Cu and (2) anincrease in Cu concentration in the solution during theevolution of the system. The former phenomenon iscontrolled by the relative solubilities of Zn and Cu sul-fides. The experiments by Seyfried and Janecky (1985)conducted under similar temperature and chemical con-ditions yielded a Zn enrichment relative to Cu(Table 10). In our model Zn is incorporated in sulfidesonly, whereas part of the Zn in natural metasomaticrocks may occur as an isomorphous admixture in alu-minosilicates. Nonetheless, the investigation of a mod-ern hydrothermal system at 21° N on the EPR sug-gested that the removal of Zn is 3–4 times more effi-cient than that of Cu (Von Damm et al., 1985).
According to the simulation results, the increase inCu concentration during the evolution of the system iscontrolled by the rapid removal of SII from basalts,which increases the solubility of Cu sulfides. The mainmineral repository of sulfur in basalts is pyrrhotite,which is much more soluble than chalcopyrite underhydrothermal conditions (Fig. 22a). Rapid disappear-ance of magmatic sulfides during basalt alteration wasrepeatedly documented (Gitlin, 1985). The efficiencyof SII removal from the hydrothermal systems of 21° NEPR is 30–60 times that of Cu removal (Von Dammet al., 1985). On the other hand, data on the composi-tion of epidosite from the Troodos ophiolites suggestthat up to 90% Cu may be removed from a stronglyaltered rock (Richardson et al., 1987). The apparentdiscrepancy between these data on modern and ancientsystems is easily explained by the simulation results:the removal of Cu is intensified in long-lived systems.
23 A better agreement between the calculated solution compositionand a particular natural object can probably be achieved by therefinement of model parameters (primarily, sulfur content in thebasalts under investigation) and accounting for the influence ofphase separation due to boiling (Chapter 6). Therefore, the datapresented here have no direct implications for the inverse prob-lem, i.e., the determination of characteristics of natural objects.
The above considerations suggest that the changefrom Zn > Cu to Cu > Zn is not an accidental effect ora specific feature of the model but is caused by thephysicochemical properties of compounds of these ele-ments. Consequently, it must be systematicallyobserved in evolving hydrothermal systems fed by sea-water. The analysis of the available data on oceanicdeposits shows that solutions from most of them haveZn > Cu (Table 6). However, the investigations of anore-forming solution from a long-lived hydrothermalsystem in the TAG field (Edmond et al., 1995) revealedCu concentrations higher than those of Zn (Table 6,analysis 35; Table 21) and higher than in any other sys-tem studied. The calculated enrichment of copper in thecentral part and zinc in the peripheral part of a matureore body is in agreement with the results of drilling in asulfide mound at the TAG field (Hannington, 1998).
4.5. Conclusions
The agreement between the results of the thermody-namic modeling of oceanic hydrothermal systems andnatural data with respect to many mineralogical andgeochemical indicators indicates that the model devel-oped in this paper is realistic. This offers the possibilityof using the model for the interpretation and predictionof some properties of natural hydrothermal systems.
The simulation of oceanic hydrothermal systemsdemonstrated that an increase in the lifetime of a hydro-thermal system is favorable for a relative copper enrich-ment in the resulting ore body. On the other hand, long-lived systems form large ore bodies. Consequently, acorrelation is expected between the size of sulfide edi-fices and their copper enrichment. Such a correlationwas observed in nature (Hydrothermal Sulfide…, 1992;Krasnov, 1993): copper enrichment was documented inlarge sulfide mounds from the TAG field and the Gal-apagos Spreading Center. The ore bodies of these areashave very long lifetimes (repeated reactivation is alsopossible) compared with the hydrothermal edifices ofthe East Pacific Rise. Thus, the reason (or one of themain reasons) for copper enrichment in large long-livedsulfide structures on the ocean floor is the extensivemetasomatic alteration of rocks within the system.
What are the geologic factors favorable for this pro-cess? One of the obvious factors is the prolonged circu-lation of hydrothermal solutions within a single blockof oceanic crust, which is possible in a slow-spreadingridge near a periodically replenished magma chamber.Such a situation occurs in the TAG region. Anotherfavorable setting is probably an off-axis volcanic cen-ter, which also possesses a magma chamber fixedwithin the oceanic crust. Copper enrichment wasreported from the sulfide bodies of off-axis volcaniccenters (Alt et al., 1987; Krasnov, 1990; Fouquet et al.,1996).
Steady-state magma chambers probably existbeneath fast-spreading mid-ocean ridges. Prolonged
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S261
hydrothermal circulation (continuous or intermittent)therefore occurs in these structures, and the hydrother-mal systems of fast-spreading ridges are formally alsolong-lived. The obvious differences of such systemsfrom those described above are the rapid withdrawal ofcrustal rocks from the zone of convective circulationand the continuous supply of fresh basalts owing todike emplacement and crystallization at the roof ofmagma chambers. Hence, the hydrothermal activity inany given crustal block appears to be shorter and meta-somatic alterations are less extensive than in slow-spreading ridges. There is a persistent excess of freshbasalts within the hydrothermal systems. In particular,this is suggested by the calculation of oxygen isotopebalance (Merlivat et al., 1984) and the isotopically lightcomposition of sulfide sulfur (Chapter 5).
Investigations of ancient objects in the well-knownSemail ophiolitic complex (Oman) showed that largemassive copper sulfide bodies are related to off-axisvolcanic centers (Haymon et al., 1989), and the orebodies are associated with abundant epidosites (Nehliget al., 1994).
The results of our simulation indicate fundamentalchanges in the behavior of many components depend-ing on the R/W value. This property of our modelimplies that it does in fact reproduce the distinctionbetween the fluid-dominated and rock-dominatedregimes of water–rock interaction (Fyfe et al., 1981).The transition between these regimes probably occursclose to the boundary between model assemblages Iand II (Fig. 22). Its calculated position at R/W ≈ 0.03 isconsistent with available experimental estimates for thebasalt–seawater system (Seyfried and Mottl, 1982). Upto R/W ≈ 0.03, the geochemical characteristics of thesteps of the model reactor are controlled mainly by theinput of Mg, SO4, and O2 with seawater; and theremoval of components from basalts becomes of primeimportance at higher R/W values.
The change in the behavior of ore elements near theboundary of the stability fields of metasomatic assem-blages I and II implies that this boundary (and the adja-cent region) is a geochemical barrier. According to thecalculations, the concentrations of many componentsare affected by this barrier, but the most pronouncedchanges are observed for variable-valence elements, H,C, S, Fe, and Cu (Figs. 21c, 21d, 23). The main drivingforce of this barrier is a solution–basalt redox reaction,which reduces SO4 to H2S, H2CO3 to CH4, and H2O toH2. FeII of the basalt is simultaneously oxidized tooccur as FeIII in epidote. Among the ore metals consid-ered, Cu is most sensitive to this geochemical barrier.The movement of this barrier during the developmentof the hydrothermal system results in the redepositionof copper sulfides, which are oxidized near the frontand deposited in the back part. The high Cu concentra-tion in the solution before the barrier (Fig. 21d) pro-motes a gradual copper accumulation on the barrier.The maximum Cu content in the ore-forming solutions
of the upwelling limb is attained when this barrierapproaches the focus of the hydrothermal system(Figs. 27b, 38).
CHAPTER 5. THERMODYNAMIC MODELSOF ISOTOPIC CHEMICAL SYSTEMS
5.1. Method of H. Ohmoto and Its Applicationin Thermodynamic Models
for the Hydrothermal Process
Isotopic geochemistry provides one of the mostpowerful tools for geochemical investigations. Uniqueresults were obtained using isotopic methods, includingthe identification of sources of hydrothermal solutionsand ores. However, when isotopic indicators are usedwithout nonisotopic ones, the interpretation of resultsmay often be ambiguous. Therefore, modern stable iso-tope geochemistry relies on a combined use of isotopicand geochemical information.
The idea that physicochemical factors in a reactionenvironment can influence the distribution of isotopesbetween reactants and, consequently, that isotopicratios are not universal signatures of sources in geo-logic processes was first advanced by Sakai (1968).Ohmoto (1972) developed this idea and proposed amethod for the calculation of isotopic variations in sys-tems with chemically reacting substances. AlthoughOhmoto (1972) and Ohmoto and Rye (1979) appliedthis method to systems with sulfur and carbon, it is aversatile tool, which has subsequently been used totackle a variety of geochemical problems (Bannikovaand Ryzhenko, 1984; Bowers and Taylor, 1985; Jan-ecky and Shanks, 1988; etc.).
Ohmoto proposed calculation of isotopic composi-tions in a closed system with chemical reactions in twostages. The first stage deals with the chemical composi-tion of the system; its aim is to calculate the concentra-tions of substances participating in isotope exchangereactions. In simple situations (Ohmoto, 1972), theseconcentrations are defined by the occurrence of a sim-ple chemical reaction or change in external parameters.The concentrations of substances in complex systemsare calculated by the methods of thermodynamic mod-eling using appropriate computer programs (Bannikovaand Ryzhenko, 1984; Bowers and Taylor, 1985). Dur-ing the second stage, the calculated concentrations ofchemical compounds are used to evaluate the equilib-rium distribution of isotopes among them. This secondstage was developed by Ohmoto.
In the general form, the Ohmoto method is reducedto the solution of the following system of equations:
(a) isotope balance equation
(33)
where mi are the molar amounts of substances involvedin the isotopic exchange, νi are the stoichiometric coef-
miνi δi δΣ–( )i
n
∑ 0,=
S262
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
ficients of the element whose isotopic exchange is con-sidered, δi are the isotopic compositions of these com-pounds (shifts relative to an isotope standard), δΣ is thebulk isotopic composition of the system, and n is thenumber of compounds involved in the isotopicexchange;
(b) (n – 1) equations for isotopic exchange equilib-rium:
(34)
where δa is an arbitrarily chosen substance from thosepresent in the system, and αi – a is the isotope fraction-
ation coefficient between the substances i and a.24
The system is linear with respect to the unknown δiparameters, and its solution is straightforward. Ohmoto(1972) solved the problem explicitly by substituting δiexpressed from the exchange equation through δa andlnα into the balance equation. Bowers and Taylor(1985) and Janecky and Shanks (1988) designed spe-cial computer subroutines for the solution of isotopicproblems, which were implemented in the EQ3/6 pro-gram for the calculation of chemical equilibria. A sim-ilar isotopic subroutine was designed by the author as amodule in the GBFLOW (v.3.1) program.
In his pioneering study, which has been extensivelycited in the geochemical literature (e.g., Faure, 1986),Ohmoto calculated variations in the isotopic composi-tions of sulfur and carbon depending on the intensiveparameters of the hydrothermal process, and quantita-tively evaluated the influence of chemical reactions onthe isotopic compositions of components in hydrother-mal solutions (Ohmoto, 1972).
Various geochemical problems were approachedusing the Ohmoto method. Bannikova and Ryzhenko(1984) studied the carbon and sulfur isotope composi-tions of compounds of the C–H–O–S system. Theyrevealed significantly nonlinear dependencies of theisotopic compositions on the initial bulk composition ofthe system. These authors constructed alignment chartson the basis of simulations and correlated the isotopiccompositions of fluids with their chemical characteris-tics, in particular, with the degree of oxidation. Theseresults were used by Bannikova (1990) for thegeochemical reconstruction of hydrothermal ore for-mation.
The most sophisticated combined isotopic andgeochemical model was constructed by Bowers andTaylor (1985) for the convective hydrothermal systemof a mid-ocean ridge (see Section 3.3). They calculatedthe oxygen and hydrogen isotopic evolution of a hydro-thermal solution in a ten-component chemical multi-system including more than 30 minerals. They showed
that the isotopic shifts δ and δ are not largein a convective system, reaching +2.6‰ and +2.0‰,
24 Bowers and Taylor (1985) used a more rigorous equilibriumcondition, δj = (1000 + δi)αij – 1000.
δi δa– 1000 αi a– , i a≠( ),ln=
DH2O O18H2O
respectively, which was in agreement with the directmeasurements by Wehlan and Craig (1983) for hydro-thermal systems at 21° N on the EPR. The model bulkisotopic compositions of secondary minerals develop-ing after basalts are also reasonably consistent with theisotopic systematics of altered ocean-floor basalts andophiolitic rocks.
Janecky and Shanks (1988) derived an isotopicgeochemical model based on sulfur isotopes for theprocess of ore deposition in oceanic hydrothermal sys-tems. Its chemical part is identical to the model of oredeposition with mixing developed by Janecky and Sey-fried (1984) on the basis of the titration concept (Sec-tion 3.3). These authors assumed that the hydrothermalsolution entering from the convective system has a con-
stant δ of +1‰ and demonstrated by calculationthat no process which accompanies mixing with seawa-ter and reactions with the country rocks is capable of
increasing δ above +4.5‰ (observed in someoceanic hydrothermal systems, see Section 5.3). Inorder to explain the heavy sulfur isotopic composition,they invoked additional reduction of marine sulfate spe-cies by interaction with basalts within the feeder chan-nel (i.e., the subsurface mixing of solutions with differ-ent histories).
The use of the Ohmoto method in the aforemen-tioned and other studies allowed a comprehensive inter-pretation of isotopic data and in some cases providedisotopic geochemical criteria for the reconstruction ofhydrothermal processes. The development and applica-tion of this method contributed significantly to the gen-eral geochemical theory of hydrothermal processes.
5.2. Method of an Isotopic Chemical System
Ryzhenko proposed an alternative approach to thecalculation of isotopic equilibria in systems with chem-ical reactions. The basic idea underlying his method isthat both chemical reactions and isotopic exchange aregoverned by the fundamental principle of chemicalthermodynamics—a tendency toward the minimumfree energy of the system. If isotopes rather than ele-ments are used as independent components, thermody-namic equilibrium can be simultaneously calculated forchemical reactions and isotope exchange reactions.
Based on this approach, Bannikova et al. (1987) andGrichuk (1987, 1988) developed a method for the ther-modynamic calculation of isotopic and chemical reac-tions, which is referred to as the method of isotopicchemical systems. In contrast to the two-stage calcula-tion procedure by Ohmoto (chemical equilibrium fol-lowed by isotopic equilibrium), the method of isotopicchemical systems achieves the desired goal in a singlecalculation step. Such equilibria can be calculatedusing standard programs designed for the solution ofnonisotopic thermodynamic problems.
S34H2S
S34H2S
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S263
It should be noted that the method of isotopic chem-ical systems is theoretically more rigorous than theOhmoto method. The validity of the two-stage compu-tational procedure relies on the negligible influence ofisotope redistribution on equilibrium in chemical reac-tions. The advantages and disadvantages of the Ohmotomethod and the method of isotopic chemical systemsare analyzed below.
The implementation of the method of isotopicchemical systems requires knowledge of the thermody-namic characteristics of the components of the system,the Gibbs free energies of the isotopic forms of com-pounds, ΔfG0I. The existing fundamental handbooks onthe thermodynamic properties of substances (ThermalConstants…, 1965–1979; Thermodynamic Proper-ties…, 1978–1982; IVTANTERMO, 1983–1985) pro-vide systematic data only for the compounds of hydro-gen isotopes, 1H, 2D, and 3T. For other elements, there areonly practical thermodynamic functions for the com-pounds of natural isotopic mixtures. They do not accountfor nuclear constituents, including those related to themixing of various isotope forms. The method of isotopicchemical systems must include a means for calculating thefree energies of formation of isotopic forms of compoundsfrom the available experimental data.
This problem was addressed by Grichuk (1987),who derived a general equation for the calculation ofΔfG0I from the known practical function, free energy offormation of compounds, ΔfG0E, and isotope fraction-ation coefficients between the given compound and theelement in a standard state, α. For the arbitrary com-pound Z composed of y chemical elements, the freeenergies of fully substituted forms, under the condi-tions of independent isotopic exchange between ele-ments in a Z molecule, can be expressed as
(35)
where ny is the number of atoms of y in Z, ly is the num-ber of isotopes of element y, Nyjst is the average abun-dance of the isotope in nature, and ϕZ and ψ are correc-tions for possible variations in the isotopic compositionof the compound.
For the relevant case of two isotopes of the ele-ment considered (1H–2D, 12C–13C, 14N–15N, 16O–18O,and 32S–34S), Eq. (35) can be significantly simplified.For the element X of the compound AX, consisting oflight (X0) and heavy (X*) isotopes:
Δf = Δf + NX*RTlnαAX (36a)
Δ f GZyj0I Δ f GZ
0E RT ny Nyjst αZlnyij( )
j
ly
∑y
∑+=
– RT ϕZ nynystψy
∑+⎝ ⎠⎜ ⎟⎛ ⎞
,
GAX
00I GAX
0E
and
Δf = Δf + RTlnαAX, (36b)
where and NX* are the average relative abundances
of the isotopes, and αAX is the isotope fractionationcoefficient for AX relative to the standard state of X.
The equations derived above are valid for both theGibbs free energies of the isotopic forms of com-
pounds, ΔfG0I, and the reduced free energies, .
Isotopic exchange between geologically importantcompounds has been studied for only a few elements(H, C, O, and S). Data on the isotopic fractionation ofmost major elements are scanty. Therefore, most of theproblems addressed in geochemical studies concernisotopic fractionation with respect to one or a few ele-ments occurring in the system. Such problems can bereferred to as semi-isotopic to discriminate them fromfull isotopic problems accounting for all isotopes of thechemical elements occurring in the system. The calcu-lation of the thermodynamic properties of isotopicforms of compounds for such semi-isotopic systemspresents no considerable difficulties, because isotopicexchanges of various elements (except for hydrogen)occur practically independently. The coupling of isoto-pic exchanges of hydrogen and other elements is a sec-ond-order effect, and its magnitude is comparable withthe errors of the experimental and theoretical determi-nation of α. Such problems can be resolved usingEqs. (36a) and (36b) (e.g., Bannikova et al., 1987). Theresults of thermodynamic calculations for the semi-iso-topic (sulfur) model of a convective hydrothermal sys-tem in the oceanic crust are discussed in Section 5.4.
Comparison of the methods of Ohmoto and isotopicchemical systems. The method of isotopic chemicalsystems considers isotopes as independent componentsof the system. The basic principles of chemical thermo-dynamics are therefore applicable, and, correspond-ingly, standard programs can be used for the simulationof such systems. As was mentioned above, the Ohmotomethod involves two stages: chemical equilibrium iscalculated in the first stage, and isotopic equilibrium iscalculated in the second stage. The first stage can beperformed using standard programs for thermodynamicmodeling, whereas special program modules areneeded for the second stage (e.g., Bowers and Taylor,1985; Janecky and Shanks, 1989). Each of the twomethods has its own advantages and limitations.
Method of isotopic chemical systems. Any programof thermodynamic modeling can be used for the simu-lation of isotopic chemical models of homogeneoussystems (liquid or gas). However, geochemicallyimportant problems usually involve isotopic exchangein heterogeneous systems (gas–solid or aqueous solu-tion–solid phases). In such a case, each solid phase par-ticipating in the isotopic exchange must be consideredas a solid solution of isotopic forms (isotopic end-mem-bers). The simulation of such systems requires pro-
GAX*0I GAX
0E NX
0
NX
0
gT0I
S264
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
grams for the calculation of equilibria in systems withseveral phases of variable composition. Such programsbelong to the highest complexity class according to thenomenclature by Shvarov (Methods of…, 1988) and arefew in number (GIBBS, SELECTOR, and HCh). Thisimposes technical limitations for use of the isotopicchemical systems method.
The input information for the method is the isotopicratios of the initial system and the thermodynamicproperties of isotopic forms of compounds. The initialisotopic ratios are defined by geological models andtheir quality is not discussed here. The thermodynamicproperties of isotopic compounds can be calculatedfrom ordinary (elemental) thermodynamic parametersand the fractionation coefficients of isotopes using theequations derived in the previous section. This can bedone without any principal difficulties using thermody-namic data banks similar to UNITHERM, because onlyΔfG0 (298 K) values and two or three coefficients ofheat capacity equations (depending on the form of thetemperature function of lnα) have to be changed. Theshortcoming of this approach is that the calculation of iso-topic exchange for several elements (e.g., H, O, and C)involves a considerable number of isotopic forms thatmust be accounted for. The thermodynamic data bankswill be expanded correspondingly. In this respect, thepotential of the method of isotopic chemical systemsappears to be somewhat excessive for the present-daystatus of the thermodynamic modeling of geologic pro-cesses.
The accuracy of calculations by this method is con-trolled by the errors of the input thermodynamic dataand uncertainties imposed by the computational pro-gram. The errors of thermodynamic data are propa-gated from (a) errors in elemental thermodynamicproperties and (b) errors in isotope fractionation coeffi-cients (obtained experimentally or calculated), whichare usually between n × 10–4 and (1–2) × 10–3 (lnα). Itshould be noted that the assumption used for the deri-vation of Eqs. (35) and (36) do not affect the quality ofthe calculation of isotope fractionation. The computa-tional error for the simulation of equilibria in theGIBBS, HCh, and GBFLOW programs is no higherthan 10–5 in the molar fractions of components, whichcorresponds to 0.01‰ for the equilibrium isotopiccomposition of compounds.
It can be seen that the main source of error in themethod under question is ordinary thermodynamicdata. The internal sources of error are several orders ofmagnitude smaller. Correspondingly, the computa-tional errors are smaller than the experimental uncer-tainties of lnα.
The Ohmoto method involves the development ofspecial isotopic blocks (subroutines) extending anyprogram for the calculation of chemical equilibria. Theprogram realization of these blocks is straightforward.The input isotopic data are not connected with the ther-modynamic properties of substances in the Ohmoto
method, and there is no need to include them in basicthermodynamic data banks. This facilitates the use ofthis method for the solution of particular geochemical
problems. The Ohmoto method utilizes the same (orequilibrium constants) and lnα values as the method ofisotopic chemical systems. Consequently, the errors ofthese parameters contribute similarly to the results ofcalculations by the both methods.
The Ohmoto method is theoretically less rigorousthan the method of isotopic chemical systems becauseof a number of inherent simplifications.
(1) The two-stage calculation procedure in the Ohm-oto method is based on the implicit assumption that iso-topic exchange does not influence chemical equilibria,i.e., there are no coupled isotopic–chemical effects. Ingeneral, this is in agreement with empirical data.
(2) In the system of Eqs. (33) and (34), the massconservation equations for each isotope (which areexactly satisfied in the method of isotopic chemical sys-tems) are replaced by the balance equation of isotopicshift (Eq. (33)).
(3) The derivation of Eq. (34) in the Ohmoto methodis based on the approximation of natural logarithmsln(1 + δ) ≈ δ.
The inspection of errors related to these assump-tions suggests that they are usually negligible. Theerrors in the equilibrium constants of chemical reac-tions related to assumption (1) (ignoring coupled isoto-pic–chemical effects) are described for each atom par-ticipating in isotopic exchange by the following equa-tion (Grichuk, 1998):
(37)
where X* is the fraction of the less abundant isotope,and X0 is the fraction of the more abundant isotope. Forgeochemically important reactions, this effect is nolarger than n × 10 J/g-atom and usually much smaller.It is much smaller than the typical errors of elementalthermodynamic parameters.
The maximum systematic error in the results of cal-culation of isotope fractionation due to assumption (2)is given by the following equation for the isotopicexchange between two compounds:
(38)
The magnitude of this systematic shift is no higherthan 0.0n‰ for all light elements (except for Li and B),and it can be ignored in geochemical applications.
Assumption (3) introduces an error in the calcula-tion of isotopic compositions specified by the equation
ξδ ≈ δavlnαij, (39)
where δav = (δi – δj)/2. If δav is strongly different fromthe isotopic standard, this assumption may lead to con-
gT0
δ pKr( ) 12303------------X* α X*
X0-------⎝ ⎠
⎛ ⎞ ,lnln≈
εδ–( )max103
4-------- X*
X0-------⎝ ⎠
⎛ ⎞ αln( )2 ‰( ).=
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S265
siderable errors, which are higher than the confidencelimits of empirical constants used in the models (up to3–9‰). It is desirable, therefore, to use the Ohmotomethod with the exact equation of isotopic exchange(instead of Eq. (34)) in the form
in combination with the mass conservation conditionfor each isotope, or in another form:
as was proposed by Bowers and Taylor (1985).Taking into account these refinements, it can be con-
cluded that the Ohmoto method and the method of iso-topic chemical systems give practically identical resultsfor geochemical problems, and the choice of a particu-lar procedure depends on practical circumstances,including the availability of software complexes andconvenience of their application. For instance, the twomethods were successfully used by Bannikova (1990).
Sulfur isotopic systematics are among the most sensi-tive indicators in the geochemistry of hydrothermal pro-cesses, because sulfur is a major ore element undergoingcomplex chemical transformations which are recorded inthe distribution of its isotopes. There are more than30 publications on the investigation of the sulfur isotopiccompositions of modern oceanic sulfides.
Early studies (Styrt et al., 1981; Arnold and Shep-pard, 1981) reported δ34S values for oceanic sulfidesbetween +1 and +4‰, and it was argued that their sul-fur had a mixed origin: most sulfur is derived by theleaching of sulfides from basalts (δ34S ≈ 0‰), and asmaller portion is produced by the reduction of seawa-ter sulfate species within the convective system (δ34S ≈+20‰). This interpretation was in agreement with themost popular concept of the source of sulfur in massivesulfide deposits (Ohmoto and Rye, 1979) and wasbased on the experimental data in the seawater–basaltsystem then available.
An alternative hypothesis for the genesis of sulfur inmassive sulfide deposits is sulfur isotope fractionationduring the reduction of marine sulfate sulfur (Sasaki,1970; Kajiwara, 1971). It appeared to be completely atodds with the data obtained for smokers. This hypothe-sis was based on an isotopic fractionation of between−21 and –17‰ for the SO4–FeS2 pair at 300–350°C. Insuch a case, the reduction of marine sulfate would haveproduced sulfides with near zero isotopic signatures,corresponding to those observed in massive sulfideores. However, such δ34S values of sulfides could beobtained only if there is an excess of sulfate sulfur. If
αij
Xi*X j0
Xi0X j*
-------------=
αij
δi 1000+δ j 1000+-----------------------,=
the inventory of sulfates is limited, a Rayleigh distilla-tion must occur and the resulting sulfur isotope compo-sition of the sulfides must become heavier, approachingthat of initial sulfate at complete reduction. The ques-tion on the content of SO4 in the zone of ore formationof ancient massive sulfide deposits has not yet beenclarified. The complete reduction of sulfates was estab-lished for smokers from the chemical compositions ofemanating solutions. Consequently, the Sasaki–Kaji-wara hypothesis cannot explain the isotopically lightcomposition of sulfides.
Figure 42 shows a generalized distribution of isoto-pic compositions in all oceanic hydrothermal objects.Although the extent of sampling varies considerablybetween the sites, the data set is still representativeowing to the great number of objects studied(845 determinations were processed). The diagramshows that the average δ34S of modern marine sulfidedeposits is +4.41‰ (Fig. 42a). The distribution of iso-topic compositions is distinctly asymmetric with amode of δ34S = +3.0‰.
The δ34S of anhydrite and barite from modernhydrothermal systems ranges between +8 and +22‰
–2
50
–1 0 1 2 3 4 5 6 7 8 9 10 11 12
60
70
80
40
30
20
10
0
N
(‡)
14 16 18 2012108
6
8
4
2
0
10
(b)
22
δ34S, ‰
Fig. 42. Isotopic compositions of sulfide and sulfate sulfurfrom oceanic sulfide edifices (excluding the GuaymasBasin). The sources of data are given in Fig. 43. (a) Sulfides,the dashed line shows the average δ34S value for the dataset; and (b) sulfates (barite and anhydrite).
S266
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
–2
N
δ34S, ‰0 2 4 6 8 10 12
10
0
10
0
50
10
0
10
0
5
0–2
δ34S, ‰0 2 4 6 8 10 12
10
0
5
0
10
0
10
0
10
0
10
0
5
15
5
15
10
0
5
15
10
0
5
15
10
0
5
15
15
5
20
25
1015202530
20
15
5
10
15
20
5
5
–2δ34S, ‰
0 2 4 6 8 10 12
10
15
20
(a) EPR, 21° N
(j) Juan de Fuca Ridge, Middle Valley
(b) 11°–13° N EPR
(k) Juan de Fuca Ridge, Escanaba Trough
(c) Green Seamount, 13° N EPR
(l) Galapagos Spreading Center
(d) 9°40′ N EPR
(m) Northern Lau Basin
(e) Juan de Fuca Ridge, southern segment
(n) TAG, 23° N MAR
(f) Juan de Fuca Ridge, Axial Seamount
(o) 14′45′ N MAR
(g) Mariana Trough
(h) Manus Basin
(p) Guaymas Basin
10
0
(i) Snake Pit, 23° N MAR
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S267
with a mode of 21‰ (Fig. 42b). This is interpreted asthe result of formation of the main mass of sulfate min-erals by mixing of hydrothermal solutions with seawa-ter.
A comparison of the sulfur isotope data for individ-ual hydrothermal fields (Fig. 43a–43p) reveals somegeneral relationships.
(1) In the majority of sampled edifices, sulfide min-erals are isotopically disequilibrated (Styrt et al., 1981;Zierenberg et al., 1984). Attempts to calculate isotopictemperatures yielded unrealistically high values (450–700°C), and some data could not be interpreted at all.25
(2) In some edifices later sulfides appeared to havehigher δ34S (e.g., at 21° N on the EPR, data by Woo-druff and Shanks, 1988).
(3) Intriguing results were obtained from a compar-ison of the δ34S values of sulfide minerals and H2S ofhydrothermal solution (Figs. 43a, 43b, 43e, 43f, 43g,43k, 43m). In the majority of hydrothermal vents,hydrogen sulfide is isotopically heavier than sulfidesfrom the edifice (Bluth and Ohmoto, 1988). The oppo-site relationship was reported only from three of27 sampled hydrothermal vents.
(4) The isotopic compositions of sulfides from thevents of the Guaymas Basin, whose hydrothermal sys-tem develops mainly in a sedimentary sequence,appeared to be systematically lighter than in other sys-tems, and even lighter than magma-derived sulfur(Peter and Shanks, 1992) (Fig. 43p). The histogram ofisotopic compositions illustrates their distinct poly-modal distribution. This is interpreted as resulting fromthe contribution of bacterial sulfides mobilized fromsediments to the formation of the isotopic composi-tions. However, this effect has not been observed inother similar systems (Escanaba and Middle Valley)where solution circulation entrains sediments(Figs. 43j, 43k).
(5) A comparison of data on various hydrothermalsites showed that sulfides from large hydrothermalstructures are isotopically heavy (compare Figs. 43a–43i and 43j–43o). This phenomenon was pointed out by
25 The minerals that must be isotopically lighter than pyrite inequilibrium, for instance, bornite, pyrrhotite, and sphalerite,appeared to be isotopically heavier, which can be exemplified bythe data from 21° N EPR (Woodruff and Shanks, 1988) and AxialSeamount and Middle Value of the Juan de Fuca Ridge (Hanning-ton and Scott, 1988; Goodfellow and Franklin, 1993).
Lein et al. (1991), Grichuk and Lein (1991), and by theauthors of Hydrothermal Sulfide… (1992).26 When allthe available data are grouped into two sets, small edi-fices and large structures (the latter includes sulfidebodies from the TAG, Middle Valley, Escanaba, andGalapagos Spreading Center with masses of n × 105–n × 106 t), their average δ34S values are +3.19 and+5.75‰, respectively. The distributions appear to bealmost normal (Figs. 44a, 44b).
The isotopic heterogeneity of modern oceanichydrothermal deposits has long intrigued researchers,and a number of hypotheses have been proposed toexplain this phenomenon (see Bluth and Ohmoto, 1988for a review). They invoked different origins of the iso-topic signatures of hydrothermal solutions and differentplaces of the generation of variations in sulfide compo-sition.
The sulfur isotope compositions of sulfides fromoceanic deposits were first reported by Styrt et al.(1981), who supposed that the observed isotopic varia-tions reflect different proportions of sulfur leachedfrom basalts and marine sulfur reduced to hydrogensulfide during water–basalt interaction. This hypothesiswas further developed by Bluth and Ohmoto (1988),who argued that the proportion of sulfur sourceschanges with time: the fraction of SII reduced from sul-fate increases and, as a result, the isotopic compositionof dissolved hydrogen sulfide is systematically heaviercompared with that of the sulfide body. These authorsexplained local variations in the composition of sulfidesby the isotopically disequilibrium processes of sulfidereplacement in the chimney walls of smokers. Thus, atany given time, each hydrothermal system shows itsown δ34S value of hydrothermal solution depending onthe water–rock interaction within it. Grichuk (1988)and Grichuk and Lein (1991) developed an isotopicchemical thermodynamic model corresponding to thishypothesis. The structure and outcomes of this modelare discussed in detail in Section 5.4.
An alternative hypothesis was proposed by Shanks
and Seyfried (1987). They supposed that isreduced during seawater circulation in basalts through
26 Sulfides with very high δ34S were recently reported from a smallsulfide edifice in the northern Lau back-arc basin (Fig. 43m).
SO42–
Fig. 43. Histograms of sulfur isotope compositions of oceanic sulfides. Arrows show the isotopic compositions of hydrogen sulfidefrom smoker solutions. Data sources: (a) 21° N EPR (Styrt et al., 1981; Arnold and Sheppard, 1981; Kerridge et al., 1983; Zieren-berg et al., 1984; Woodruff and Shanks, 1988); (b) 11°–13° N EPR (Bluth and Ohmoto, 1988; Hydrothermal Sulfide…, 1992);(c) 13° N EPR, Green Seamount (Alt, 1988); (d) 9°40′ N (Hydrothermal Sulfide…, 1992); (e) Juan de Fuca Ridge, southern segment(Shanks and Seyfried, 1987); (f) Juan de Fuca Ridge, Axial Seamount (Hunnington and Scott, 1988); (g) Mariana Trough (Kusakabeet al., 1990); (h) Manus Basin (Lein et al., 1993); (i) Snake Pit, 23° N MAR (Kase et al., 1990); (j) Juan de Fuca Ridge, MiddleValley (Goodfellow and Blaise, 1988; Goodfellow and Franklin, 1993); (k) Juan de Fuca Ridge, Escanaba Trough (Koski et al.,1988; Zierenberg et al., 1993); (l) Galapagos Spreading Center (Skirrow and Coleman, 1982); (m) northern Lau Basin (Bortnikovet al., 1993); (n) TAG field, 23° N MAR (Lein et al., 1991; Hydrothermal Sulfide…, 1992; Herzig et al., 1998b); (o) 14°45′ N MAR(Bogdanov et al., 1997a); and (p) Guaymas Basin (Koski et al., 1985; Lein et al., 1988; Peter and Shanks, 1992).
S268
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
the conversion of magmatic pyrrhotite to pyrite via thereaction
7FeS + 8H+ + 4FeS2 + 4H2O +3Fe2+.
They supposed also that the newly formed pyrite hasa mixed isotopic composition, which was estimated as+2.9‰ by isotope balance calculations in accordancewith the stoichiometry of this reaction. It was also sug-gested that the hydrothermal solution has to be in isoto-pic equilibrium with this pyrite, which constrains itsδ34S = +2.1‰ at 370°C. This value must characterizehydrothermal solutions produced within any oceanichydrothermal system. The logic of this hypothesisimplies that the only factor affecting the δ34S value ofsolution within the system is temperature: the lower thetemperature, the higher the fractionation coefficient αfor the pyrite–hydrogen sulfide pair, and the lighter theisotopic composition of hydrogen sulfide (however, thedegree of fractionation is rather low: at 300°C the δ34Sof solution in the interior part of the system is +1.7‰).
It can be easily seen that this scenario is based onseveral arbitrary assumptions. The pyrrhotite–pyriteconversion is not the only possible mechanism of sul-fate reduction during solution–basalt interaction. Bothexperimental data and calculations suggest that thereduction occurs with concurrent transformations of
SO42–
iron-bearing aluminosilicates (FeII of pyroxenes andolivine to FeIII of epidote). Shanks and Seyfried (1987)mentioned a δ34S value of between +2.9 and +5.0‰ forvein pyrite from DSDP Hole 504B (Honnorez et al.,1985). However, this vein pyrite was precipitated fromsolutions ascending in a stockwork zone at a relativelylow temperature (about 200°C). The subsequent analy-ses of disseminated pyrite from the rocks of the dikecomplex (Alt et al., 1989) yielded much lower δ34S val-ues (+0.69‰ on average) (Fig. 45c), which are onlyslightly higher than that of the initial monosulfides(+0.2‰ on average, Fig. 45a). Many analyses of oce-anic sulfides (more than 15% of the data set, Fig. 42a)have δ34S values below a critical value of +2.1‰.
In a later study, Janecky and Shanks (1988) dis-missed the idea of the pyrrhotite–pyrite conversion andpostulated that the isotopic composition of an ascend-ing hydrothermal solution is close to that of magmaticsulfides, δ34S = +1.0‰. The negligible contributionfrom marine sulfate into the isotopic composition of
40
50
60
70
30
20
10
0
N
(‡)
0 1 2 3 4 5 6 7 8 9 10 11 12–1–2
40
50
60
70
30
20
10
0
δ34S, ‰
(b)
Fig. 44. Histograms of the isotopic compositions of sulfidesfrom ore structures of (a) small and (b) large sizes. Thedashed lines show average values for the data sets.
543210–1–2–3–4–5
6420
N (‡)
–20 –16 –8–12 –4 0 4
4
6
2
8
10
0
(b)
543210–1–2–3–4–5
6
4
2
0
8
10
12 (c)
543210–1–2–3–4–5
246
0
(d)
δ34S, ‰
Fig. 45. Selected results of the investigation of oceanicbasalts recovered from ODP Hole 504B in the Costa Ricarift. (a) Monosulfides from the basalts of the dike complex(Alt et al., 1989); (b) disseminated pyrite from the transi-tional zone (Belyi et al., 1984); (c) disseminated pyrite fromthe rocks of the dike complex (Alt et al., 1989); and (d) veinpyrite (Kawahata and Shikazono, 1988; Alt et al., 1989).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S269
solutions generated within the hydrothermal systemwas explained by the removal of sulfate from the solu-tion by anhydrite precipitation at the beginning of thedownwelling convective flow. According to the modi-fied hypothesis, isotopic variations are already gener-ated in the ascending channel beneath ore bodies andwithin ore bodies at the expense of mixing with seawa-ter and/or anhydrite reduction. The different magni-tudes of δ34S increase in sulfides from various hydro-thermal systems are explained by variable mixing pro-portions.
As was mentioned in Section 3.1, the scales of sub-surface mixing cannot be reliably constrained from thedata on solution chemistry, because the currentlyaccepted method of analysis recalculation involvesextrapolation to Mg = 0. The fraction of seawater (X)required, according to the Shanks hypothesis, for sucha mixing can be estimated by chemical and isotopic bal-ances calculated using equations similar to Eq. (33):
Co = Cd(1 – X) + CwX
and
Coδo = Cdδd(1 – X) + CwδwX,
where the subscripts d, w, and o refer to the deep hydro-thermal solution, seawater, and the values observed insmokers, respectively; C is the concentration of sulfurin the solution (hydrogen sulfide in hydrothermal solu-tions and sulfate in seawater). This system can bereadily transformed to
XCo
Cw------
δo δd–δw δd–----------------.=
Following the model of Shanks, let us accept δd =+1.0‰, δw = +20‰, and Cw = 28.23 mmol/kg. Then, Xcan be calculated for any given δo and Co. Such calcu-lations for some sampled vents are shown in Table 22.It can be seen from these results that the observed vari-ations in the isotopic composition of hydrogen sulfideimply a relatively small fraction of seawater in the mix-ture, only 0.25–6.0%. However, even such small addi-tions of seawater should have resulted in the precipita-tion of considerable amounts of magnesium silicates(talc and serpentine, Section 4.3.1) comparable with theamounts of newly formed sulfides, which is not the casein the objects studied.
The isotopic chemical simulation carried out by Jan-ecky and Shanks (1988) showed that the disequilibriumreduction of sulfates during mixing provides δ34S val-ues of solutions no higher than +2.5‰, and sulfidesdeposited from them have δ34S values no higher than+4.5‰. A comparison of these results with natural data(Figs. 42a, 43) shows that the mixing model does notexplain the observed diversity of isotopic composi-tions. Janecky and Shanks proposed that a more exten-sive process of sulfate reduction must occur in thefeeder zone with the participation of hydrothermalsolutions, fresh basalts, and anhydrite. The supposedreduction of isotopically heavy sulfate can provide anyincrease in the δ34S of solution discharging on theocean floor. Thus, the Shanks hypothesis implies thatthe isotopic variations are related to factors that areaccidental with respect to processes in the interior ofhydrothermal systems (subsurface mixing and anhy-drite occurrence in the channel). Therefore, the exist-ence of the aforementioned persistent relationships in
Table 22. Estimation of the fraction of seawater in mixtures according to the Shanks hypothesis
Hydrothermal system Sampledobject Reference
,
mmol/kg, ‰ X, %
Juan de Fuca Ridge, Cleftsegment, 45° N
Plume [Shanks and Seyfried, 1987] 0.72 5.6 0.66
" 2.617 4.2 1.66
" 0.2 7.3 0.25
vent 1 0.544 6.4 0.58
Juan de Fuca Ridge,Escanaba Trough, 41° N
– [Zierenberg et al., 1993] 1.1 7.8 1.48
EPR, 21° N NBS [Woodruff and Shanks, 1988] 3.8 3.4 1.80
OBS 7.4 1.5 0.73
OBS 6.9 1.3 0.41
SW 4.0 3.5 1.98
HG 3.9 2.3 1.00
HG 3.0 3.2 1.31
EPR, 13° N vent 6 [Bluth and Ohmoto, 1988] 8.2 4.7 6.00
vent 8 4.5 3.3 2.05
vent 10 8.0 4.6 5.70
cH2S δH2S
S270
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
the isotopic composition of sulfides cannot beexplained by this hypothesis.
As possible indirect indicators, the Janecky–Shanksmodel predicted a decrease in copper content in the sys-tem with increasing δ34S and a decrease in δ34S withdecreasing temperature, which must result in an expan-sion of the δ34S range toward isotopically light compo-sitions in large hydrothermal structures. These authorsnoted that the modern ore-forming systems of mid-ocean ridges are too small, too young, and too poorlysampled to display such effects (Janecky and Shanks,1988, p. 820). The data of subsequent studies were notcompatible with these predictions. Large hydrothermalsystems are relatively enriched in copper (Section 3.1)and show the heaviest isotopic compositions of sulfides(Figs. 43i–43l). The Shanks hypothesis does notexplain the systematic increase in δ34S of dissolvedhydrogen sulfide relative to the sulfides of hydrother-mal edifices.
Bowers (1989) presented isotopic chemical modelsfor the downwelling limb of a hydrothermal convectionsystem. She considered four types of models. Types 1and 2 reproduced isothermal seawater–basalt interac-tion by the method of the degree of reaction progress,i.e., in a closed system. Types 3 and 4 simulated poly-thermal step flow reactors with different R/W distribu-tions between temperature steps. Owing to the closedcharacter of the systems (Section 2.1), marine sulfurwas not removed from the system in isothermal modelstype 1 and 2, and δ34S of hydrogen sulfide increasedaccordingly to +15–17‰ at R/W ≈ 0.2. δ34S decreasedto +5.2 and +8.8‰ at higher R/W = 2 owing to the dilu-tion by sulfur from basalts. Since it is clear that such amodel (closed system with back-reactions) is notappropriate for convective systems, Bowers carefullyinterpreted these results. The mineral assemblages andcompositions of solutions obtained in type 3 and 4models are very similar to those of the model of Bowersand Taylor (1985). These calculations suggest that sul-fate is exhausted at a step of 250°C with concurrentanhydrite precipitation and, to a smaller extent (15–40% of initial content), is reduced to H2S. The δ34S ofhydrogen sulfide and sulfide minerals increases in thisstep up to +(7–10)‰. The dissolved sulfur is diluted bymagmatic sulfur during the following steps, and δ34Sdeclines rapidly in the solution to +0.5‰.
Proceeding from the results of these calculations,Bowers concluded that the reduction of marine sulfatewithin hydrothermal systems does not exert a stronginfluence on the isotopic composition of solutions andδ34S < +1‰ at R/W > 1. Thus, she supported the con-clusions of Janecky and Shanks (1988) on the near-zerosulfur isotopic composition of solutions produced inthe interior parts of hydrothermal systems. As to thereasons for the systematically higher δ34S values of thesulfides and solutions of smokers, Bowers alsoaccepted the explanation of Janecky and Shanks andrejected the idea of Bluth and Ohmoto (1988). Similar
conclusions were derived by Kusakabe et al. (1990)from a study of hydrothermal systems in the MarianaTrough and, with some variations, by Herzig et al.(1998) for the Valu Fa Ridge.
The entrainment of isotopically heavy sulfate intothe ore process through subsurface mixing was activelydebated in connection with the results of drilling in theTAG hydrothermal field. The participants of this projectaccepted the Shanks model (e.g., Herzig et al., 1998b).A detailed evaluation of their data is beyond the scopeof our study. Note only that the main differencebetween the Shanks and Bluth–Ohmoto hypotheses isrelated to the place where the sulfur isotopic composi-tion of fluid becomes heavier. According to the Bluth–Ohmoto hypothesis, this occurs in the downwellinglimb of the system, and the feeder channel contains iso-topically heavy hydrogen sulfide. According to Shanks,hydrogen sulfide entering the channel is isotopicallylight and its δ34S increases through a reaction withanhydrite or subsurface mixing. In the TAG field, theisotopically heaviest pyrite (δ34Sav = +8.0 ± 0.7‰) wasfound in the lowest parts of boreholes, below the zoneof anhydrite occurrence (Knott et al., 1998; Gemmelland Sharpe, 1998). Pyrite associating with anhydrite islighter (δ34Sav = +5.6 ± 0.7‰). In an attempt to recon-cile these facts with the Shanks hypothesis, the afore-mentioned authors suggested a mixing of two fluids: afluid of deep circulation with δ34S ≈ 5.5‰ (this ismerely the minimum value for sulfides from pyrite–anhydrite veins and is higher by 1‰ than the maximumvalue admitted by the model of Janecky and Shanks,1988) and a hypothetical fluid of subsurficial circula-tion with isotopically heavy hydrogen sulfide (δ34S ≈9.0‰). How can this mixing be correlated with the hightemperature of the fluid emanating from edifices? Howcan the shallow convection cell, which is supplied bycold seawater, be heated up to the temperature of deepcirculating fluid? The fact is that isotopically heavypyrite was found in the same zones where the highesttemperatures were recorded (Petersen et al., 1998).
It can be easily seen that all these complex scenariosproposed by the authors cited stemmed from their con-viction that a fluid with δ34S higher than +4.5‰ cannottheoretically be derived in the downwelling limb of aconvection cell.
Summarizing the current state of the problem ofinterpretation of oceanic sulfur isotope data, we con-clude that a relatively simple and clear-cut qualitativemodel was proposed by Bluth and Ohmoto (1988).According to this model, sulfur from oceanic sulfideores has a mixed (magmatic + marine) origin and themixing proportion changes in time toward increasingcontribution from the marine source. The local varia-tions in isotopic compositions within edifices arerelated to the isotopic disequilibrium processes of min-eral replacement. In principle, this scheme provides anexplanation for all the observed relationships. How-ever, the published quantitative models for the genera-
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S271
tion of sulfur isotopic compositions in oceanic hydro-thermal systems are based on an alternative scenario,which cannot explain all the persistent relationships inthe isotopic compositions of oceanic sulfides.
5.4. Isotopic Chemical Model for a Convective Hydrothermal System
Using the method of isotopic chemical systemsdescribed in Section 5.2, the author developed a quan-titative model for the behavior of sulfur isotopes in thedownwelling limb of an oceanic convective system(Grichuk, 1988; Grichuk and Lein, 1991; Hydrother-mal Sulfide…, 1992). The chemical and dynamicaspects of this model are identical to those described inchapters 3 and 4. The processes of sulfur isotope redis-tribution occurring within the ore body are not consid-ered in this model, because the available data suggestthat they are strongly disequilibrated.
The main process reproduced by the isotopic chem-ical model is the mixing of marine sulfur (δ34S =+20.0‰) with sulfur extracted from basalts by solu-tions (δ34S = + 0.3‰; Sakai et al., 1984). The mixingis complicated by isotope fractionation between dis-solved sulfur species and sulfur-bearing mineralsand a portion of the sulfur being expelled from min-
erals not in contact with solution, due to the transfer ofsolution portions (waves) to the next step of the flowreactor.
In order to simulate this model, the dimensionalityof the chemical multisystem described in Section 3.4.2was increased to 16 components through the separateconsideration of 32S and 34S. The expanded set of spe-cies included seven 34S isotope forms, and the sulfur-bearing minerals of the multisystem were regarded asideal solid solutions of isotopic end-members formingeight series: anhydrite, pyrite, pyrrhotite (troilite),galena, sphalerite, chalcocite, bornite, and chalcopy-rite. The thermodynamic properties of the isotopicforms were calculated from Eqs. (36a) and (36b) using
elemental values from the UNITHERM databaseand αT values taken from Table 23. Equilibria in isoto-pic chemical systems were calculated by the GIBBSprogram designed by Shvarov.27
Figures 46a and 46b show the calculated chemicaland isotopic compositions of sulfur compounds in thedownwelling limb of the convection cell for a case withΣR/W(400°C) = 1.26 (wave no. 1). These results sug-
27 These calculations were repeated using the GBFLOW v.3.1 pro-gram based on the Ohmoto method.
gT0
Table 23. Reference values of 103lnαAX–BX used for calculations; (after Ohmoto and Rye, 1979)
Substance A C Estimated error, * Note
S (g) –0.16 0 ±0.5 Consistent with Grinenko and Thode (1970)
CaSO4 5.26 6 4–2 Similar to estimates by Friedman and O’Neil (1977) and Golyshev et al. (1983)FeS2 0.4 0 0.4–0.1
FeS 0.1 0 0.3–0.1
PbS –0.63 0 0.3–0.1
ZnS 0.1 0 0.3–0.1 Consistent with compilation by Friedmanand O’Neil (1977)
Cu2S –0.75 0 0.5–0.15 Approximate estimate
CuFeS2 –0.05 0 0.4–0.15
Cu5FeS4 –0.25 0 0.5–0.15 Approximate estimate
5.26 6 4–2 =α for CaSO4, anhydrite
Ca
Mg
Na
H 6.3 1.9 4–2 After Igumnov et al. (1977)
H2S – – –
HS– –0.06 –0.6 0.4–0.2
* The uncertainty of α decreases with increasing T.
103 αi-H2S g( )ln
A
T2----- 106 C+×=
SO42–
SO40
SO40
SO4–
SO4–
S272
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
gest that within the stability field of assemblage I, theisotopic signature of sulfate sulfur remains almost con-stant. Hydrogen sulfide in equilibrium with the sulfateshows very low δ34S values (but its concentration isvery low). The concentration of sulfur in the solutiondecreases owing to anhydrite precipitation. In theregion of transitional assemblages, the sulfate ionremaining in solution is reduced. During this processthe δ34S of hydrogen sulfide increases up to positivevalues, and that of residual sulfate may be as high as+40.7‰ owing to isotope fractionation. The isotopiccompositions of sulfides are similar to that of H2S, andthe difference between them is in accordance with the
fractionation coefficients in the mineral–hydrogen sul-fide system. An increase in R/W in the stability field ofassemblage II results in the entrainment of sulfur fromigneous rocks, which lowers the δ34S of solutions to≈+0.5‰. These results were obtained by Grichuk(1988) and supported by the independent calculationsof Bowers (1989) for a similar model. The bulk isotopiccomposition of solution (δ34SΣ) is summed over variouschemical forms and appears in the diagram (Fig. 46b)as a characteristic curve with a logistoid-like shape.
A comparison of the results corresponding to vari-ous R/W values (Figs. 47a, 47b) suggests that the tem-
0
0.005
0.010
0.015
0.020
0.025
Conce
ntr
atio
n, m
ol
Anh
Py
S(VI)
S(II)
220150 260 300 340 380Temperature, °C
–20
–10
0
10
20
30
40
50
δ34S
, ‰
Assemblage IIAssem-
Transitional region
Bulk system
Anhydrite
Sulfides
Bulk solution
S(II)
S(VI)
(a)
(b)
blage I
Fig. 46. Calculation of an isotopic chemical model for sulfur in the downwelling limb of a hydrothermal system at ΣR/W(400°C) =1.26 (first wave). (a) The concentrations of major sulfur species and (b) the isotopic compositions of sulfur compounds.
1
10–1 100
R/W, kg/kg
5
010–3
δ34S, ‰
2
3
4
5
6
7
0101
Transitional region
(b)Final R/W values
in series0.126
0.42
0.84
1.26
2.110
15
20
25
30
35
10–2 10–1 100 101
R/W, kg/kg
(a)δ34S, ‰
Fig. 47. Variations in the sulfur isotope composition of solution as a function of the cumulative R/W ratio calculated for variableintensity of interaction. (a) The dependency of isotopic composition on R/W and (b) the same for large R/W.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S273
perature of the system and R/W distribution in themodel steps exert a noticeable influence on the form ofthe dependency of δ34SΣ on ΣR/W in a hydrothermalsolution. During a weak water–rock interaction(ΣR/W(400°C) < 0.8), the curves of the isotopic com-positions of solutions practically coincide, whereas astronger interaction significantly increases the bulkδ34S of solution in the zone of transitional assemblages(Fig. 47a). The analysis of the simulation shows thatthis effect is related to the appearance of pyrite in thesolid phase. In the presence of pyrite, the amounts ofreduced and oxidized sulfur species become compara-ble, and isotopic fractionation between pyrite and dis-solved sulfates results in a sharp increase in the δ34S ofsolution. This increase is also evident in Fig. 46b. Thestability field of pyrite in the downwelling limb isshown in Fig. 22a. It can be seen that the pyrite field isintersected by flow lines calculated for extensivewater–rock interactions. The larger the pyrite stabilityfield in the flow reactor, the more pronounced theresults of fractionation. The scale of δ34S increase in thebulk isotopic composition of solution depends ondynamic relationships rather than on the temperaturedependencies of fractionation constants: in the caseswith ΣR/W (400°C) of 0.84, 1.26, and 2.1 (Fig. 47a), themaximum δ34SΣ values are attained at almost equaltemperatures of 250, 240, and 230°C, respectively. Themaximum δ34SΣ values are significantly different inthese cases, whereas the fractionation coefficient isalmost constant within such a narrow temperature inter-val: lnα(pyrite–sulfate) is –0.02376 at 250°C and−0.0252 at 230°C.
The influx of magmatic sulfur from the rock into thereaction volume increases at high R/W. As a result,δ34SΣ decreases gradually, and the transitional zoneeffects are strongly mitigated. Figure 47b presents amagnified portion of this diagram for R/W > 0.1. In thestability field of mineral assemblage II, the curves ofthe sulfur isotope composition of solution form a nar-row band. The weak dependence of δ34SΣ on R/W dis-tribution in the temperature steps is a favorable prop-erty of the curve, because this distribution cannot nowbe reliably determined in natural systems and is esti-mated in the model considered by theoretical methods(Section 2.2).
A surprising feature of the model is the pressuredependency of δ34SΣ in the interior part of the system(Grichuk, 1988) (Fig. 48). This effect is also related tothe chemistry of the process rather than to the thermo-dynamics of isotope exchange (our model uses pres-sure-independent fractionation coefficients). The rea-son resides in the strong pressure sensitivity of anhy-drite solubility (Blownt and Dickson, 1969). Thesolubility of this mineral controls the concentration ofresidual marine sulfate affected by reduction. Thehigher the solubility, the larger the contribution of iso-
topically heavy marine sulfur, and the gentler the slopeof the δ34SΣ dependency.28
The correlation of the δ34SΣ value of solution withΣR/W (Fig. 47b) makes it possible to independentlydetermine ΣR/W from data on the isotopic compositionof oceanic hydrothermal solutions. The analysis of theerrors of the model (Grichuk, 1988) yielded an esti-mated accuracy of ΣR/W of ±0.2 logarithmic units.However, the ΣR/W value obtained from Fig. 47b forthe best studied hydrothermal system at 21° N on theEPR (average δ34S = +2.59 for 129 determinations,Fig. 43a) gives ΣR/W = 0.15–0.20. This is significantlydifferent from the ΣR/W values obtained by other meth-ods (0.90–0.55, Tables 11, 12). The reason for this dis-crepancy is that the curve in Fig. 47b was constructedfor the model of the first wave, i.e., for a very younghydrothermal system.
In order to assess the influence of time, we simu-lated a multiwave scenario corresponding to the evolu-tion of a long-lived hydrothermal system. Figure 49shows variations in the isotopic composition of sequen-tial solution portions (waves) for a model with initialΣR/W = 1.26. This diagram shows a distinct progressiveincrease in δ34SΣ during the lifetime of the system. Thereason for the increase in δ34S of the model solution isobvious. The influx of isotopically heavy marine sulfurinto the convective system is constant, whereas that ofisotopically light magmatic sulfur declines in accor-dance with the accepted model dependency of R/W on
1/ (Eqs. 2.10, 2.11). Thus, the isotopic compositionof the solution becomes heavier in response to thedeceleration of the diffusion metasomatic process con-trolling water–rock interaction and sulfur extractionfrom the basalts. This is fully consistent with the con-clusion of Bluth and Ohmoto (1988) that the increase inthe δ34S of smoker solutions is related to the formationof altered layers (armoring) on the walls of fractures,through which seawater circulates.
28 Pressure estimates for the interiors of convective systems werediscussed in Section 3.2.
τ
10–210–3 10–1 100
R/W, kg/kg
0
5
10
15
20
25
30δ34S, ‰
500 bar
250 bar
Fig. 48. Sulfur isotopic composition of solution as a func-tion of pressure in the downwelling limb of the convectivesystem.
S274
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
The chemical evolution of the system in the zone ofassemblage I is additionally marked by a decrease inthe rate of Ca extraction from the basalts. Correspond-ingly, less marine sulfate species are deposited in theinitial steps of the reactor, and the sulfate ion can pene-trate deeper into the system (Section 4.1.2). In contrast,the extraction of isotopically light SII is depressed withdecreasing current R/W values in the reactor steps. Thecombined action of these two factors results in the nearparallel displacement of δ34SΣ curves in Fig. 49.
In isothermal sections of the system (reactor steps),δ34SΣ increases gradually with time, i.e., with increas-ing number of waves (Fig. 50a). The simulation of this
process revealed the flaw of the method of a step flowreactor that was discussed in sections 2.1.4 and 4.1.2: ajump of a boundary between mineral assemblages fromone step to another is accompanied by oscillations inthe sulfur content of the solution, which is reflected inthe calculated isotopic compositions. This is illustratedby small variations in both the isotopic compositions ofsulfur species and the bulk composition of solution inFig. 50a. Nonetheless, it can be clearly seen in this dia-gram that the isotopic variations are separated into twostages:
(a) δ34S of hydrogen sulfide increases systematicallywith time from +0.61‰ in the first wave to +16‰ inwave no. 530, δ34S of sulfate increases from +20‰ to+34‰, and the bulk isotopic composition of solution isidentical to that of hydrogen sulfide; and
(b) after wave no. 531, the δ34S of solution decreasesabruptly to almost zero, and that of sulfate decreases to+18‰, and the bulk composition of the solution contin-ues to change gradually and approach the compositionof sulfate.
A comparison of this diagram with Fig. 27a sug-gests that the isotopic composition changes abruptlywhen the oxidizing sulfate solution reaches the givenreactor step. It is characteristic that the isotopic compo-sition of sulfates does not become identical to that ofthe initial seawater: the sulfates are isotopically lighterowing to an admixture of sulfur extracted from basaltsin previous steps.
After passing through the high-temperature reactorsteps, the solutions are introduced through the ascend-ing channel into the ore deposition zone, where a sul-fide edifice gradually forms. Its average isotopic com-position will correspond to the bulk composition ofhydrogen sulfide of all sequential solution portions.29
Figure 50b illustrates the calculation of the cumulativeisotopic composition of SII emitted from the 370°Creaction step (focus of the system). It can be seen that,at a small number of waves, the bulk sulfur isotopic
29 Ignoring the loss of hydrogen sulfide during the early stages ofblack smoker development and anhydrite reduction in the sulfidebody.
Fig. 49. Variations in the δ34S of solution for the model of a long-lived hydrothermal system with initial ΣR/W (400°C) = 1.26.
S(II)S(VI)Bulk S of solution
–505
10152025303540 (a)
1011 201 301 401 501 601 701 801Wave number
H2S in solution
Cumulative
–202468
1012141618
δ34S
, ‰
(b)
Fig. 50. Variations in the isotopic composition of solution inthe 370°C isothermal section obtained by the simulation ofa long-lived hydrothermal system with the initial ΣR/W =0.732. (a) The isotopic compositions of sulfur species and(b) the cumulative curve of the isotopic composition ofhydrogen sulfide removed from the downwelling limb ofthe system.
δ34S
, ‰
(instantaneousvalue)
mean value
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S275
composition of the sulfide edifice ranges between +1and +4‰. With the increasing lifetime of the system,the δ34S of the edifice will increase (in the case consid-ered, up to +10‰). The quantitative parameters of iso-topic evolution of the sulfide body will depend on theΣR/W value reached in the system, because it affects thecompositions of parent solutions. The general tendencyof a δ34S increase with time will be retained.
In addition to the cumulative curve, Fig. 50b showsthe isotopic composition of hydrogen sulfide in a cur-rent solution portion, i.e., the instantaneous δ34S value.A comparison of these curves suggests that, owing tothe regular evolution of the system, the composition ofthe sulfide edifice, which is a time-integrated character-istic, will inevitably be delayed relative to the variationsin solution composition, which has the latest and, cor-respondingly, the highest δ34S value. Such relationshipsare observed in most hydrothermal systems (Figs. 43a,43b, 43e, 43g, 43k).
It should be noted that some hydrothermal edificesshow evidence for a reverse evolution trend: a decreasein δ34S with time (Woodruff and Shanks, 1988; Bluthand Ohmoto, 1988). Taking into account the results ofsimulation, such a trend can be interpreted as a conse-quence of the rejuvenation of the hydrothermal systemby some tectonic events, when considerable amounts offresh basalt come into contact with the solution (forinstance, owing to rock fracturing).
Verification of the model. A comparison of theresults of simulation with the data obtained for naturalcomplexes shows an overall consistency between themodel and natural processes.
The data on the secondary alteration of oceanicbasalts show that magmatic sulfur is rapidly mobilizedby hot seawater. This is indicated by the absence of pyr-rhotite in the model, except for the region of very highR/W. The upper part of the crustal section penetrated byHole 504B contains pyrite with a strongly variable iso-topic composition, from –18.3 to +0.5‰ (average−3.1‰, Fig. 45b) (Belyi et al., 1984). Sulfides from theunderlying dike complex are isotopically heavier(Figs. 45c, 45d): pyrite from stringers shows an averageδ34S of +3.15‰ (from 0 to +5.8‰), and disseminatedpyrite, +0.7‰. This is correlated with the abruptincrease in δ34S of hydrogen sulfide and sulfides fromnegative to positive values obtained in the model for thetransitional zone (Fig. 46b).
Anhydrite from stringers in basalts shows a δ34Svalue similar to that of seawater, from +20 to +22‰,which is in good agreement with the calculations for thestability field of assemblage I (left part of Fig. 46b).The model predicts a heavier isotopic composition ofsulfates in the transitional zone. It is difficult to assessthe adequacy of this model effect to natural processes,because an increase in the δ34S value of sulfates occursdirectly before their complete disappearance. Nonethe-less, sulfates with δ34S of up to +36.7‰ were found indrill cores from Hole 504B in the Costa Rica Rift (Alt
et al., 1983), which is in agreement with the model val-ues for the beginning of the assemblage II zone(Fig. 46b).
In general, the isotopic composition of edificesapproximately corresponds to the composition of theirparent solutions. The calculations imply that oceanichydrothermal systems must produce solutions with asmall positive δ34S value (Figs. 46–50). This result is inagreement with observations in natural solutions(Figs. 42, 44a).
The magnitude of an increase in δ34S dependsmainly on ΣR/W and rises over the lifetime of the sys-tem (Fig. 50). In the context of the model considered,the observed diversity of δ34S values in sulfides fromoceanic hydrothermal systems can be explained byvarying degrees of maturity of the hydrothermal sys-tems, i.e., different extents of transformation of thecrustal materials by hydrothermal processes. Accordingto this criterion, the hydrothermal systems of 21° N,11°–13° N, and 9° N on the EPR; Cleft segment andAxial Seamount on the Juan de Fuca Ridge; and SnakePit on the MAR, having an average δ34S of sulfide bod-ies from +2 to +4‰ (Figs. 43a–43g), are classified asyoung or short-lived hydrothermal systems containingexcess fresh basalts in their interior parts. The hydro-thermal vents of the TAG field, Galapagos SpreadingCenter, and, presumably, Middle Valley and Escanabatroughs with average δ34S values higher than +6‰(Figs. 43j–43l, 43n) are classified as long-lived sys-tems, whose deep parts were extensively altered byhydrothermal processes. This result provides a simpleexplanation for the correlation mentioned in Section5.3 between the isotopic compositions of sulfide sulfurand the size of sulfide edifices (Figs. 44a, 44b). It is alsoin agreement with the results of simulation of the chem-ical composition of solutions and ores (Chapter 4) andgeological evidence for these hydrothermal sites.
The isotopic compositions of sulfides from the off-axis sulfide occurrences of the EPR correspond in gen-eral to the range obtained by our simulation: their aver-age δ34S values are +4.3‰ for Green Seamount (21° NEPR) (Alt, 1988) and +1.8‰ for the marginal high andSE Seamount (13° N EPR) (Fouquet et al., 1996). In thelatter case, there is no correlation between the isotopiccomposition and size of edifices, but the data for suchobjects are still very limited.
Sulfides from the hydrothermal ores of theLogatchev and Rainbow fields, which sit on ultrabasicblocks, are characterized by heavy isotopic composi-tions (Bogdanov et al., 1997a; Lein et al., 2003). Amodel similar to that described above can probably beused for the interpretation of these data, taking intoaccount the specific features of seawater–peridotiteinteractions (cf. Alt and Shanks, 2003).
The available limited data on the isotopic composi-tions of sulfides from hydrothermal occurrences inisland arcs do not contradict the model describedabove. Samples of massive sulfides from the JADE field
S276
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
in the Okinawa Trough yielded δ34S values between+3.2 and +6.1‰ (Marumo and Hattori, 1999) andbetween +5.2 and +7.2‰ (Glasby and Notsu, 2003).
Sulfides from the hydrothermal fields of back-arcspreading centers yielded the following average δ34Svalues: +2.65 at 18° N in the Mariana Trough (Kusak-abe et al., 1990), +3.35 in the Manus Basin (Lein et al.,1993), and +4.2‰ in the Lau Basin in the northern ValuFa Ridge (Herzig et al., 1998a). In the context of themodel considered, these values are in agreement withthe small sizes of hydrothermal edifices observed inthese localities. Sulfides from the Vai Lili field in thecentral part of the Valu Fa Ridge show a heavier isoto-pic composition, δ34Sav = +8.6‰ (n = 27) (Bortnikovet al., 1993; Herzig et al., 1998). This could be due tothe high intensity of hydrothermal processes in thisfield, which lowered the ΣR/W values in the interiorpart of the system and accelerated the evolution of iso-topic ratios compared with other systems.
It should be noted that not all the available data areconsistent with the above-described model. Isotopi-cally light sulfides with δ34Sav = –5.06‰ (n = 13) werefound in the Hine Hine hydrothermal field (Lau back-arc basin) (Herzig et al., 1998). Negative values cannotbe obtained by mixing of marine and magmatic sulfur.In our model negative δ34S(H2S) values are generated inthe transitional region (Fig. 46) owing to fractionation
with ; however, sulfate must be isotopicallyheavier than seawater. Barite from this field showsδ34Sav = +16.4‰ (n = 5). Such relationships cannot beelucidated by the model developed.
Herzig et al. (1998) explained their data by theinflux of SO2-bearing magmatic gas into the Hine Hinehydrothermal system, which is situated in the caldera ofa submarine volcano. The disproportionation reaction
4SO2 + 4H2O H2S + 3H2SO4
is accompanied by the fractionation of sulfur isotopesand produces sulfides with peculiar isotopic character-istics. This explanation is supported by the unusualmineral composition of the edifice, where sulfides andbarite associate with pyrophyllite, kaolinite, alunite,and native sulfur. The numerical model makes no pro-vision for a contribution from magmatic gas, and itmust be modified to be applicable to such objects.
CHAPTER 6. SIMULATION OF BOILINGIN OCEANIC HYDROTHERMAL SYSTEMS
6.1. Role of Boiling in Ore Formation:Overview of the Problem
Boiling (more generally, phase separation in fluid)in hydrothermal systems is traditionally regarded as one of
SO42–
the main factors controlling their ore potential.30 Boilingcan influence ore deposition via several mechanisms.
Loss of complex-forming species. In the geochemi-cal literature, the most popular mechanism of the influ-ence of boiling on ore metal deposition is the removalof complex-forming ligands into the gas phase. Thedecomposition of complex species results in reactionsof the type
Zn(HS ZnS↓ + H2S↑. (40)
This mechanism can be exemplified by uranium pre-cipitation from carbonate complexes in response tosolution degassing and CO2 loss.
Loss of a precipitant. Some anion-forming ore com-ponents (primarily H2S and CO2) are highly volatileand rapidly escape into the gas phase during boiling.The loss of precipitants may result in the dissolution ofore minerals occurring in contact with a heterogeneousfluid via reactions of the type
ZnS + 2H+ Zn2+ + H2S↑. (41)
It is evident that the loss of precipitants due to boil-ing prevents ore deposition. Such a mechanism isimportant for chalcophile elements, whose precipitant,SII, readily escapes into the vapor phase as H2S. It isessential that if this mechanism operates, ore depositionis associated with vapor condensation, which is theopposite of boiling. This process was addressed byBychkov and Grichuk (1991) by the example of As andSb deposition in the hydrothermal system of the Uzoncaldera (Kamchatka).
Change in pH and Eh. The migration into the gasphase of substances containing variable-valence ele-ments and substances showing electrolytic dissociationin the aqueous phase may change the redox and acid–base states of the system, which, in turn, will affect thebehavior of ore components. This point can be illus-trated by the reaction
2Au(cr) + 2H+ 2Au+ + H2↑. (42)
A shift in pH and Eh has repeatedly been invoked asa possible reason for gold precipitation (Cunningham,1985; Cline et al., 1992). However, in contrast to theloss of complex-forming ligands, this mechanism mayexert different influences on the behavior of ore compo-nents: ore minerals may either precipitate or dissolve.
Among the compounds of variable-valence ele-ments occurring in hydrothermal solutions, those oflow valence states show higher affinities to the gasphase (H2S, CH4, and H2). Therefore, boiling usuallyshifts the redox state of hydrothermal solutions toward
30 Boiling and degassing can be regarded as particular cases of het-erogenization, the separation of a liquid into liquid and gasphases. Boiling refers to the situations when the main componentof the gas phase is H2O, and degassing implies that another vola-tile component is predominant in the gas phase. There is no fun-damental physical difference between the terms boiling, degas-sing, and phase separation.
)20
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S277
more oxidized states. This is favorable for the dissolu-tion of the most common minerals of some ore ele-ments (Au, Ag, and U), whereas minerals of other ele-ments (Fe and Mn) can be precipitated.
Some important components of hydrothermal solu-tions, including H2S and H2CO3, are characterized byacid-type dissociation. The partitioning of their dissoci-ated forms into the gas phase will increase the pH of theresidual solution:
HC CO2↑ + OH–. (43)
This effect is favorable for the precipitation of sul-fide, carbonate, and oxide ore minerals.
Loss of a solvent. The boiling off of liquid and anaccompanying increase in the mineralization of the liq-uid phases leads to mineral deposition. This is probablyone of the most important deposition mechanisms forsome components of hydrothermal solutions, forinstance, SiO2. Correspondingly, the opposite process,vapor condensation, produces a low-mineralized solu-tion, which may be aggressive with respect to the countryrock. In contrast, an increase in solution concentrationmay enhance complexation and result in dissolution ofsolid phases of those ore components that form complexdissolved species with a large number of ligands (prima-rily Fe, Zn, Cu, and Pb chloride complexes).
Temperature and pressure effects. In addition to purelychemical interactions, boiling is related to strong thermaleffects. During the discharge of a two-phase hydrothermalfluid, its temperature is constrained by the T–P boilingcurve. The ascent of such fluid toward the surface will beaccompanied by decompression and, consequently,strong cooling. The majority of ore components precipi-tate during cooling. Such P–T constraints were used, inparticular, as an important indicator of sea depth duringthe formation of massive sulfide deposits (Ridge, 1973;Finlow-Bates and Large, 1978; Krasnov, 1987).
In addition to the above-described mechanisms,boiling can probably affect the behavior of ore mineralsin other ways. For instance, the influence of the major-component composition of hydrothermal solutions onore components via changes in the dielectric constantof solution has recently been discussed in the geochem-ical literature (Kolonin et al., 1994; Akinfiev, 1994).The escape of CO2 into the gas phase is probablyaccompanied by such effects, but it is currently difficultto assess their scales in natural environments.
Competitive mechanisms. The complexity of theproblem of boiling is largely related to the diversity ofmechanisms through which important components ofhydrothermal solutions, such as H2S and CO2, canaffect the behavior of ore metals during boiling. Forinstance, H2S may be simultaneously a precipitant anda complex-forming substance. Since these mechanismsoperate in opposite directions, they compete for influ-ence, and the resulting effect cannot be determined apriori. It will depend on the conditions of boiling andthe composition of the system.
O3–
In some cases, theoretical analysis of such a compe-tition is instructive. For instance, if the prevailing zincspecies in a solution are free ions or chloride com-plexes, the boiling of the solution intensifies reac-tions (41) and (44):
ZnS + 2H+ + 2Cl– Zn + H2S↑. (44)
There is no ore deposition, and moreover, the oremineral will be dissolved. The process proceeds simi-larly in the case of the predominance of ZnHS+:
ZnHS+ + H+ Zn2+ + H2S↑. (45)
When Zn(HS is the major dissolved species, reac-tion (40) occurs and the ore phase will be depositedduring boiling.
A more complex problem was encountered in theanalysis of gold behavior during boiling (Drummondand Ohmoto, 1985). If there are no complex-formingsubstances, boiling promotes reaction (42), whichresults in gold dissolution. However, within a wide
range of conditions, Au(HS is the predominant spe-cies of hydrothermal solutions, and the loss of com-plex-forming components causes the reaction
(46)
Assuming that liquid pH and the composition of thereleased gas phase are constant, the equilibrium con-stant of reaction (46) can be written as
(47)
where Pi is the partial pressure, Ptot is the total pressure,and xi is the molar fraction of the ith component in thegas phase.
Thus, a pressure drop (enhancement of boiling)shifts the reaction to the right, toward gold precipita-tion. The loss of complex-forming components appearsto be more important than the change of redox state ofthe system. However, the assumption of constant pHand gas composition (especially the latter) degrades theimportance of this conclusion. Note also that a decrease inthe pH value of the system promotes gold dissolution inreaction (42) and deposition in reaction (46).
The influence of three of the aforementioned mech-anisms on ore matter is unambiguous: the loss of com-plex-forming components and adiabatic cooling pro-mote ore mineral deposition, whereas the loss of a pre-cipitant results in dissolution. The action of othermechanisms may be different depending on the identityof ore components.
Drummond and Ohmoto (1985) showed that pHbuffering by dissolved substances and/or solid phasescan also influence the behavior of ore components dur-ing boiling. In addition, these authors demonstrated
Cl20
)20
)2–
2Au(HS + H2↑ + 2H2O
2Au(cr) + 4H2S↑ + 2OH–.
)2–
KPH2S
4 aOH–2
PH2
---------------------xH2S
4 aOH–2
xH2
--------------------Ptot3 const × Ptot
3 ,= = =
S278
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
that, if the system is open to the vapor phase (which isremoved), the loss of highly volatile components ismost extensive during the initial stage of boiling. Cor-respondingly, the relative contributions of differentmechanisms may vary during boiling.
To take into account this diversity, one has to con-sider simultaneously a number of components, and thisexercise cannot be done “manually” at the level of the-oretical constructions. The investigation of such sys-tems is only possible by means of computer thermody-namic modeling.
Thermodynamic modeling of boiling of hydrother-mal solutions. The pioneering study of Drummond andOhmoto (1985) played a crucial role in the develop-ment of modern concepts on the connection betweenboiling and ore formation. These authors performedthermodynamic calculations of equilibria in a het-erophase system using a method based on the coeffi-cients of interphase partitioning (Henry constants) ofthe components. They demonstrated that open-systemboiling (with vapor phase removal) is more efficientwith respect to ore generation than the closed-systemprocess. They proposed geochemical indicators for theidentification of the products of ore deposition duringboiling.
A thermodynamic model with adiabatic boiling wasconstructed by Spycher and Reed (1989) using theBroadlands hydrothermal system as an example. They
demonstrated that the separation of volatile compo-nents changes the pH value of solution and, in combi-nation with temperature variations, results in the depo-sition of sulfides, sulfosalts, carbonates, and nativegold.
Bowers (1991) evaluated the influence of phase sep-aration in CO2-rich systems on the deposition of goldand base-metal sulfides. Her model was restricted to athree-component gas phase (H2O, CO2, and H2S), butthe liquid and solid phases were described by a com-plex 17-element multisystem. It was found that theeffect of boiling on the behavior of ore elements is dif-ferent in the presence and absence of a rock because ofthe buffering action of aluminosilicate minerals.
The aforementioned studies focused on the problemof ore deposition. The mobilization of ore matter duringsolution boiling was not addressed by these authors.Models of hydrothermal systems with ore dissolutionduring boiling were considered by Bychkov and Gri-chuk (1991) and Grichuk and Shvarov (1992). Someresults of the investigation of these models are dis-cussed below.
6.2. Boiling in Oceanic Hydrothermal Systems
The possibility of subsurface boiling in oceanichydrothermal systems was first evaluated by Delaneyand Cosens (1982), who used the phase diagram of sea-water (Fig. 51) to predict the formation of boiling zonesin the hydrothermal systems of seamounts and volca-noes at shallow depths (<1.6 km). The first naturalobservations for the existence of boiling zones (varia-tions in the total salinity of hydrothermal solutions)were reported from a hydrothermal system at 13° N onthe EPR (Michard et al., 1984).31 However, the oceandepth is more than 2.5 km in that site. According to thephase diagram of the H2O–NaCl system, boiling of asolution with a salinity of 3.5 wt % requires tempera-tures higher than 390°C at this depth. The observed val-ues were no higher than 350°C and prevented the inter-pretation of salinity variations resulting from phaseseparation. Various alternative explanations have beenproposed (e.g., Seyfried et al., 1986). The subsequentstudies of hydrothermal systems in the Juan de FucaRidge (Von Damm and Bischoff, 1987), especially inthe system of Axial Seamount (Massoth et al., 1989;Butterfield et al., 1990) clarified this problem. A num-ber of indirect lines of evidence for subsurface boilingwere found, including hydrothermal vents emittingsolutions less mineralized than seawater and enrichedin volatile components (CO2, H2S, etc.), which wereproduced by vapor condensation (Butterfield et al.,
31 There is an extensive literature on the problem of fluid boiling inthe deep levels of the oceanic crust, which involves the investiga-tions of gas–liquid inclusions (Vanko, 1988; Kelley and Delaney,1988; etc.). Since the correlation of these data with the structureof convective hydrothermal systems is not quite clear, they arenot considered in this study.
5004003002001000T, °C
4000
3000
2000
1000
Depth, m
1
2
3
Juan de Fuca Ridge
21° N EPR
Fig. 51. Analysis of the possibility of boiling in oceanichydrothermal systems on the basis of the phase diagram ofseawater (Delaney and Cosens, 1982). (1) Boiling curve forseawater, (2) adiabat for ascending hydrothermal solutions,and (3) minimum ocean depths for the EPR and the Juan deFuca Ridge.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S279
1990; Von Damm et al., 1997). Note, however, thatemanations of heterogeneous vapor–water jets weredirectly observed to date only in the shallow-water (400m) hydrothermal system of Grimsey near the northerncoast of Iceland (Hannington et al., 2001).
Table 24 shows solution compositions for a group ofclosely spaced vents at the ASHES hydrothermal field(Butterfield et al., 1990). The concentration of Cl– inthe Hell vent is almost identical to that of seawater, andthere was probably no phase separation within it. TheInferno vent is enriched in chlorides by 14% relative toseawater, which is indicative of the partial boiling off ofsolution. This vent is also significantly enriched in Feand Cu. The Crack and Virgin Mound vents are, in con-trast, depleted in chlorides, and their Cl– contents areonly 47 and 32% of the seawater value. This suggests acontribution from condensed water vapor to the solu-tions of these vents. These observations suggest that thedecrease in salinity is associated with an increase in theconcentrations of volatile components, CO2, H2S, andHe. The He/CO2 ratio is almost constant, which sug-gests a simultaneous separation of these componentsinto the gas phase. Similarly, the constant Br/Cl value(1.553 in seawater) reflects the conservative behavior oftheses halogens during phase separation. The freshenedvents are strongly depleted in chalcophile elements,especially Zn, probably owing to the high content ofhydrogen sulfide. Given this variability, of special inter-est is the lack of difference between the vents withrespect to B and SiO2 contents. Boron is rather volatile
under hydrothermal conditions (Ellis, 1979), and its rela-tive enrichment (compared with more conservative Cl) inthe Virgin Mound vent is geochemically understand-able. Similar boron contents in water vapor wereobtained in the experiments by Berndt et al. (1990).The relatively high content of silica in the freshenedvent suggests that the condensed water vapor probablyhad enough time to react with the fracture walls andextract SiO2 from them. Taking this fact into account,the similarity of pH values in high- and low-concentra-tion vents can be interpreted in different ways: eitherphase separation under the P–T conditions of theASHES hydrothermal system did not significantlyaffect the pH of solutions (in the experiments of Bis-choff et al., 1996, the pH of the vapor phase decreasedat the expense of hydrolytic reactions), or wall-rockreactions depressed the excess acidity of the conden-sates.
Von Damm and Bischoff (1987) proposed a three-component mixing scenario for the interpretation ofobserved variations in the chemical composition ofhydrothermal solutions from the southern segment ofthe Juan de Fuca Ridge. This scenario included the fol-lowing components: (1) boiled-off hydrothermal solu-tion, which is a brine with a salinity of about 3 mol/kgdepleted in volatile components; (2) vapor condensatewith a low mineralization and high content of volatiles;and (3) seawater introduced during sampling and bysubsurface mixing. The concentration of the brine wasestimated by these authors from the deficit of dissolved
Table 24. Compositions of hydrothermal solutions from ASHES hydrothermal field, Axial Seamount, Juan de Fuca Ridge(Butterfield et al., 1990)
Hydrothermal vent
Inferno Hell Crack Virgin MoundComponent,parameter Unit
T °C 328 301 217 299
pH(25°C) 3.5 3.5 – 4.4
Cl mmol/kg 624 550 258 176
SiO2 " 15.1 – – 13.5
CO2 " 50 90 179 285
H2S " 7.1 – – 18
He μmol/kg 2.45 – 8.44 11.2
Fe mmol/kg 1.065 0.868 0.0133 0.012
Mn " 1.150 1.136 0.287 0.142
Zn μmol/kg 111 134 2.6 2.2
Cu " 9.9 1.2 0.1 0.4
Pb " 0.302 – – 0.101
Br " 956 856 401 250
B " 590 – – 450
Br/Cl n × 10–3 1.532 1.556 1.554 1.420
He/CO2 n × 10–5 4.9 – 4.72 3.93
S280
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
gases in the hydrothermal solutions studied. This valuemay be different in other hydrothermal systems, but theprincipal mixing scheme is probably of general signifi-cance (Edmonds and Edmond, 1995).
Phase separation within hydrothermal systems affectsthe fate of ore elements. As can be seen in Table 24, theresidual brines are enriched in ore metals, whereas thecondensed vapor is strongly depleted in them butenriched in hydrogen sulfide. Hence, the mechanism ofprecipitant loss is in this case operative for chalcophileelements. This is in good agreement with the experi-ments by Bischoff and Rosenbauer (1987) simulatingseawater boiling in contact with a rock at P–T condi-tions approaching those of oceanic hydrothermal sys-tems. Table 25 shows the results of one of their experi-ments. It can be easily seen that a three-fold increase inbulk mineralization due to boiling is accompanied byan enrichment in Fe by a factor of two; Mn, by a factorof three; and Cu and Zn, by approximately two ordersof magnitude. This is indicative of extensive mobiliza-tion of chalcophile ore elements from the rock inresponse to H2S partitioning into the vapor phase.
Using the concentration of Cl– as an indicator of thedegree of phase separation, the whole set of availabledata on the compositions of oceanic hydrothermalsolutions (Table 6) was found to exhibit a distinctpositive correlation of this parameter with the con-centration of ore elements and a negative correlationwith H2S (Figs. 52a, 52b).32 The observed exponentialdependency indicates that, in addition to metal concen-tration due to evaporation, there is also an additionalinput into the solution.
32 The analyses of Cu in hydrothermal solutions do not show aclear relationship, probably because Cu is prone to losses duringcooling before and in the course of sampling.
Thus, the combined natural data set for all oceanichydrothermal systems indicates the leading role of themechanism of precipitant loss during subsurface boil-ing. Subsurface boiling serves in these systems as a fac-tor to increase the metal-bearing potential of hydrother-mal solutions.
The available data indicate the existence of phaseseparation (boiling) zones in the interiors of many oce-anic hydrothermal systems, including 9° N, 11° N,13° N, 17° S, 18° S, and 21° S on the EPR; the Cleft andEndeavor segments and Axial Seamount in the Juan deFuca Ridge; the Logatchev, TAG, Broken Spur, Rain-bow, Lucky Strike, and Menez Gwen fields on theMAR; the Kairei field in the Indian Ocean; the Lau,North Fiji, Manus, and Woodlark back-arc basins; andhydrothermal sites in the Okinawa Trough (Table 6).This raises the problem of assessing the factor of boil-ing for the geochemistry and metallogeny of oceanichydrothermal systems.
Treatment of this problem by the thermodynamicmodeling method leads to a number of specific issues,which are not typical of the classic problems of water–rock interaction considered in the previous chapters.Among them are
—a method for the calculation of thermodynamicequilibria in heterophase systems gas + aqueous solu-tion + solid phases (accounting for nonideality of thegas phase, thermodynamic properties of solution com-ponents for the P–T conditions of the existence of mul-ticomponent two-phase fluids, and programs for thesimulation of systems with several nonideal solutions);
—methods for the construction of geological andphysicochemical models of systems with boiling.
These issues are discussed in Sections 6.3 and 6.4,respectively.
6.3. Calculation of Thermodynamic Equilibria in Systems with the Gas Phase 6.3.1. Accounting for nonideal
mixing in the gas phase
Gas phase as an ideal mixture. The simplestapproach to the description of gas properties is obvi-ously an ideal gas approximation. There are no chemi-cal interactions between molecules in ideal gas (activitycoefficients are equal to 1), and the following expres-sions hold for a pure gas component
(48)
and for the component of a gas mixture
gT, P, i = + RTlnXi, (49)
where Xi is the molar fraction of the component in themixture.
The ideal gas approximation is adequate for the cal-culation of the properties of low-density gases, but it isnot always plausible for hydrothermal conditions. At
gT P i, ,0 gT 1 bar i ,, RT P ln+=
gT P i, ,0
Table 25.
Experimental data of Bischoff and Rosenbauer(1987) on the simulation of seawater boiling in contact witha rock
T
,
°
C 406 392
P
, bar 330 251
Phase stateof the system
Homo-geneous
Heterogeneous
Phase analyzed Water Water Vapor
Cl, g/kg 18.97 55.13 1.173
Fe, mg/kg 973 1791 37
Mn, mg/kg 97 280 2.4
Zn, mg/kg 0.1 4.9 0.04
Cu, mg/kg 0.01 1.9 0.01
H
2
S, mmol/kg 3.3 1.0 6.0
CO
2
, mmol/kg 5.8 1.9 8.6
H
2
, mmol/kg 0.6 0.1 1.1
Note: Rock/water = 1, the duration of the experiment is 20 days.
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S281
pressures of tens to hundreds of bars or higher, thefugacity of a pure gas component (
f
i
) is significantlydifferent from
P
, and this factor must be taken intoaccount in calculations. Then, Eq. (48) must bechanged to
= +
RT
ln
f
i
. (50)
The combined use of Eqs. (50) and (49) gives theapproximation of an ideal solution of real gases(Lewis–Randall rule). It is usually believed that thisapproximation can be applied within a wide range ofconditions, because the effects of nonideal mixing ofgas components are small.
33
Computer programs and models for ideal gas–min-
eral systems appeared even earlier than those for sys-tems with an aqueous phase (Volkov and Ruzaikin,1969). However, their application in geochemistryappeared to be limited and involved cosmochemicalproblems (e.g., Khodakovsky
et al.
, 1978) and calcula-tion of the state of volcanic gases (Symonds
et al.
,1992; Symonds and Reed, 1993). A mixed model issometimes used: the deviation of fugacity from pres-sure is taken into account for major components (H
2
Oand CO
2
) but ignored for minor components (Bowers,1991).
Gas phase as a real solution.
A provision for non-ideal mixing must be made for a more accurate descrip-tion of the properties of gas mixtures. Three mainapproaches are used to this end in the modern thermo-dynamics of gases. (1) The properties of gases and theirmixtures (in a general case, equations of state) areexpressed as polynomials of independent parameters ofthe system (
T
,
P
or
V
, and
X
i
). (2) Equations of state areconstructed using physical model approximations con-necting the
T
,
P
, and
V
of gases and their mixtures (thevan der Waals equation and its modifications). (3) Themolecular dynamics method employs calculationsbased on the model potentials of intermolecular inter-action.
The molecular dynamics method probably has supe-rior prognostic properties, but its application isrestricted mainly to the calculation of gas propertiesunder high parameters of state because of computa-tional difficulties. To our knowledge, there have beenno attempts to use this method for the thermodynamicmodeling of hydrothermal ore formation.
The method of polynomial functions was used forthe thermodynamic modeling of geochemical pro-cesses in the CHILLER program (Spycher and Reed,
33
Fernandez-Prini and Crovetto (1985) showed that Henry’s con-stants for water–nonpolar gas systems calculated in the Lewis–Randall approximation deviate from experimental values morestrongly than those obtained in the ideal gas approximation. Thisphenomenon is related to the opposite signs of corrections for thefugacity of real gases and for real gas mixing. Thus, the resultingerrors are cancelled in the ideal gas approximation. Similar rela-tionships of errors were reported for some other instances bySpycher and Reed (1988).
gT P i, ,0 gT 1 bar i , ,
0
1988, 1989). This method has the highest accuracy forthe interpolation of experimental data sets. However,more important for the multicomponent systems ofgeological interest are extrapolation of properties bothwith respect to intensive parameters (
T
and
P
) and thecomponent composition of the system. Holloway(1986) argued that the empirical parameters of polyno-mials obtained by the least squares method are usuallystrongly correlated. Therefore, the quality of dataapproximation decreases near the boundaries of theexperimentally studied domain and extrapolationbeyond it requires considerable care. As to the numberof system components, the CHILLER program consid-ers a gas phase consisting of five components, and non-ideal mixing effects are taken into account only forH
2
O, CO
2
, and CH
4
. Even such a simple gas mixturerequired the use of two-dimensional and three-dimen-sional arrays of cross-coefficients, each of which is apolynomial in
T
and
P
. This approach evidentlyrequires a tremendous amount of experimental infor-
–3.0
–2.5
–2.0
–1.5
–1.0
log(H
2
S)
0 0.5–6
–4
1.0 1.5
–3
–2
–1
Cl
–5
log(Me)
Fe
Zn
(a)
(b)
Fig. 52.
Correlation of the concentrations of ore elementswith the chlorinity of oceanic hydrothermal solutions.(a) Hydrogen sulfide and (b) iron and zinc. Concentrationsare given in mol/kg.
S282
GEOCHEMISTRY INTERNATIONAL
Vol. 42
Suppl. 2
2004
GRICHUK
mation, and its complexity increases as a power func-tion of the number of components in the system. Thus,the price for the accuracy of the method is difficultapplication for complex systems.
The method of equations of state is based on the vander Waals equation, which extends the Clapeyron–Mendeleev equation of ideal gases (
PV
=
RT
, where
R
is the gas constant) by accounting for the nonzerovolume of gas molecules:
(51)
This equation was modified in 1948 by O. Redlichand J. Kwong:
(52)
The coefficients
a
and
b
of this equation can bereadily determined on the basis of the correspondingstate principle from the critical point parameters of puresubstances.
The very popular Redlich–Kwong equation hasrepeatedly been modified by many authors, both withrespect to the form of the second term and the methodof
a
and
b
derivation. For instance, in the Soave equa-
tion,
a
/ is changed to
a
(
T
) =
a
cr
α
(
T
),
α
(
T
) = 1 +
m
,
m
= 0.480 + 1.574
ω
– 0.176
ω
2
,
where
ω
is the Pitzer acentric factor, which is specificto each gas. A modified Redlich–Kwong equation was
PRT
V b–------------
a
V2------.–=
PRT
V b–------------
a
T V b+( )V------------------------------.–=
T
1 TTcr------–⎝ ⎠
⎛ ⎞
incorporated into a version of the EQ3/6 programdesigned for the calculation of heterogeneous systemswith the gas phase (Bowers, 1991).
Peng and Robinson (1976) developed a modifiedequation of state, which showed important propertiesfor the description of gas mixtures,
(53)
where
b
= 0.086640350 ,
and
m
= 0.37464 + 1.54226
ω
– 0.26992
ω
2
.
The Peng–Robinson equation describes the proper-ties of a pure gas using only three individual properties:
T
and
P
of the critical point of the gas and the Pitzeracentric factor
ω
. These properties have been deter-mined for a wide spectrum of substances, includingmany compounds of geochemical interest. The valuesof critical parameters and
ω
for some gases are listed inTable 26.
For the calculation of gas mixtures, Peng and Rob-inson made the following assumptions: (1) the proper-ties of a gas mixture are identical to those of a hypothet-ical gas having the same critical
T
and
P
as the mixture;(2) the and parameters of a mixture are functionsof mixture composition and the parameters of themixed components,
b
i
and
a
(
T
)
i
:
(54)
(55)
and
(56)
where
δ
ij
is the cross-coefficient,
δ
ij
=
δ
ji
and
δ
ii
= 0.Thus, only one two-dimensional array of coeffi-
cients is used for gas mixtures (
δ
ij
). Peng and Robinson(1976) argued that the
δ
ij
parameters can be regarded astemperature-invariant. Therefore, the Peng–Robinsonequation must be very robust with respect to tempera-ture extrapolation. In order to improve the accuracy ofcalculations, subsequent studies approximated thecross-coefficients by polynomial functions of tempera-ture for some systems. The values of
δ
ij
have beendetermined for many pairs of geochemically importantcompounds (Table 27).
PRT
V b–------------
a T( )V V b+( ) b V b–( )+-------------------------------------------------,–=
RTcr
Pcr-----------
a T( ) 0.42748002327R2Tcr
2
Pcr------------- 1 m 1 T
Tcr------–⎝ ⎠
⎛ ⎞+2
,=
a b
b xibi,i
∑=
a xix jaij,j
∑i
∑=
aij aia j 1 δij–( ),=
Table 26. Critical parameters and acentric factors ofgeochemically important gas compounds
Com-pound Tcr, K Pcr, bar Vcr,
cm3/mol Zcr ω
H2O 647.3 220.5 56.0 0.229 0.344
CO2 304.2 73.8 94.0 0.274 0.225
CO 132.9 35.0 93.1 0.295 0.049
CH4 190.6 46.0 99.0 0.288 0.008
H2S 373.2 89.4 98.5 0.284 0.100
SO2 430.8 77.8 122.0 0.268 0.251
HCl 324.6 83.1 81.0 0.249 0.120
HF 461.0 64.8 69.0 0.120 0.372
H2 33.2 13.0 65.0 0.305 –0.220
N2 126.2 33.9 89.5 0.290 0.040
NH3 405.6 112.8 72.5 0.242 0.250
O2 154.6 50.5 73.4 0.288 0.021
Ar 150.8 48.1 74.9 0.291 –0.004
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S283
Equation (53) cannot be directly used in programscalculating equilibria under isobaric–isothermal condi-tions, because its arguments are T and V. Therefore, thePeng–Robinson equation, as well as other equations ofthe Redlich–Kwong type, must be recast by changingvariables into a cubic polynomial in compressibility, Z:
(57)
where
and
The Peng–Robinson equation allows calculation ofthe partial parameters of mixed components, amongwhich the most important are the fugacity coefficientsof individual components in the mixture, ϕi = fi/xiPtot.Peng and Robinson (1976) transformed Eqs. (54)–(57)to obtain
(58)
where Z, A, B, , , and aij are given by Eqs. (57) and(54)–(56), respectively. Thus, the fugacity coefficientof an individual component is defined by the composi-tion and compressibility of the gas mixture (Z), theadditive functions ( and ) of the properties of purecomponents of the mixture (Tcr, i, Pcr, i, and ωi), and thecross-coefficients (δij).
Owing to its simplicity and the low number ofempirical parameters, the Peng–Robinson equation isvery attractive for the thermodynamic modeling ofmulticomponent systems. Mironenko and Kosorukov(1990) implemented the Peng–Robinson equation intheir computer program for geochemical modeling.
The downside of the simplicity of the equation ofstate is its lower accuracy compared with polynomialequations. Therefore, the accuracy of the description ofthe properties of gas phase must be specially evaluated.
Estimation of errors related to the use of the Peng–Robinson equation. Similar to other model equations,the Peng–Robinson equation does not account for all
Z3 1 B–( )Z2– A 3B2– 2B–( )Z+
– AB B2– B3–( ) 0,=
A aP
R2T2------------,=
B bP
RT-------,=
ZPVRT--------.=
ϕilnbi
b---- Z 1–( ) Z B–( )ln–
A
B2 2--------------–=
×
2 xiaij
j
∑a
---------------------bi
b----–
⎝ ⎠⎜ ⎟⎜ ⎟⎜ ⎟⎛ ⎞
Z 2 1+( )B+
Z 2 1–( )B–-----------------------------------⎝ ⎠
⎛ ⎞ ,ln
b a
a b
physical properties of molecules. According to the cor-responding states principle, it includes temperature andpressure normalized to the critical parameters (T/Tcrand P/Pcr, respectively) and the acentric parameter ω,which accounts for the deviation of the geometry of gasmolecules from spherical symmetry. However, there isno explicit provision for possible dipole–dipole interac-tions. Correspondingly, the Peng–Robinson equation(as well as other Redlich–Kwong type equations) ade-quately describes nonpolar gases but is not so good forpolar gases. The geochemically most important polargas is water vapor.
The properties of water vapor are well known. Thereare rather precise models allowing calculation of itsproperties, among which is the Haar–Gallagher–Kellmodel used in the modified HKF model and in theUNITHERM data bank. This provides an opportunityto check the accuracy of the Peng–Robinson equationfor gaseous H2O. Table 28 shows H2O fugacity calcu-lated along the boiling curve of pure water by means ofthe Haar–Gallagher–Kell model and the Peng–Robin-son equation. According to these results, the error inwater vapor fugacity calculated by the Peng–Robinsonequation goes up to 2% relative and is highest at 250°C.The energy equivalent to this error (gT(H2O)gas) is 30–
Table 27. Cross-coefficients, δij, for geochemically impor-tant gas compounds
Mixturecomponents δij
Refer-ence
T inter-val, °C
H2O–CO2 –0.5572 + 0.001879T – 1.274 × 10–6T2
(a) >100
0.49852–0.0008T (b)
0.057 (e)
H2O–H2S –0.3897 + 1.565 × 10–3T – 1.142 × 10–6T2
(a) 0–220
–0.4860 + 2.092 × 10–3T –1.87 × 10–6T2
(d)
0.157 (e)
H2O–H2 0.30 (c) <363
H2O–N2 0.53 (c) <350
0.45 (f)
H2O–NH3 –0.260 (a)
H2O–He 0.80 (f)
CO2–H2S 0.096 (a)
0.093 (e)
CO2–CH4 0.65 (b)
CO2–NH3 0.096 (a)
H2S–NH3 0.324 (a)
Note: (a) Roberts and Tremaine (1985), (b) Mironenko andKosorukov (1990), (c) Alvarez et al. (1988), (d) Carroll andMather (1989), (e) Shvedenkov and Savinov (1994), and(f) Fernandes-Prini and Crovetto (1989).
S284
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
80 J/mol, which seems to be rather small comparedwith other error sources in multicomponent models.However, it should be kept in mind that models withboiling consider gas to be in equilibrium with liquidwater. The uncertainty obtained in Table 28 corre-sponds to a displacement of the boiling curve byapproximately one bar. This may be rather embarrass-ing in the calculations of isobaric–isothermal equilib-ria. An obvious way to circumvent this difficulty is theuse of a combined method, involving the more accuratemodel of Haar–Gallagher–Kell for the fugacity of purewater vapor and the Peng–Robinson equation for othercomponents and nonideal mixing effects (Japas andLevelt-Sengers, 1989).
A very important circumstance for the assessmentof the plausibility of the Peng–Robinson equation isthat H2O shows a very strong dipole interaction. Thisinteraction is characterized by the dipole moment of themolecule (μ). Table 29 presents dipole moments for themolecules of geochemically important gases. Thedipole moment of the H2O molecule is the highest. Cor-respondingly, the effects related to dipole interaction inwater vapor are much higher than those for other sub-stances. Therefore, the errors of the Peng–Robinsonequation for H2O are the maximum possible values innatural systems.
The substitution of a gas mixture by a hypotheticalgas with the same critical parameters that is made in thePeng–Robinson equation (as well as in other Redlich–Kwong type equations) is theoretically not strictly cor-rect (this issue was discussed by Namiot, 1991). Theresult of this substitution is that the equation betterdescribes the homogeneous region than the liquid–gasphase boundaries.
The accuracy of various equations of state wasassessed by Seitz et al. (1994) by the example of theCO2–CH4–N2 system (200°C and 1000 bar), for whichcomprehensive experimental data are available. Theyshowed that the accuracies of the Redlich–Kwong andPeng–Robinson equations are inferior to those of poly-nomial equations. One important conclusion by theseauthors is that the main contribution to the errors of theRedlich–Kwong and Peng–Robinson equations isrelated to the uncertainties in the properties of pureboundary components, which were discussed above inthe example of H2O. The mixing volume effect in theternary system considered was very small (<3% of thevolume of mixture), and the Peng–Robinson equationreproduced it with an error of ±45% relative, which iscomparable to the error for pure components. ThePeng–Robinson equation appeared to be more accuratethan the Redlich–Kwong equation. Japas and Levelt-Sengers (1989) used the precise Haar–Gallagher–Kellequation for water and calculated the compositions ofequilibrium gases in binary systems with water bymeans of the Peng–Robinson equation. Their results forthe molar fractions of gas components were accurate towithin about 1%.
Peng and Robinson (1976) assumed that the δij val-ues are temperature independent. A detailed study ofthe H2O–H2S system (Roberts and Tremaine, 1985;Carroll and Mather, 1989) showed that the properties ofthis system can be better described using temperature-dependent values. The polynomials obtainedby these authors are given in Table 27. According to theresults of Carroll and Mather, δij varies from –0.06 at0°C to +0.08 at 220°C. These variations in δij for theH2O–H2S system result in variations in the activity
δH2O–H2S
Table 28. Comparison of the calculated H2O fugacities along the line of pure water boiling with the Haar–Gallaher–Kellmodel and the Peng–Robinson equation
Temperature,°C
Pressure,bar
Fugacity, bar Deviation
Haar–Gallaher–Kell method*
Peng–Robinsonequation Δf, % , J/mol ΔP, bar
150 4.758 4.604 4.54 –1.39 49.3 –0.07
200 15.537 14.321 14.48 1.11 –43.4 0.17
250 39.728 34.178 34.84 1.94 –83.4 0.77
300 85.805 67.442 68.33 1.32 –62.3 1.13
350 165.125 115.632 115.0 –0.55 28.4 –0.90
* Calculated using the UNITHERM data bank.
ΔGH2O
Table 29. Dipole moments of geochemically important gasmolecules (Holloway, 1986)
Substance Dipole moment, Debye
H2O 1.86
CO 0.11
HCl 1.07
H2S 0.9
SO2 1.59
NH3 1.47
Note: Nonpolar gases (CO2, CH4, H2, N2, O2, etc.) show a dipolemoment of zero owing to the symmetry of their molecules.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S285
coefficient of H2S within ±30% in the water-richregion, which is no higher than other errors in the ther-modynamic models of multicomponent systems.
Several attempts have been made to improve thePeng–Robinson equation, for instance, the PRSVmodel by Stryjec and Vera (1986). However, all theseattempts introduced new empirical parameters. Theimprovement of accuracy always reduced the possibil-ity of practical application of the method.
Thus, the Peng–Robinson equation can be used inthe thermodynamic models of multicomponent systemsas a not very accurate but universal tool for the calcula-tion of gas equilibria.
6.3.2. Specific features of the thermodynamic description of phase properties in the model
of boiling hydrothermal systems
Phase diagram of the H2O–NaCl system and itsapplications to seawater boiling. The H2O–NaCl sys-tem has been extensively studied in the range of hydro-thermal conditions. Early studies were surveyed byHolland and Malinin (1979). In the 1980s, detailedexperimental investigations of the H2O–NaCl systemwere performed by Bischoff et al. (1986), Bischoff andRosenbauer (1988), and other authors. The results ofthis work were summarized by Bischoff and Pitzer(1989). Bischoff and colleagues paid considerableattention in their studies to the position of the criticalline in this system, because the critical parameters ofsolutions play an important role in the correspondingstates theory and their experimental determination is asignificant challenge. The H2O–NaCl system has beenstudied to such an extent that an equation of state wasproposed for it (Tanger and Pitzer, 1989), which isunique for water–salt systems. The projection of theboiling surface on the P–T plane is constructed inFig. 53a using the data of Bischoff and Pitzer (1989).Its isothermal sections, boiling curves for temperaturesof 350, 370, 380, and 400°C, are shown in Fig. 53b.These diagrams exhibit the main property of water–saltsolutions: the addition of electrolyte decreases the pres-sure (increases the temperature) of boiling. For theregion of interest, the depression of boiling pressure isan almost linear function of solution salinity.
Since sodium chloride is the predominant salt com-ponent of seawater, the H2O–NaCl system has tradi-tionally been regarded as a reference system for thephysicochemical analysis of geochemical processesinvolving seawater. The compositions of hydrother-mal solutions only approximately correspond to theH2O–NaCl binary and contain admixtures of K and Cachlorides. However, these admixtures probably have aminor influence on the position of the boiling surface(Holland and Malinin, 1979).
The presence of dissolved gases will increase boil-ing pressure (decrease boiling temperature). The funda-mental thermodynamic reason for this effect is the fol-
lowing: dissolved gases are extensively partitioned intothe gas phase, which reduces the molar fraction of H2Oin the vapor and, correspondingly, the chemical poten-tial of gaseous water. This stabilizes the gas phase.
Figure 53a shows the points of temperature mea-surements in oceanic hydrothermal systems. Accordingto these data, almost all solutions (except for two cases)occurred in the single-phase field at the moment ofejection. Extremely high for oceanic hydrothermal sys-tems temperatures were recorded in smokers at thehydrothermal system of 9° N EPR in 1991 immediatelyafter a submarine eruption, 388–403°C at 253–258 bar(Von Damm et al., 1995). The samples of these solu-tions showed a low salinity (up to 10–15% of the nor-mal seawater value, Table 6) and high gas contents,which indicates their relation to phase separation (boil-ing) phenomena (Von Damm, 2000). Even a higher
400
150
300 32050
100
(a)
340 360 380 420Temperature, °C
200
250
300
350
400
Pressure, bar
40
300
302010 500
260
220
180
140
100
Pressure, bar
NaCl, wt %
350°C
370°C
380°C
400°C
Critical
Critical
Critical
300
NaCl, g/kg
(b)
0
35
100
line
points
line
Fig. 53. Phase diagram of the H2O–NaCl system with theboiling surface after Bischoff and Pitzer (1989). (a) Con-centration isolines in the T–P coordinates and (b) isothermsin the P–X coordinates. Crosses show the T–P parameters ofoceanic hydrothermal solutions, rhombs are the data forvents at 9°17′–9°46′ N on the EPR, April 1991 (Von Damm,2000), and triangles are the data for the Brendon vent(21°34′ S EPR, November, 1998) (Von Damm et al., 2003).
S286
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
temperature of 405°C (287 bar) was documented in1998 in the Rapa Nui vent field (21°34′ S EPR) (VonDamm et al., 2003). These parameters almost exactlymatch the critical point of seawater (407°C and298 bar). Fluid from three chimneys of the hydrother-mal edifice was less saline than seawater, whereas twoother chimneys discharged a more saline fluid, which isindicative of phase separation within the edifice.
Thermodynamic description of the properties of theliquid phase in a two-phase field. As was mentioned inSection 3.3, the current methods of derivation of thethermodynamic properties of dissolved substances arebased on the modified Helgeson–Kirkham–Flowersmodel, which relies on the thermodynamic propertiesof pure water (density, dielectric constant, and theirderivatives with respect to temperature and pressure).The simulation of boiling processes in water–salt sys-tems encounters the problem of the existence of a liquidphase in the presence of salts under the P–T conditionswhere pure water exists in the vapor state. Standardprograms for the calculation of thermodynamic proper-ties (including SUPCRT 92 by Johnson et al., 1992) arenot appropriate for such conditions.
A pragmatic way to overcome this difficulty is to usethe extrapolated thermodynamic properties of dis-solved substances. The extrapolation of properties inpressure is evidently advantageous, because the pres-sure dependency of gT is less steep than the temperaturedependency. A check by the properties of substancesused in our model of an oceanic hydrothermal system(Table 18) showed that the data obtained immediatelynear the boiling curve of pure water (from Pboiling toPboiling + 20 bar) are adequately described by a quadraticfunction of pressure. The error of such an approxima-tion is no higher than ±100 J/mol. This allows extrapo-lation of gT of dissolved substances within a limitedrange below Pboiling. It should be taken into account thatthe error of extrapolation definitely increases awayfrom the boiling curve. Note, however, that the ordinaryuncertainties of thermodynamic data, which are
involved in the extrapolated values as systematic errors,are significantly higher. In general, the variations in gTwith pressure are rather small: at 350°C the gT value ofdissolved species changes by a few kilojoules as pres-sure decreases from 166 to 140 bar. The maximumeffects were obtained for the OH– ion and relatedhydroxo complexes (about –3.5 kJ/mol per onehydroxyl ion). Unfortunately, there are still no data thatwould allow the assessment of the accuracy of extrapo-lated values.
Thermodynamic properties of water vapor in thestability field of liquid water. As was mentioned above,the description of vapor phase properties in boiling sys-tems is based on the thermodynamic properties of gas-eous water calculated by an accurate model (forinstance, by the Haar–Gallagher–Kell equation). Stan-dard programs for the calculation of the thermody-namic properties of substances provide these parame-ters only for the stability field of the pure water vaporphase. If boiling is calculated in multicomponent sys-tems at pressures above the boiling curve of the single-component system (which is very probable in systemswith high gas content), the properties of H2Ogas have tobe extrapolated into the metastable field. It was shownthat extrapolation of gT in pressure using a third degreepolynomial yields satisfactory results for the subcriticalregion.
Quality control of thermodynamic data for boilingsystems. Currently, such a control can be performedonly through the simulation of boiling in the H2O–NaClsystem. The simulation was carried out for the 350°Cisotherm using the experimental data of Bischoff andPitzer (1989) (Fig. 54). The diagram indicates that thecalculated values are in general consistent with theexperimental points. The model and experimental pres-sures of the beginning of boiling are identical,165.2 bar. The calculated curve for salt solutionsimplies somewhat lower degrees of boiling comparedwith the experiments. The discrepancy increases withincreasing salinity, and the maximum value of 5% NaClwas obtained at 150–140 bar.
Thus, the simulation of boiling is in general consis-tent with experimental data, and the model somewhatunderestimates the fraction of water vapor in the sys-tem. The discrepancy may be as high as 5 bar in pres-sure or 5% in electrolyte concentration.
6.3.3. Software
The calculation of thermodynamic equilibria inboiling systems requires computer programs designedfor systems with several phases of variable composi-tion. To the author’s knowledge, such programs areGIBBS (for large computers), SELECTOR, and HCh(for personal computers).
Boiling simulation was performed by the authorusing the PENG program, which is a modification ofthe GBFLOW software. PENG allows calculation of
170
160
150
140
130
120
110
10010 20 30 40 500
NaCl concentration, wt %
P, bar
Calculation
Experiment
Fig. 54. Comparison of the thermodynamic calculation ofboiling in the H2O–NaCl system at 350°C with the experi-mental data of Bischoff and Pitzer (1989).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S287
equilibria in systems with two phases of variable com-position, liquid and gas. According to Shvarov’snomenclature (Methods of…, 1988), PENG is a pro-gram for equilibrium calculations of the second class,i.e., a program for equilibrium calculation in systemswith a partially constrained phase composition,because the liquid water phase must always be presentin equilibrium. The PENG program utilized a modifiedalgorithm of Gibbs free energy minimization, whichincluded improved procedures for the calculation ofactivity coefficients (they are computed for the gasphase by a special subroutine based on the Peng–Rob-inson equation) and chemical potentials of components(transition from the molar amount of a component inthe system to molar fractions requires separate summa-tion for each phase of variable composition).
6.4. Thermodynamic Model of Boiling in Oceanic Hydrothermal Systems
The main obstacle to the simulation of boiling inoceanic hydrothermal systems is related to the almostcomplete lack of information on the situation in boilingzones. Neither the T–P conditions, the character ofphase separation, nor the residence time of fluid in boil-ing zones are reliably known. Although the conse-quences of boiling and phase separation were tracked inthe compositions of hydrothermal solutions, the realmechanism remains obscure. There were attempts tocreate conceptual scenarios of phase separation (two-layer convection cell by Bischoff and Rosenbauer, 1989and separation in an oblique fault by Fox, 1990), but aquantitative dynamic model of the process has not yetbeen developed. Therefore, our study of the problemsof the thermodynamic modeling of boiling in hydro-thermal systems was restricted to the determination ofsuch characteristic that are independent of the particu-lar structure of boiling zones, i.e., invariants of the boil-ing process.
Some observed properties of oceanic hydrothermalsolutions, an analogy with terrestrial hydrothermal sys-tems, and general theoretical considerations constrainthe diversity of possible scenarios of the boiling pro-cess. Among such constraints are the following.
(a) Two-phase fluids have a low bulk density. There-fore, the appearance of a boiling zone in a convectioncell sharply increases the expelling force of convectionand, correspondingly, the velocity of fluid movement.The boiling zone can be preserved only near the heater(magma chamber).
(b) The process can be regarded as adiabatic in theupwelling limb because of the rapid fluid movement(Bischoff and Pitzer, 1985).
(c) Since the vapor component of the two-phasefluid has a low viscosity, the boiling zone can propagatefar upward along the channel of the convective system.However, as was mentioned above, the discharge oftwo-phase fluids in deep oceanic hydrothermal systems
has never been observed. The reason for this peculiarityof oceanic hydrothermal systems is probably related tothe rapid condensation of the vapor phase due to sub-surface mixing (vapor bubbles are collapsed by theaddition of only a small percent of cold seawater).
(d) The observed variations in the salinity of oceanichydrothermal solutions suggest a possible phase sepa-ration (at least partial) of vapor and residual liquid.
(e) According to the calculations of Bowers et al.(1988), hydrothermal solutions with salinities higherthan that of seawater are in equilibrium with the majorminerals of metasomatic rocks. This suggests that theresidual brines of many hydrothermal systems havebeen in contact with rocks for a fairly long time.
Taking into accounts these constraints, we consid-ered the following boiling scenarios:
(1) isothermal boiling in contact with a rock withoutvapor phase elimination, which is an analog of boilingin the high-temperature focus of the system;
(2) phase separation in contact with a rock;(3) adiabatic cooling of a two-phase fluid, which is
an analog of upward movement along an extended con-duit; and
(4) condensation of a two-phase fluid during mixingwith seawater.
The principal schemes of these scenarios are shownin Fig. 55.
6.4.1. Isothermal boiling of solutionin contact with a rock: a model for the focus
of a hydrothermal system
In this scenario (Fig. 55a), the equilibrium composi-tions of the system tholeiitic basalt–hydrothermal solu-tion–vapor were calculated as functions of confiningpressure. The calculations were performed for a tem-perature of 350°C, which is sufficiently high to simu-late the focus of a hydrothermal system but still allowsfor the acquisition of reliable thermodynamic data fordissolved species. The initial solution composition cor-responded to that calculated by the simulation of thedownwelling limb of a convection cell (Chapter 4) forthe following conditions: T = 350°C, P = 500 bar,ΣR/W = 0.496, and wave no. 1 (Table 30). Pressure inthe boiling model varied between 190 and 130 bar(Pboiling = 165.2 bar for pure water at 350°C), whichallowed us to explore a wide range of the degree of boil-ing off. The amount of rock in contact with the boilingsolution was taken as 0.1 kg per one kilogram of solu-tion; the same value was accepted for this temperaturein the model of the downwelling limb. The obtained setof equilibrium states of the system can be interpreted asa stepwise decrease in pressure in a closed system(autoclave with a pressure drop).
The results of simulation are shown in Fig. 56. Boil-ing in the model system begins when pressure falls to
S288
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
165 bar. This pressure is close to the boiling point ofpure water (165.2 bar), which results from the superpo-sition of two opposite effects, a salinity-related increasein boiling pressure (by approximately 5 bar) and adecrease in pressure due to the presence of gases. Asubsequent pressure decrease results in a rapid increasein the fraction of the vapor phase in the system(Fig. 56a), which reaches 50% of the mass of fluid at155 bar and 88% at 130 bar. Boiling is accompanied byan increase in the salinity of the liquid phase: from aninitial mineralization of 33 g/kg, it becomes twice ashigh at 155 bar and reaches 260 g/kg at 130 bar(Fig. 56b).
For the sake of comparison, Fig. 56b also presentsthe calculated mineralization of solution during boilingin the NaCl–H2O system (at an initial NaCl concentra-tion of 0.587 mol/kg H2O, corresponding to the miner-alization of the model hydrothermal solution) and theexperimental data for boiling in the NaCl–H2O systemobtained by Bischoff and Pitzer (1989). The calcula-tions for the binary system place the beginning of boil-ing (established by an inflection in the diagram) at apressure of 159 bar, which is in good agreement withthe experiments (161 bar for a salinity of 35 g/kg). Boil-ing in the multicomponent model system begins earlier,at a total pressure of 165 bar, because the partitioningof dissolved gases into the vapor phase reduces the frac-
Fig. 55. Principal schemes of models for boiling scenarios. (a) Isothermal boiling, (b) phase separation, (c) adiabatic boiling, and(d) rapid discharge with vapor condensation.
(a) Heterogeneous
Homogeneous fluid
fluid
(b) Vapor
Solution
Vapor Vapor
(c)
Seawater
(d)
Heterogeneousfluid
S = const
P1 P2> > P3
T = const
T1, P1
P2 < P1
T2 = T1
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S289
0
–16
170
log(X)
Pre
ssure
, bar
(g)
160
150
140
130
180
Wat
er
Hom
ogen
eous
–7
–8
logC [mol]
(f)
20 0
Fraction, %
(e)
0.0
05 0
mol/kg
(d)
10
–2
10
–3
170
mol/kg
Pre
ssure
, bar
(c)
160
150
140
130
180
50 0
g/kg
(b)
20 0
%
(a)
40
60
80
100
Beg
innin
gV
apor
of
boil
ing
Model
cal
cula
tion
100
150
200
250
300
Exper
imen
ts i
n N
aCl–
H2O
syst
em
Cal
cula
tion f
or
NaC
l–H
2O
syst
em
Beg
innin
gof
boil
ing
Cal
cula
ted b
egin
nin
gof
boil
ing i
n N
aCl–
H2O
syst
em
10
–1
10
0
10
1
Het
erogen
eous
Dis
appea
rance
regio
nre
gio
n
K Na
Ca
Fe
Si
Cl
0.0
10
0.0
15
Hom
ogen
eous
Het
erogen
eous
regio
n
regio
n
H2
CH
4S
(II)
40
60
80
100
120
Beg
innin
gof
boil
ing
Soli
d p
has
e
Solu
tion
Vap
or
–6
–5
–4
–3
–2
–1
of
Sph
Hom
ogen
eous
Het
erogen
eous
regio
n
regio
n
Fe
Cu
Zn
Pb
S(I
I)
–2
–4
–6
–8
–10
–12
–14
Hom
ogen
eous
regio
n
CO
2C
OC
H4
H2S
SO
2H
Cl
H2
Fig
. 56.
Com
posi
tion
of h
ydro
ther
mal
sol
utio
ns d
urin
g bo
iling
in c
onta
ct w
ith a
roc
k ca
lcul
ated
at T
= 3
50°C
and
R/W
= 0
.1. (
a) T
he f
ract
ion
of li
quid
wat
er, w
t %; (
b) m
iner
aliz
atio
n of
resi
dual
solu
tion
afte
r boi
ling
off,
g/kg
; (c)
con
cent
ratio
ns o
f maj
or c
ompo
nent
s in
solu
tion,
mol
/kg;
(d) c
once
ntra
tions
of v
olat
ile c
ompo
nent
s in
solu
tion,
mol
/kg;
(e) p
artit
ion
of S
II b
etw
een
the
phas
es o
f the
sys
tem
; (f)
con
cent
ratio
ns o
f ore
ele
men
ts in
sol
utio
n; a
nd (g
) mol
e fr
actio
ns o
f gas
es in
the
vapo
r pha
se. T
he d
ashe
d lin
e in
Fig
. 56b
sho
ws
expe
rim
enta
l dat
a fo
r the
350
°Cbo
iling
isot
herm
in th
e N
aCl–
H2O
sys
tem
(Bis
chof
f and
Pitz
er, 1
989)
. The
arr
ows
in F
ig. 5
6f s
how
the
initi
al c
ompo
sitio
ns in
the
mod
el o
f adi
abat
ic c
oolin
g.
S290
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
tion of water in it to 98.6% in the boiling point of themodel system. As pressure decreases to 160–155 bar, themodel multicomponent system becomes similar to thebinary system, because with increasing mass of thevapor phase, its composition approaches 100 mol %H2O. When the pressure of the solution decreasesbelow 150 bar, the simulated multicomponent solutiondeviates significantly from the binary system towardhigher mineralization, perhaps owing to the accumula-tion of Ca and K chlorides in the solution (their contentrises above 20% of the sum of salts). In general, theconcentrations of solutions in the model of hydrother-mal fluid boiling fall between the calculated and exper-imental curves of the binary NaCl–H2O system. Takinginto account the complex composition of the modelsolution and the model assumptions, such an agreementcan be regarded as satisfactory.
This calculation revealed a characteristic feature ofmulticomponent systems with volatile components:boiling begins in them at higher pressure than in thebinary system (by 5 bar), but the mass fraction of vaporincreases rather slowly as pressure decreases within165–160 bar (Fig. 56a). This is the net result of twoopposite effects: the presence of admixture gases invapor promotes boiling, whereas the enhancement of
boiling dilutes the gas phase with water vapor and rap-idly diminishes the role of the admixtures (Fig. 56g).
The concentrations of major components (Cl, Na,Ca, and K) in the solution increase proportionally dur-ing boiling. The concentration of Fe grows faster: whileCl concentration increases by a factor of 10, Fe showsa 30-fold increase and reaches 47 mmol/kg at 130 bar.This effect is a consequence of complexation: the main
Fe species, Fe and FeOH , depend on the squareof the chloride ion concentration. In contrast to themajor cations, the concentration of Si in the solutiondecreases. The reason for such a behavior is that theactivity of silica in solution is connected with Na activ-ity through the activity product of albite, which is oneof the major mineral phases in the system. Since theconcentration of Na increases during solution evapora-tion, the equilibrium concentration of silica declines.
According to the results of simulation, the pHT valueof the solution is practically constant during boiling andonly slightly increases (from 6.46 to 6.87) with consid-erable evaporation. In contrast, the redox state of thesolution changes significantly during boiling: log( )declines from –1.85 at 166 bar (before boiling) to –2.73at 50% evaporation (155 bar). This effect is related tothe partitioning of volatile components between the liq-uid and vapor phases. Figure 56d shows the concentra-tions of reduced volatile components (H2, CH4, andH2S) in the liquid phase. It can be seen that these com-ponents are already removed from the aqueous solutionduring incipient boiling, when the mass of the vaporphase is still low. As a result, the concentrations ofhydrogen and hydrogen sulfide decrease in the liquidphase by an order of magnitude, and that of methane,by three orders of magnitude. The distribution ofhydrogen sulfide between phases is shown in Fig. 56e,which suggests that half of the initial H2S inventory hasalready migrated from the liquid phase into vapor whenthe degree of solution evaporation is 10%. The deple-tion of H2S in the aqueous solution is much more pro-nounced in this diagram than in the previous one: Fig.56e shows the partition of SII between the phases of thesystem, whereas Fig. 56d shows its concentration. Adecrease in the mass of the liquid phase intensifies theprocess of redistribution.
Figure 56g shows the composition of the equilib-rium vapor phase. Within the whole range of condi-tions, it is strongly dominated by water vapor. Minorcomponents are (in descending order) H2, H2S, CO2,and, at the initial stage of boiling, CH4. The molar frac-tion of these substances is 0.n% at low degrees of evap-oration and declines with decreasing total pressure to0.0n%. An exception is methane, whose concentrationin the vapor phase decreases to 10–5% owing to oxida-tion. The acid gases considered in the problem, HCl andSO2, show very low equilibrium concentrations andhave a negligible influence on the character of the pro-cess.
Cl20 Cl2
–
aH2
Table 30. Composition of the model hydrothermal solutionthat was used as the initial state in models with boiling; thiscomposition was obtained by the simulation of the down-welling limb of the convective system at T = 350°C, P =500 bar, and ΣR/W = 0.496 for the first wave
Component Amount, moles
H 0.117304
O 0.060189
K 0.028856
Na 0.467914
Ca 0.024079
Mg 1.04 × 10–7
Fe 0.001177
Al 8.12 × 10–6
Si 0.013993
C 0.006694
S 0.008247
Cl 0.5459
Cu 1.26 × 10–6
Zn 2.47 × 10–5
Pb 3.27 × 10–7
H2O 54.59448
pHT 5.73
Equilibrium mineralassociation
Gn, Sph, Ccp, Ab, Ep60, Chl75, Dph, Act80
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S291
The most interesting question in the numerical sim-ulation is the behavior of ore elements. Figure 56f dis-plays the concentrations of ore elements in the liquidphase. This diagram reveals the main tendency in theboiling process: the ore metals are rapidly accumulatedin the liquid phase during boiling. The analysis of themodel results suggests that different factors are respon-sible for the increase in the concentrations of variousmetals. The main carriers of iron in the model systemare chlorite, epidote, and actinolite, and the increase intotal iron concentration in the solution is explained byits enhanced complexation with chloride owing to anincrease in solution mineralization. Before the begin-ning of boiling, the behavior of chalcophile elements(Cu, Zn, and Pb) is controlled by the presence of hydro-gen sulfide through sulfide solubilities. The increase intheir concentrations during boiling is related to theremoval of H2S into the vapor phase (mechanism of theloss of precipitant). This effect varies from element toelement. Zinc is characterized by the highest mobilityin the model system and is completely extracted intothe solution at only a small degree of evaporation(≈10%). The disappearance of sphalerite from the equi-librium assemblage produces an inflection in the Znconcentration trend (Fig. 56f), and its subsequent accu-mulation is related to the boiling off of the liquid phase.Lead concentration increases owing to a combinedeffect of the loss of precipitant (hydrogen sulfide) andthe enhancement of complexation. The effect of boilingis most pronounced for the copper concentration. Inaddition to the loss of precipitant and stronger com-plexation, the behavior of copper is influenced by adecrease in Fe++ activity due to the oxidation of the sys-tem. At P = 130 bar, about 1/3 of the total amount ofcopper occurring in the system migrates into solution.If we compare the states of the system before the onsetof boiling (calculation for 166 bar) and after extensiveboiling off (130 bar), the total mineralization of thesolution increases by a factor of about 10, the concen-tration of H2S decreases by a factor of eight, Fe and Znincrease by a factor of 30, Pb increases by a factor of
150, and Cu increases by a factor of 700. The initialrock suffered almost complete leaching of sulfides.Thus, phase separation appears to be most importantfor the least mobile ore metal, copper.
The results presented here on the influence of boil-ing on the properties of isothermal systems cannot bedirectly correlated with the sequential development ofa natural process, because gradual decompression in aclosed system could hardly be expected in nature. Adecrease in hydrostatic pressure owing to tectonic pro-cesses is much slower than the formation of hydrother-mal sedimentary ores. In contrast, pressure changes dueto opening of channels for solution flow or boiling in aliquid column must have catastrophic consequencessimilar to phreatic explosions. The model resultsobtained here should be regarded as a qualitativeassessment of the influence of boiling on the transpor-tation of metals by solutions formed in the high-tem-perature focus of a hydrothermal system.
6.4.2. Phase separation in contactwith a rock (open system)
The gradual separation of a vapor phase from a solu-tion (stepwise distillation) is known to be a more effi-cient fractionation mechanism. The Rayleigh distilla-tion law is widely used for the description of variousgeochemical processes. As to the multicomponent sys-tem addressed in this study, the Rayleigh distillationequation cannot be directly used because of the vari-ability of fractionation coefficients due to phase transi-tions in the mineral part of the system. We performed,therefore, numerical simulation of phase separationwith a pressure increment of 1 bar. The fraction of thevapor phase generated and released at such a pressureincrement is only about 10% (thus, this process is notmuch different from the continuous Rayleigh model).The composition of the model system at a given stepwas equal to the composition of solution from the pre-vious step. Such a scenario can be viewed as a slow
166 164–8
–7
162 160 158 156 154
–6
–5
–4
–3
–2
Pressure, bar
logC
[m
ol]
Sph
Gn
Fractional
Closed system
FeSCuZnPb
SCu
Beginning of boilingseparation
Fig. 57. Concentration of ore elements in solution during the fractional separation of the vapor phase at T = 350°C and R/W = 0.1.For the sake of comparison, the data on S and Cu behavior during boiling in a closed system are shown by dashed lines. The arrowsmark the disappearance of sulfide minerals.
S292
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
ascent of boiling liquid accompanied by gradual phaseseparation within the high-temperature zone of the sys-tem. At each step 10 g of fresh rock per one kilogram ofinitial fluid were added to the system. This addition issufficient to buffer the pH value of fluid but is still toolow to change the metasomatic mineral assemblage(Ab + Ep + Chl + Act) through an increase in R/W.
The simulation of the influence of phase separationon the behavior of metals in boiling hydrothermal solu-tions is shown in Fig. 57. In accordance with the theory,the gradual separation of the vapor phase efficientlyremoves H2S from the liquid phase. Already at P = 155 barand a degree of evaporation of 50%, the liquid phaseretains only 0.8% of the initial amount of sulfide sulfur(8.4% in the above-described closed-system model).Correspondingly, the ore metals are also mobilized atsignificantly lower pressures. Zn is completelyextracted from the rock at P = 161 bar (at a water lossof 9%); and Pb, at P = 157 bar (at a water loss of 32%).When the concentration of the solution increases by afactor of two (P = 155 bar), 26% Cu is removed fromthe rock, whereas the same degree of boiling off in aclosed system corresponds to the very beginning ofcopper mobilization. It can be clearly seen in Fig. 57that the concentrations of H2S and Cu change muchmore drastically when the vapor phase is removed thanin a closed system. The concentrations of zinc appearedto be practically identical, because this element wascompletely extracted into solution in both scenarios.The concentrations of Fe are always controlled bysolid-phase assemblages and are insensitive to volatileloss.
Hydrogen and methane also escape much more rap-idly from solution in the scenario with phase fraction-ation. However, the pHT value of the solution is practi-cally constant (6.4 ± 0.05) within the range of modelconditions owing to the buffering effect of solid phases.
Thus, boiling accompanied by the removal of thegas phase can lead to extensive heavy metal mobiliza-tion from the rock. This result agrees with the conclu-sion by Drummond and Ohmoto (1985).
6.4.3. Adiabatic cooling of two-phase fluid
The rapid movement of hydrothermal solutions inthe upwelling limb of the convection cell, which isestablished from many observations (Chapter 3), mini-mizes heat exchange between the solution and channelwalls, and the T–P–x parameters of solution vary nearlyadiabatically. Because the conduit is open and the solu-tion can expand freely, this process can be regarded asisentropic (Bischoff and Pitzer, 1985): hydrothermalfluid expands during decompression and the energy forwork against external pressure and the latent energy ofphase transformations are supplied by the system cool-ing without any external heat sources. In such a rapidprocess, chemical interactions between the heteroge-neous fluid and wallrocks can be ignored and only min-
eral deposition from the cooling solution needs to beconsidered.
Pressure is the only independent parameter of thismodel, and all other parameters can be expressed asfunctions of pressure through the condition of entropyconservation in the system. Such problems cannot beeasily simulated using programs designed for the calcu-lation of isobaric–isothermal equilibria (such asGIBBS, GBFLOW, and PENG), especially if boilingoccurs in the model. A simple method for the approxi-mate solution of isenthalpic problems was proposed byBychkov and Grichuk (1991). In this method, an iso-thermal series of states of the system is calculated, anda state showing the target value of thermodynamicpotential (H or S) is selected from them.
The entropy of a heterogeneous system is an addi-tive function of the masses of phases:
S = mlSl + mvSv,
where the subscripts l and v denote liquid and vapor,respectively. The entropies of phases depend on theircompositions. Their direct calculation would make theproblem cumbersome and inefficient, because most ofthe substances in this problem are trace components,which make negligible contributions to the bulkentropy of the system. The H2O–NaCl system (Bischoffand Pitzer, 1985) is a suitable prototype for the modelof hydrothermal solution boiling, which allows us tosimplify the task. Given the properties of the binarysystem, the mass proportions of phases providing thetarget value of the entropy of the system can be calcu-lated, and then a state of the multicomponent systemwith the same phase proportions can be found.
The thermal properties of the H2O–NaCl systemalong the boiling surface at temperatures up to 300°Cwere tabulated by Haas (1976a, 1976b), and the entro-pies of phases under higher parameters of state werereported by Tanger and Pitzer (1989). The results ofcalculations of isothermal boiling in contact with a rockat 350°C and pressures of 160 and 150 bar (composi-tions 1 and 2 in Table 31) were used as initial states forthe simulation of adiabatic cooling.
The calculated parameters of adiabatic cooling witha temperature step of 50°C for model solutions 1 and 2are shown in Table 32. These calculations revealed aninteresting feature of adiabatic cooling in a water–saltsystem. Composition 1 initially contained 20 wt % ofthe vapor phase, and adiabatic decompression resultedin an increase in the amount of vapor up to 40%, i.e.,boiling continued in the heterogeneous system duringits ascent. The same change in pressure increased thefraction of liquid in composition 2, i.e., partial vaporcondensation occurred in that case. This seemingly par-adoxical result is related to the high initial content ofvapor in composition 2, about 70 wt %. The gas phaseexpands considerably as the pressure decreases (at200°C it occupies up to 99.4% of the system volume).The work of vapor expansion results in its partial con-
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S293
densation (mass fraction of vapor decreases from 70 to60%).
The results of the thermodynamic modeling ofchanges in the composition of solution during isen-tropic cooling are shown in Fig. 58. These results indi-cate that isentropic cooling is accompanied by the dep-osition of ore minerals (Fig. 58a): pyrite, sphalerite, andchalcopyrite. The sulfides are deposited despite a two-fold increase in the fraction of the vapor phase in themodel within the cooling interval (Fig. 58b), and thecontinuous partitioning of hydrogen sulfide into thevapor phase (Fig. 58c). Simultaneously, the amount ofiron in the solution decreases by a factor of two; zinc,by an order of magnitude; and copper, by two orders ofmagnitude (Figs. 58c, 58d). The concentration of leaddoes not change, and the system remains unsaturatedwith respect to galena. These changes are accompaniedby a significant decrease in pHT from 6.5 to 3.9(Fig. 58e) and a decrease in H2 activity from 8 × 10–3 to1 × 10–5. The characteristic maximum of pyrite deposi-tion at 200°C is related to the fact that the concentrationof Fe in the residual solution becomes higher than thatof H2S below this temperature. Note that the totalamount of hydrogen sulfide in the heterogeneous sys-tem does not vary considerably.
Similar results were obtained for the adiabatic cool-ing of initial composition 2, which was produced byisothermal boiling at a pressure of 150 bar (Fig. 59c).Pyrite is again the most abundant phase. A specific fea-ture of mineral formation for this vapor-rich initialcomposition is that sphalerite saturation is achievedonly after cooling to 150°C.
Figures 59a–59c compare the results of simulationof mineral formation during the stepwise cooling of ahomogeneous solution and heterogeneous fluids. Fig-ure 59a shows a cooling path of the model solution thatwas produced in the simulation of isothermal boiling at350°C (Section 6.4.1) and P > Pboiling (166 bar). The T–P parameters during cooling of this composition werevaried along the boiling line within the homogeneousfield, and the calculation procedure corresponded to theslow cooling scenario (Section 4.3.1). Figures 59b and59c display the results of the adiabatic cooling of heter-ogeneous fluids described above. It can be seen that theadiabatic cooling of initially heterogeneous fluidsyields much larger volumes of ore precipitates (in thecases considered, by a factor of 6, on average). Themain contribution is that of pyrite, while sphalerite con-tent shows only a twofold increase. Compared with thecase of a homogeneous solution, the mass of chalcopy-rite deposited during the cooling of heterogeneous fluidincreases by factors of 5 and 20 for compositions 1 and2, respectively. In the homogeneous model, the massfraction of chalcopyrite in the bulk precipitate was0.36%, whereas this value is 0.38% for composition 1(showing a fivefold increase of the total mass of sul-fides) and 1.5% for composition 2.
The main reason for the enhancement of ore deposi-tion from heterogeneous fluids is the higher initialmetal content in them compared with homogeneousfluid. This can be seen in Fig. 56f, which displays theconcentrations of ore components in the initial solu-tions that subsequently underwent cooling in themodel. The considerable loss of metals from solution
Table 31. Initial compositions of hydrothermal fluids in themodel of adiabatic cooling obtained by the calculation of iso-thermal boiling in contact with rock at T = 350°C, ΣR/W =0.596, and wave no. 1
Composition no. 1 2
Pinit 160 150
Gas/liquid, mass ratio 0.253 2.22
Liquid phase
Mass, kg 0.810 0.315
Salinity, g/kg 40.72 103.3
pHT 6.47 6.60
Amount of component, mole
H 0.035424 0.008811
O 0.023567 0.007291
K 0.032676 0.032676
Na 0.450453 0.4481
Ca 0.030586 0.031363
Mg 1.21 × 10–8 7.52 × 10–9
Fe 0.001459 0.00214
Al 1.17 × 10–5 6.43 × 10–6
Si 0.004528 0.001152
C 0.001405 0.000169
S 0.0031 0.000445
Cl 0.545898 0.54589
Cu 2.07 × 10–6 9.7 × 10–6
Zn 0.000135 0.000135
Pb 2.28 × 10–7 8.45 × 10–7
H2O* 43.11551 15.65974
Gas phase
Mass, kg 0.205 0.699
Amount of component, mole
H 22.73958 77.67893
O 11.33757 38.80279
C 0.006169 0.007405
S 0.007144 0.009805
Cl 2.15 × 10–6 1.02 × 10–5
Equilibrium mineralassociation
Gn, Bn, Ab, Ep75, Ep60, Chl75, Act80
Bn, Ab, Ep75, Ep60, Chl50, Chl75, Act80
S294
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
suggest that temperature is a more significant factor forthe behavior of ore elements than hydrogen sulfide par-titioning into the vapor phase. On the other hand, acomparison of residual concentrations in ore-formingsolutions shows that phase separation is favorable formetal transportation by hydrothermal solutions. Withinthe whole temperature interval, the concentrations ofmetals in the liquid phase of heterogeneous systemswere higher than those in the homogeneous system, andtheir temperature dependencies were less steep. This isespecially pronounced for Zn concentrations(Fig. 59d).
However, in order to estimate the applicability of themodel results to natural objects and the role of adiabaticcooling in ore formation to convective systems in gen-eral, it is necessary to evaluate the thermal and hydro-dynamic aspects of this process. It is known that ther-mal convection is driven by the difference of hydro-static pressures generated by water columns in thedownwelling and upwelling limbs. This differencearises from the density contrast between cold and hotwater. Phase separation results in a dramatic decrease inbulk fluid density owing to the large molar volume ofthe vapor phase. According to the data of Haas (1976a,1976b), at 300°C and a bulk NaCl content of0.5 mol/kg, the densities of homogeneous solution andtwo-phase systems with 20 and 70 wt % of vapor are0.744, 0.180, and 0.059 g/cm3, respectively. Conse-
quently, the formation of a boiling zone in theupwelling limb and the focus of a hydrothermal systemmust strongly intensify solution circulation. Somehydrodynamic aspects of this problem were discussedby Cathles (1977) and Brikovski and Norton (1989).The hydrostatic pressure of the heterogeneous fluid col-umn in the channel is small compared with the seawaterpressure at the mouth of the channel. For a verticalchannel size of 1 km, an ocean depth above the channelmouth of 1.6 km, a temperature of 350°C, and a vaporfraction of 20 wt %, its contribution will be only 30 baras compared with 103 bar for cold seawater. As a result,the channel filled with ascending boiling fluid serves asan efficient intake manifold. The reduced pressure inthe channel will inevitably intensify the lateral influx ofcolder water from fractures in the country rocks andsubsurface mixing. This factor must strongly limit thedevelopment of boiling zones, because even a smalladdition of cold water will promote vapor condensationand cessation of boiling.
On the other hand, the adiabatic cooling of hetero-geneous fluid from 350 to 300°C corresponds to a pres-sure decrease from 165–150 to 85–80 bar (Table 32).Given a density of the heterogeneous mixture of0.2−0.3 g/cm3, the length of the ascending flow channelmust be 2–4 km. This is close to the maximum possiblevalue for oceanic hydrothermal systems (perhaps alsofor ancient hydrothermal systems producing massive
Table 32. Parameters of the adiabatic cooling of solutions
Temperature, °C Pressure, barMasses of phases, g Bulk entropy
of system, J/KSalinity
of liquid, g/kgliquid gas
Adiabat for composition 1 (Table 31); mass of system 1014.3 g and salt content 32.97 g
350 160.56* 809.7 204.6 3961.591 40.72
300** 83.38 693.2 321.1 3961.601 47.56
300*** 83.18 703.3 311.0 3961.597 46.88
Average 83.28 698.2 316.0 3961.599 47.22
250 38.58 675.5 338.7 3961.585 48.81
200 15.05 627.6 386.7 3961.594 52.54
150 4.59 583.4 430.9 3961.587 56.52
Adiabat for composition 2 (Table 31); mass of system 1014.0 g and salt content 32.50 g
350 153.66* 314.6 699.4 4796.890 103.3
300** 80.87 369.4 644.5 4796.888 88.0
300*** 80.84 373.6 640.4 4796.881 87.0
Average 80.85 371.5 642.5 4796.885 87.48
250 37.81 424.1 589.9 4796.899 76.63
200 14.79 426.0 588.0 4796.898 76.29
150 4.52 417.4 596.6 4796.897 77.86
* Calculated pressure for the H2O–NaCl system for the gas/liquid ratio obtained for the initial composition by the simulation of a mul-ticomponent system.
** Calculated using the data by Haas (1976a, 1976b).*** Calculated using the data by Tanger and Pitzer (1989).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S295
sulfide deposits). The lengths of real vertical channelscorrespond to temperature variations of only 15–25°Cduring adiabatic discharge. These values are similar tothe estimates obtained by Bischoff and Pitzer (1985).The model calculations showed that the maximum ore
deposition from heterogeneous fluids occurs at muchlower temperatures and more extensive cooling. Thenatural scales of adiabatic cooling allow rapid ascent ofheterogeneous solutions through the conduit almostwithout loss of ore components.
Fig. 58. Compositions of hydrothermal solutions during the adiabatic (isentropic) cooling of a heterogeneous fluid. The initial com-position of solution is given in Table 31 (no. 1), and the adiabat parameters are listed in Table 32. (a) The deposition of mineralsduring adiabatic cooling; (b) the degree of water evaporation, %; (c) iron and sulfur in solution; (d) heavy metals in solution; and(e) solution pHT.
0
20
40
60
80
100
0
0.002
0.004
0.006
0.008
0.012
Fe
H2S
–8
–7
–6
–5
–4
–3
350 3003
4
250 200 150
5
6
7
Temperature, °C
CuZnPb
(c)
(d)
(e)
3000
250 200 150
40
60
80
20
Mas
s, m
g/k
g
SphCcpDph
(a)
solu
tion
Py
Fra
ctio
n o
f li
quid
wat
er, %
mol/
kg
pH
logC
[m
ol/
kg]
(b)
S296
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Thus, the model of an adiabatic (isentropic) coolinghas limited applications for the reproduction of ore for-mation in the ocean. The natural process either does notfollow an adiabatic path (owing to subsurface mixing)or produces no significant deposition of ore matter.
6.4.4. Condensation of two-phase fluid during rapid discharge and ore deposition
As was shown in Section 4.3.1, ore deposition dur-ing rapid discharge of hydrothermal fluids can be repro-duced in the scenario of instantaneous cooling of
hydrothermal solution without any chemical interac-tion with seawater. The cooling of a heterogeneousfluid will be accompanied by the condensation of thevapor phase and the interaction of vapor-borne sub-stances with components from the liquid phase. Thatthis case is realistic is suggested by the fact that fluidheterogeneity was never visually observed during thedischarge of solutions in black smokers with definiteevidence of boiling.
Within the scenario of instantaneous cooling, modelcalculations were performed for heterogeneous fluidformed (a) in the case of isothermal boiling without
(a)
Homogeneous solution
80
60
40
20
0
Mas
s, m
g/k
g s
olu
tion
Dph
Ccp
Sp
Py
Dph
Ccp
Sp
Py
300 250 200 150
Po
Ccp
Sp
Py
(b)80
60
40
20
0
Mas
s, m
g/k
g s
olu
tion
(c)80
60
40
20
0Mas
s, m
g/k
g s
olu
tion
(d) –3
–4
–5
–6
–7
logC
[m
ol/
kg]
300 250 200 150350Temperature, °C
Heterogeneous solution (no. 1)
Heterogeneous solution (no. 2)
Po
Ccp
Sph
Py
Py
Dph
CcpSph
Sph
Py
SphCcp
Dph
Fig. 59. Compositions of ore precipitates and hydrothermal solutions during the cooling of homogeneous and heterogeneous fluids.(a) Homogeneous initial solution, (b) 20% of vapor phase, (c) 70% of vapor phase; and (d) residual concentrations of zinc in solutionduring fluid cooling.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S297
phase separation (Section 6.4.1, Fig. 56) at a varyingfraction of the gas phase and pressures of 166, 160, 150,and 130 bar; and (b) in the case with phase separation(Section 6.4.2) for various degrees of boiling off.
The temperature of cooling was taken to be 150°C,similar to the calculations presented in Section 4.3.1.The pressure at this temperature was taken to be 100 barin order to provide the complete condensation of thevapor phase.
The results of fluid cooling simulation withoutphase separation are shown in Fig. 60a together withthe cooling of the homogeneous solution that was pro-duced in the isothermal model at P = 166 bar. As can beseen in this diagram, within a wide range of vapor con-tent (20–70 wt %), the cooling and condensation of het-erogeneous fluid do not significantly affect the mass ofprecipitate but change its mineral composition toward ahigher content of heavy metal sulfides. The fraction ofsphalerite in the precipitate is 3% in the case of homo-geneous solution and increases up to 10–13% duringthe cooling of heterogeneous fluids; the content of chal-copyrite increases by a factor of 20, from 0.06 to 1.3%.Similarly, the precipitate from a heterogeneous fluidcontains up to 0.13% galena as compared with 0.01%calculated for a homogeneous solution. This enrich-ment is related to an increase in the concentrations ofheavy metals in hydrothermal fluids owing to phaseseparation (Section 6.4.1, Fig. 56f). An interesting
effect was observed during the cooling of fluid with themaximum vapor fraction (89 wt %). In that case, pyrite,which was predominant in all other cases, was replacedby pyrrhotite. The phenomenon is connected with thestrong partitioning of H2 into the vapor phase duringboiling in the focus of the hydrothermal system(Fig. 56g). Although the molar fraction of hydrogen inthe vapor phase is not high (about 0.1%), the vaporphase is abundant in this model. In the bulk system, theamount of H2 increases up to 49 mmol as comparedwith 14 mmol in the homogeneous case. The condensa-tion of such a gas phase results in the precipitation of amore reduced phase, pyrrhotite. This process also pro-duced the maximum chalcopyrite fraction of 3%.
Thus, the condensation of heterogeneous fluids isaccompanied by the precipitation of ore minerals, andthe main feature distinguishing this case from the cool-ing of homogeneous hydrothermal solutions (withoutboiling) is the elevated fraction of heavy metals, Cu,Zn, and Pb.
Interesting results were obtained from the simula-tion of rapid cooling of solutions that suffered frac-tional separation during boiling (Fig. 60b). A minorvapor loss (5–10% of the mass of water) increases theamount of ore precipitate, but a high degree of boilingoff has the opposite effect. This is surprising, becausethe concentrations of ore metals in the solution increaseduring phase separation (Fig. 57). The reason for this
(a)
(b)
Homogeneous20% vaporsolution
Homogeneous
70% vapor 89% vapor
100 mg
5.4% system vapor loss
30.5% 12.2% 40.0% 47.4% 53.5%
Sph Prl
Py
PrlCcpSph
PyPy
GnSph
CcpKln Gn
SphCcp
Po
Po
Gn
Sph
Sph
CcpBn
Dph
Fe-TrFe-Tr
CcpDph
DphFe-Tr
Ccp
SphGnPy
Po
Sph
Ccp Mc
Ccp PrlSph
Py
PyPy
SphSph PrlPrlCcp
Fig. 60. Compositions of ore precipitates obtained during the rapid cooling of a heterogeneous fluid. The initial compositions offluids were obtained by the simulation of boiling in contact with a rock at T = 350°C (Figs. 56, 57). (a) Cooling after boiling withoutphase separation (initial compositions were obtained at P of 166, 160, 150, and 130 bar); and (b) boiling with phase separation (ini-tial compositions were calculated at P of 166, 162, 160, 158, 157, 156, and 155 bar). Circle area is proportional to the mass of pre-cipitated matter for 1 kg of fluid.
S298
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
effect is related to the deficit of sulfide sulfur: H2S iseliminated with vapor during phase separation. Alreadyat 20% evaporated, the residual H2S concentrationbecomes equal to the total content of iron and heavymetals, and the advanced phase separation generates anSII deficit in the solution. Because of this, at largedegrees of evaporation, the amount of sulfides pro-duced by cooling of the residual solutions decreasescontinuously. Pyrite is replaced by pyrrhotite (althoughthe environment does not become more reducing thanin the beginning of boiling) and later by Fe-chlorite andFe-actinolite. The preservation of Zn and Cu sulfides inthe precipitate suggests that they are less soluble thaniron sulfides under such conditions. After 50% evapora-tion, cooling of the solution produces sphalerite, chal-cocite, and gangue minerals.
6.4.5. Verification of the model
A comparison of the data obtained by the simulationof boiling in the interiors of hydrothermal systems withnatural observations shows that the model provides anadequate description of the main properties of hydro-thermal solutions that are usually connected with boil-ing. The model implies extensive partitioning of vola-tile components into the vapor phase during boiling.The liquid phase is simultaneously enriched in iron andheavy metals; this effect is related not only to evapora-tive concentration but also (and to a larger extent) tometal mobilization from the country rock. This resultsin a positive correlation of iron and heavy metal con-tents with chloride and a negative correlation with H2S.Such relationships exist in natural hydrothermal sys-tems (Section 6.2). The concentrations of metals inmodel solutions are of the same orders of magnitude asthose in the natural prototypes: from n mmol/kg to20 mmol/kg of Fe; 0.n mmol/kg of Zn; 0.0n mmol/kgof Cu, and 0.00n mmol/kg of Pb (Figs. 52, 56, 57).Exact quantitative correspondence between the modeland the natural prototypes could hardly be expected,because the temperature of our simulation (350°C) islower than that in the interiors of oceanic hydrothermalsystems (370–400°C). Nonetheless, the calculated andobserved trends in the development of boiling pro-cesses are evidently identical.
6.5. Conclusions
Analysis of the results of simulated boiling in oce-anic hydrothermal systems showed that this factor isfavorable for the transfer of metals in hydrothermal sys-tems. Phase separation in solutions results in the parti-tioning of metals and hydrogen sulfide into differentphases and a significant increase in the solubility of sul-fides in heterogeneous systems. This effect correspondsto the action of the mechanism of precipitant loss. Itmust evidently influence not only the elements consid-ered in our simulation (Fe, Cu, Zn, and Pb) but also anyelement with a sulfide form of precipitation and a non-
sulfide form of migration in solutions. Among the ele-ments considered in our model, copper is most sensitiveto boiling, followed by zinc and lead. Boiling in theinterior of a hydrothermal system can increase the con-centrations of these metals in solution by up to twoorders of magnitude. If the residual brine and vapor arespatially isolated, the mobilization of metals from therock will further increase. Consequently, boilingincreases the potential of hydrothermal systems withrespect to base metal ore formation.
The cooling of heterogeneous solutions accompa-nied by condensation (during submarine discharge) canyield considerable amounts of sulfide precipitates.They will be enriched in heavy metals relative to theores formed by the cooling of homogeneous fluids.
CHAPTER 7. POSSIBLE ROLE OF MAGMATIC FLUID: A COMBINED
EXHALATION–RECYCLING MODEL
7.1. Problem Formulation
Magmatic fluids are traditionally regarded by geol-ogists as the main ore generating agent in hydrothermaldeposits. This point concerns, in particular, massivesulfide deposits. Starting from the pioneering studies byZavaritskii (1943), a volcanic exhalation hypothesisand its modifications have been used in genetic recon-structions by many authors. However, an analysis of theliterature on this problem shows that this concept isbased on the observations of spatial and temporal linksbetween magmatism and ore mineralization, whereasthere are no direct lines of evidence for the formation ofmassive sulfide ores from magmatic gases. The findingand investigation of modern ore-forming systems onthe ocean floor resulted in the development of an alter-native recycling concept for the genesis of massive sul-fide ores (Rona, 1984; Krivtsov, 1987; etc.). Seawatercirculating in heated rocks is regarded as the main ore-forming agent in this concept. Both heat and ore com-ponents are supplied to the hydrothermal system fromthe country rocks. The exhalation and recycling con-cepts are not mutually exclusive; a superposition ofthese processes can be expected during functioning ofa magma-related hydrothermal system. Hence, there isa need for indicators of the exhalation and recyclingmechanisms which would allow us to distinguish theircontributions to the formation of massive sulfide depos-its.
The investigations of modern hydrothermal systemsin oceans have not yet provided unambiguous evidencefor the presence of a significant exhalation componentin the ore-forming process (Chapter 3). However,almost all the oceanic hydrothermal systems studied todate are associated with divergent plate boundaries, andthe melts reaching the crust in these regions are usuallywater-poor. According to the results of geodynamicreconstructions, most ancient massive sulfide depositswere formed in island-arc settings with a different style
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S299
of magmatism, a different type of magma chambers(central-type volcanoes rather than fissure eruptions), andmagmas richer in water. Therefore, it would be inappropri-ate to extend the conclusions obtained for oceanic hydro-thermal systems to ancient massive sulfide systems.
A paradoxical situation is revealed by the analysisof the literature on the problem of the volcanic exhala-tion origin of massive sulfide ores: the researchersnever addressed the problem of quantitative estimationof the chemical effect of magmatic gases on the wall-rocks of massive sulfide deposits and characterizationof the ores that could be formed during magmatic fluidejection on the seafloor. One reason for this paradox isthe lack of methods allowing for such quantification.Therefore, a search for new ways to tackle this problemis very important. The thermodynamic modeling ofore-forming processes is a promising approach.
In this section we address calculation of the thermody-namic model of a submarine hydrothermal system whoseevolution included the stages of magmatic exhalation andconvective recycling. This simulation was aimed at esti-mating the relative contributions of these mechanisms tothe formation of submarine massive sulfide deposits.
7.2. Characteristics of the Exhalation Convective Model
The caldera of a submarine volcano filled with basicextrusive rocks and underlain by a magma chamber wasregarded as a geologic prototype of the model (Fig. 61).During the first (magmatic) stage of development, thehydrothermal system is fed by the fluid released from themagma. As the fluid ascends toward the surface, it reactswith the country rocks forming a footwall channel, andore minerals are deposited during fluid discharge on theocean floor. The second (convection) stage produces arecycling hydrothermal system within the volcano sup-plied by seawater. This system inherits the dischargeconduit of the magmatic stage, and the ore-forming pro-cess is superposed on the products of previous hydro-thermal activity. Since there is no geological informationon the mechanism of the change of system supply, thestage of a mixed magmatic–recycling source is not con-sidered in the model, which simplifies the interpretation.
Model parameters are the compositions of rocks, mag-matic gas, and seawater; temperature and pressure in thedownwelling limb of the convection cell, feeder channel,and ore deposition zone; intensities of solution interactionwith the country rocks (specified by the rock/water ratio);production of the system during the magmatic and recy-cling stages; and durations of these stages.
The model composition of the country rocks wasapproximated by the average composition of calc-alka-line basalts, and present-day seawater composition wasused in the calculations. The estimation of the compo-sition of magmatic gas is discussed below. For the sakeof comparison, the ranges of temperature, pressure, R/W
ratio, and discharge rate of the hydrothermal system weretaken to be identical to those used in Chapter 4.
The mass proportions of magmatic and recycled flu-ids can be estimated from the heat balance of the hydro-thermal systems, by calculating the amount of seawaterrequired to cool the unit mass of magmatic melt andcomparing it with the abundance of water dissolved inthis melt. The heat released during the formation ofigneous rocks includes the latent heat of crystallizationand thermal effects of solid rock cooling. The proce-dure of heat balance calculation is similar to thatdescribed in Chapter 3, but the two contributions areconsidered separately:
=
= 0.25 + 0.46 = 0.71 g H2O/g rock.
The concentration of water in the initial island-arc meltwas taken to be 1% or 0.01 g per one gram of basalt. If thiswater is completely released during crystallization (whichis certainly an overestimation) and there is no explosivedegassing, the mass proportion of magmatic and convec-tive components will be 1 : 70. It should be kept in mindthat this proportion is valid for the whole period of magmachamber cooling. Cann et al. (1985/1986) calculated thathydrothermal activity is vigorous when melt is present inthe magma chamber and attenuates gradually after meltdisappearance. Therefore, the effective mass proportionwill be less contrasting during the active phase of the pro-cess (but no higher than 1 : 25).
The discharge rate of the system in the convectionstage was taken to be 10 kg/s, which is equal to that ofthe model described in Section 4.3.2. In order to facili-tate comparison, the same discharge rate was assignedto the magmatic stage, which is not constrained by nat-ural data. Therefore, the proportion of the durations ofthe magmatic and recycling stages corresponded to theratio of solution masses (no less than 1 : 25).34 The sim-
34 An increase in the prescribed discharge rate value makes themagmatic stage shorter without changing the mass of releasedfluid and the corresponding mass of transported ore matter.
MH2O0.4
370 × 0.0044-------------------------------- 0.75
370 × 0.0044--------------------------------+
Magmatic fluid
Magma chamber
Fig. 61. Principal scheme of the model of a convective sys-tem involving magmatic fluids.
S300
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
ulation was carried out with a time increment of 108 s(about 3 y), and the lifetime of the model system wasn × 1000–n × 10 000 y. The product of the dischargerate and the lifetime of the system is equal to the totaloutput from the system, which controls the final size ofthe growing ore body.
Given the approximate character of many parame-ters, only those that affect the qualitative results of sim-ulation are of real significance. An analysis of our sim-ulation distinguishes a limited number of such modelparameters: the maximum temperature in the interiorpart of the system, ΣR/W value in the downwellinglimb, and relationships of the discharge rate of the sys-tem and the duration of the process. Variations in otherparameters influence the quantitative results of simula-tion but do not change the general qualitative patterns.
Initial composition of magmatic gas. The determi-nation of the initial composition of magmatic fluid is aconsiderable challenge, because there is no direct evi-dence for massive sulfide deposits. The composition ofvolcanic gases from island-arc volcanoes can be used asa proxy. The most comprehensive and reliable resultson the composition of volcanic gases were recentlyobtained for Kudryavyi Volcano (Iturup Island, Kurils)by Taran et al. (1995). These authors managed to sam-
ple fumaroles with a temperature of up to 940°C. Thegas from these events changed negligibly after its sepa-ration from a melt. The roof of the magma chamber ofKudryavyi Volcano is probably situated only 200–300 m below the bottom of the crater. The composition ofmagmatic gas from this volcano is shown in Table 33.
It appeared, however, that simulation with mag-matic fluid of such a composition does not producemassive sulfide ores. Reconnaissance calculationsyielded mineral assemblages containing considerableamounts of anhydrite and elemental sulfur, which arecommon in subaerial exhalative volcanic deposits. Thisresult is due to the SO2-rich composition of gas fromKudryavyi Volcano. The reason for this effect lies in theconditions of gas separation from magma. Wallace andCarmichael (1994) showed that the relationships of oxi-dized and reduced sulfur species in a gas equilibrated withmagma depend on the redox conditions and total pressurein the system. The equilibrium constant of the reaction
SO2 + 3H2 H2S + 2H2O (59)
gives
/ = K59Ptot( / ). (60)XH2S XSO2XH2
3 XH2O2
Table 33. Compositions of fluids used in the model, mol/kg
Compo-nent
Volcanic gas (Taran et al., 1995) Convecting solution (calculated for the downwelling limb)
initial composition composition corrected to adegassing pressure of 500 bar
beginning of convection(calculation step no. 200)
end of convection(calculation step no. 900)
H 1.626999 3.436776 0.071243 0.071482
O 3.747455 1.201559 0.072429 0.072470
K 0.001400 0.001400 0.019807 0.005699
Na 0.002500 0.002500 0.476855 0.471885
Ca 0.026262 0.026262 0.023889 0.033377
Mg 0.000020 0.000020 0.000010 0.000016
Fe 0.000131 0.000980 0.001194 0.001393
Al 0.069808 0.069808 0.000008 0.000007
Si 9 × 10–8 9 × 10–8 0.014545 0.015788
C 0.9371 0.937100 0.004330 0.002624
S 1.030381 0.189000 0.001188 0.000480
Cl 0.252391 0.252391 0.545900 0.545900
Cu 0.000004 0.000004 0.000012 0.000031
Zn 0.000047 0.000047 0.000122 0.000033
Pb 0.000006 0.000006 0.000002 0.000001
H2O 49.14576 48.200000 55.327085 55.405107
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S301
In order to take this factor into account, we refinedthe composition of magmatic fluid. It was assumed thatgas separating from the melt at a temperature of1100°C is in thermodynamic equilibrium with theQFMP (quartz–fayalite–magnetite–pyrrhotite) buffer.The composition of gas from Kudryavyi Volcano in theH–O–C–S–Fe–SiO2 system was used for simulation(with excess solid buffer phases). The calculations wereperformed by the GBGAS program using the approxi-mation of ideal mixing of real gases, and the thermody-namic properties of substances were taken from theUNITHERM data bank. Figure 62 shows the results ofthe simulation, indicating that an increase in the degas-sing pressure strongly reduces the molar fraction ofSO2 in the gas, whereas the molar fraction of H2Sremains constant. Hence, the SO2/H2S ratio in a deepmagma chamber will be lower than that observed involcanoes. The simulation of the formation of a mas-sive sulfide system was carried out using the modelcomposition of magmatic gas, whose major-componentconcentrations were corrected to a degassing pressureof 500 bar.35 This composition is given in Table 1.
The interaction of magmatic fluid and convectingseawater with the walls of a feeder conduit was calcu-lated by the procedure described in Section 4.2. Theformation of an ore body was simulated by the methodsimilar to that used in Section 4.3.2, except that duringthe first 100 steps of calculation, magmatic gas wasintroduced into the channel and altered seawater fromthe downwelling limb (its composition was calculatedby the method described in Sections 2.2 and 4.1.2) wasused in subsequent steps.
7.3. Results of Simulation 7.3.1. Formation of footwall metasomatic zoning
According to our simulation, the interaction of mag-matic fluid with the walls of the feeder channel corre-sponds to acid leaching. The back zones of resultingcolumns (Fig. 63) are made up of the Prl + Py and Qtz +Prl + Py associations. Away from the channel wall,chlorite and talc appear and are further joined by actin-olite, epidote, wairakite, and albite. The front zones arecomposed of assemblages typical of rocks affected bypropylitic alteration: Chl + Ep + Act + Ab sulfides at350°C, and carbonates occur instead of epidote andwairakite at 300 and 250°C. Another notable feature ofthe calculated columns is the magnesian chlorite com-position (clinochlore) in their assemblages. Copper isfixed as chalcopyrite in the Qtz + Prl + Py zone,whereas zinc and lead form sphalerite and galena,respectively, in the clinochlore-bearing zone.
The results of simulation shed light on the nature ofmetasomatic zone boundaries and the regime of com-ponents within the column. Figure 64 gives some data
35 The concentrations of chlorides and ore metals were not cor-rected.
on the composition of solutions moving through themodel metasomatic column at a temperature of 350°C.The back zones of the column, including the Qtz + Prl +Py zone, are formed in contact with acid solutionsshowing pHT 1.5–1.7. The transition to propyliteassemblages is associated with the neutralization of theacid magmatic fluid by the country rocks (up topHT 4.5–4.6). Sulfide sulfur is efficiently transported bythe fluid up to the frontal zone of propylites andstrongly prevails over ore metals. The sequence of ele-ment mobility in the metasomatic halo is S > Fe, Zn > Cu.
These results are fundamentally different from thoseobtained in the convection model (Section 4.2, Fig. 28).Altered seawater produces chloritization and albitiza-tion in the country rocks. The resulting chlorite has aniron-rich ferromagnesian composition. The pHT value(about 5.4) and proportions of ore elements in the per-colating solution are practically invariable. The hydro-thermal fluid of an exogenous origin passing throughthe convective system weakly reacts with basalt,because it equilibrated with the country rocks in thedownwelling limb. Only a small shift from the equilib-rium state takes place in the feeder channel owing tochanges in T and P. The formation of a distinct footwallmetasomatic zoning is obviously impossible in thiscase.
Consider now what happens when recycling over-prints the metasomatic column formed by a magmaticfluid, i.e., a two-stage model. Figure 65 illustrates thesimulation of the passage of 100 portions of magmaticgas through the metasomatic column (this case isshown in Fig. 64) followed by 100 portions of recycledsolution from the downwelling limb (Tmax = 350°C andΣR/W = 0.47). The main effect of such a superpositionis the formation of sericite during the percolation ofrecycled solution. This produces the typical metaso-matic zoning of massive sulfide deposits: Qtz + Ser +Py → Qtz + Ser + Chl + Py → propylite.
1010–10
10–8
100 500 1000 20001
10–6
10–4
10–2
100
Pressure, bar
Mole fractions of components
** * * *
**
H2CO2
CH4SO2H2SS2
CO
Fig. 62. Calculated composition of a magmatic gas phase inequilibrium with the QFMP buffer at various pressures andT = 1100°C.
S302
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
7.3.2. Chemical evolution of the ore body in the combined exhalation–recycling model
The calculations presented in this section were per-formed for the following scenario of ore-forming pro-cesses. During the first (exhalative) stage, an ore bodywas formed at the expense of magmatic fluid discharge(ejection) on the ocean floor (its composition is shownin Table 33, column 2). This stage lasted 100 computa-tion steps, with a discharge fluid rate of 10 kg/s, whichcorresponds to the degassing of a magma body, about
4 km3 in volume, within 300 y. During the secondstage, degassing ceased and a convection system waslaunched. This system had the same discharge rate, andits activity lasted 2500 steps (about 8 ky).36
Figures 66a and 67 show the structure and bulkcomposition of the ore body formed by the end of themagmatic stage (after 100 computation steps). The cal-culations suggest that the body has a pyritic composi-
36 The proportion of the durations of the two stages, 1 : 25, wasdetermined above from heat balance conditions.
Fig. 63. Mineral composition of metasomatic columns developing around a feeder channel under the influence of magmatic fluidat (a) 350°C, (b) 300°C, and (c) 250°C and P = 500 bar.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S303
tion. It grows rapidly and reaches about 1 Mt by the endof this stage. The concentrations of zinc and copper inthe body are not high, no more than 0.7% Zn and 0.04%Cu. This base metal deficiency in the ore body is a con-
sequence of the increased solution acidity. Magmaticfluids exiting from the channel show pHT values ofabout 1.5 and are, therefore, undersaturated withrespect to base metal sulfides, despite the rather high
0
1
2
3
4
5
pH (a)
1–9
–8
4 7 10 13 16 19 22 25 28 31 34 37 40 43 46 49
–9
–6
Reactor step
–5
–4
–3
–2
–1
–0
S
Fe
Zn
Cu
(b)log(Me) [mol/kg]
Fig. 64. Composition of solution generated during the formation of a metasomatic column corresponding to Fig. 63a at T = 350°Cand P = 500 bar (computation step no. 20). (a) pH and (b) concentrations of dissolved components (logarithmic scale).
1Reactor step
4 7 10 13 16 19 22 25 28 31 34 37 40
100
80
60
40
20
0
%
Py
Dsp
Act80
Prl
Ep70
Cch
Dph
Qtz
Tr
Ser
Ep60
Wai
Chl50
AbAct80
Ep75
Wai
Dph
Chl50Chl50
Ep60
Cch
Ser
Prl
DspQtz
Py
Tr
Fig. 65. Mineral composition of metasomatic rocks formed around a feeder channel in the combined exhalation–recycling modelat T = 350°C, P = 500 bar, ΣR/W = 0.47, after the passage of 100 portions of magmatic fluid followed by 50 portions of recycledseawater.
Ab
S304
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
concentrations of copper, zinc and sulfide sulfur in thesolutions. The precipitation of these components beginsonly after considerable dilution of the fluid with seawa-ter.37
When the ore-forming fluid changes from a mag-matic to a recycling source, the character of the ore pro-cess also changes. The ore body grows several timesmore slowly during the recycling stage (Fig. 67a), andits composition changes dramatically. During computa-tion step nos. 100–350, the ore body rapidly accumu-lates zinc extracted from the downwelling limb of the
37 Another possible mechanism of neutralization is the metaso-matic alteration of calcareous bottom sediments, but it is not con-sidered in our model.
convection cell, and copper gain begins at step no. 720.Figures 66b and 66c show the composition of the orebody in the course of its growth. Zinc is mainly precip-itated in the peripheral part of the ore body (Fig. 66b),whereas copper is accumulated in its hottest centralzone. During the mature stage of development, zinc isgradually removed from the ore edifice, and magnetiteappears near the channel mouth. Figure 66c shows zon-ing formed in the ore body by step no. 950.
7.4. Discussion
The simplified model developed in this chapter isbased on a number of very approximate parameters,and the results of our simulation should be considered
Sph
0
20
40
60
80
100Mineral, %
0 10 20 30 40 50
Ab
Prl
Anh
Qtz
Ccp
Sp
Py
Ab Prl
Prl
Qtz
Py
(a)
(b)0 10 20 30 40 50 60
0
20
40
60
80
100Ab Prl
Qtz
Py
(c)
0 10 20 30 40 50 60 70 80Radius, m
Anh
QtzCcp
Py
0
20
40
60
80
100
Fig. 66. Proportions of minerals in the radial cross-section of an ore body during various stages of its growth. (a) The end of themagmatic stage (computation step no. 100); (b) the beginning of the convective stage (computation step no. 300); and (c) the endof the convective stage (computation step no. 950).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S305
preliminary. Nonetheless, the following conclusionscan be drawn from the qualitative results of calcula-tions.
(1) Hydrothermal systems with different fluidsources will probably be significantly different in thecharacter of metasomatic zoning around the feeder con-duit. It should be least pronounced in the convectivesystems with small contributions from magmatic fluids.
The interaction of magmatic gas with country rocksforms a classic metasomatic column of acid leaching.Such columns are not common in massive sulfidedeposits and were never reported from modern oceanichydrothermal systems.
The mineral assemblages of ancient massive sulfidedeposits were reproduced only when the model columnof acid leaching was overprinted by the action ofaltered seawater.
(2) The magmatic and convection stages of ore bodyformation produce substantially different ore materials.The acid solutions of the magmatic stage rapidly forma pyrite–silica ore body, and the growth rate decreasessignificantly during the convection stage, when copperand, to a lesser extent, zinc are accumulated within thebody.
In general, the available data do not contradict theseconclusions, although there are some discrepancies.
As was noted in Section 4.2, there are only minormetasomatic alterations beneath the ore edifices on theoceanic floor, which is in agreement with the results ofsimulation. Drilling at the TAG hydrothermal siterevealed a small zone of paragonitization directlybeneath the body (Humphris et al., 1998; Hanningtonet al., 1998). Paragonite was never obtained in themodel. Its development at the TAG field is probablyrelated to the extensive process of subsurface mixing,which was not taken into account in the model. Thezones of paragonite development were not documentedin ancient massive sulfide occurrences.
The calculation of the model with magmatic fluidsyielded pyrophyllite-bearing metasomatic assem-blages. In natural deposits such assemblages are indic-ative of the conditions of intense acid leaching (Meta-somatism of…, 1998). Pyrophyllite occurs sporadicallyas a minor component in ancient massive sulfide depos-its, and its appearance is often connected with the pres-ence of silicic rocks.
Monomineralic pyrite bodies deposited from acidsolutions have not been found on the ocean floor. Suchobjects are known in ancient deposits (Franklin et al.,1981). The sequential deposition of pyrite and basemetal sulfides was suggested by many authors for mas-sive sulfide deposits (e.g., Smirnov, 1982), but thisproblem is yet not fully resolved because of the difficul-ties of reconstructing ancient ore-forming processes.
The results of simulation allow us to suppose thefollowing.
(1) The footwall zoning of massive sulfide depositscould only be formed by the superposition of seawaterconvection on the products of the magmatic stage ofsystem development. No single solution source (neithermagmatic nor exogenous) could produce such a zoningin submarine hydrothermal systems.
(2) During the stage of exhalation activity, the orebody has a pyrite composition. Base metals are accu-mulated within it by the superposition of a convectiverecycling process. On the other hand, the major portionof sulfides is introduced into massive sulfide depositsduring the magmatic stage.
These suggestions can be used as a starting point forfuture investigations of submarine hydrothermal sys-tems with a combined fluid source.
Fig. 67. Mass of a model ore body and quantitties of ore ele-ments. (a) Total mass, (b) amounts of iron and sulfur, and(c) amounts of base metals.
S306
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
CONCLUSIONS
The results obtained in Chapters 4 and 6 suggest thatthe ore-forming potential of hydrothermal systemsdepends on the concentration of ore elements in thesolutions formed within them. The question of the fac-tors controlling these concentrations is pivotal for theunderstanding of the geochemistry of oceanic ore min-eralization.
The metal content of hydrothermal systems dependson diverse factors which may be interdependent. It istherefore difficult to discriminate the contributions ofindividual processes by the empirical generalization ofnatural data. Let us consider how such a problem can beresolved by a combined analysis of results obtainedfrom the examination of natural objects, experimentaldata, and thermodynamic modeling.
The analysis of the logical scheme of oceanic hydro-thermal systems (Section 3.2) showed that the mainprocess responsible for the accumulation of ore ele-ments in hydrothermal solutions is solution interactionwith the country rocks. Hence, the conditions of water–rock interaction are the main factors that control themetal content in oceanic hydrothermal systems. Themost important parameters are temperature, pressure,and the intensity of interaction (specified by the R/Wratio). The results of simulation (Section 4.1.2) suggestthat the character of water–rock interaction changesduring the evolution of a hydrothermal system, and,consequently, the lifetime of the system (τ) is also avery important factor. The occurrence of boiling withinhydrothermal systems also influences the behavior ofmetals in solutions (Section 6.2), and the fraction ofreleased gas (η) is a factor of metal concentration insuch systems.
In addition to the aforementioned factors, a numberof other parameters that could affect the behavior ofmetals in oceanic hydrothermal systems have beendiscussed in the literature. The possible role of mag-matic gases in the supply of metals was addressed inSection 3.2, where we showed that this contribution isnegligible in the hydrothermal systems of mid-oceanridges. The composition of hydrothermal solutions canbe affected by the composition of oceanic crustal mate-rials. In the majority of oceanic hydrothermal systemsstudied, water reacts with rather homogeneoussequences of tholeiitic basalts, which show little chem-ical variation.38
Each of the three sources of information (naturaldata, experiments, and computer simulation) relevant tothe problem considered has certain advantages and dis-advantages.
Observations of natural objects provide the mostvaluable evidence. A considerable body of data has
38 The role of hydrothermal solution interaction with sedimentaryrocks overlying basaltic sequences and the participation of ultra-basic rocks in the generation of metal concentrations in hydro-thermal systems should be addressed in future studies.
been accumulated to date on the chemical compositionof hydrothermal solutions from submarine systems.The interpretation of these data must allow for the factthat only solutions discharged on the ocean floor areavailable for analysis, whereas the metals are extractedfrom rocks into the solution in deep parts of the system.The solution may lose some of its ore components onthe way to the surface. The contributions of these fac-tors cannot be directly measured, because they operatein the interior part of the system. They have to be recon-structed, and while such reconstructions are reliable forT and R/W, other factors require some assumptions.
The experimental simulation of seawater–basaltinteraction allows characterization of the process undercontrolled conditions (T, P, and η). This approach pro-vides unbiased data independent of the knowledge ofthe interaction mechanism. The experimental equip-ment that was used in the majority of recent studies(autoclaves allowing liquid sampling during the exper-iment) minimizes systematic errors. However, the inter-pretation of experimental data must account for the fol-lowing points: (a) the time span of the natural process(τ) cannot be reproduced in experiments; (b) the avail-able equipment is designed for static experiments,whereas the natural hydrothermal systems are of a flow-through character; and (c) the experimental results donot always correspond to the target R/W values, and thesynthetic assemblages of secondary minerals are notfully consistent with natural ones (Section 3.2).
The thermodynamic modeling of hydrothermal sys-tems provides unique opportunities for solving theproblem, because all the factors in question are pre-scribed as model parameters. It allows the evaluation ofthe role of each factor separately, while keeping allother parameters constant, which is practically impos-sible by other methods. However, thermodynamic mod-eling is most sensitive to the completeness and qualityof our knowledge on the geochemical processes occur-ring in hydrothermal systems. Nonetheless, a compari-son of the results obtained by computer simulation withthe characteristics of natural objects showed adequateagreement with respect to a wide range of parameters(Section 4.4).
From the above considerations, it is evident that thepossible shortcomings and difficulties of interpretationof various sources of information are mutually indepen-dent and uncorrelated. An important consequence ofthis result is that the robustness of a conclusionincreases substantially if it is drawn by fundamentallydifferent methods.
Influence of temperature on the concentration ofmetals in oceanic hydrothermal systems. According tonatural observations, the highest concentrations of oreelements are characteristic of the hottest black smokers(Table 6), whose measured discharge temperature isalmost always higher than 300°C. Colder white andtransparent smokers with temperatures below 300°Cshow much lower concentrations of ore metals. How-
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S307
ever, this relationship does not necessarily reflect thecharacter of processes in the interior part of hydrother-mal systems but may be related to the discharge condi-tions on the ocean floor. White and black smokers oftenoccur in close proximity and probably have a commondeep source. Strictly speaking, the natural data are onlyindicative of the decreasing transportation capacity ofhydrothermal solutions during their cooling. However,since the transportation of hydrogen sulfide togetherwith heavy metals in a hydrothermal solution is possi-ble only at high temperatures, extensive ore generationcan be achieved only in high-temperature systems(probably, higher than 300°C in the interior parts).
The experimental investigation of seawater–basaltinteraction provides compelling evidence that anincrease in T intensifies the mobilization of ore metalsfrom rocks. Figure 68a shows the experimental resultsby Seyfried and Janecky (1985), which were obtainedholding all other interaction factors constant (P = 400 barand R/W = 1). The concentrations of ore metals in theexperimental fluids appeared to be similar to those ofnatural hydrothermal solutions at T = 375–400°C.
The results of simulation (Fig. 68b) support thisinference. At high R/W, while sulfides are present in thesolid phase, the equilibrium concentrations of heavymetals in the model solutions increase with increasingtemperature.
There is no such effect at low R/W, because basemetals are completely extracted from the basalts andtheir concentrations are functions of R/W only (how-ever, the low-R/W range probably does not play a sig-nificant role in the formation of ore-forming solutionsin the natural systems considered).
The explanation of the temperature dependency ofore metal concentrations is straightforward. At hightemperatures the degree of base metal complexation(primarily with the chloride ion) rises and this enhancesthe solubility of sulfides and the mobility of heavy met-als.
Influence of pressure. Natural observations are notsufficient to reveal the influence of this factor. Most ofthe hydrothermal systems studied are situated at theaxis of the EPR, where the ocean is about 2.5 km deep,and their magma chambers probably occur at approxi-mately equal depths. This parameter is different for thesystems of the Mid-Atlantic Ridge (about 3.6 km deep)and Axial Seamount of the Juan de Fuca Ridge (about1.5 km deep), but these systems are distinguished fromthose of the EPR by other important characteristics. Inparticular, the hydrothermal system of Axial Seamountshows evidence of boiling.
Experimental data indicate that, in general, pressurehas a relatively weak effect on water–basalt interaction,except for the P–T region near the critical point of water(T = 370–400°C and P < 500 bar). According to theexperiments by Seyfried and Janecky (1985), a pressuredecrease within this region provides extensive mobili-zation of heavy metals. Figure 69a shows experimental
results (Mottl et al., 1979; Rosenbauer and Bischoff,1983; Seyfried and Janecky, 1985) obtained at variouspressures holding other experimental parameters con-stant (T = 400°C and R/W =1).
The results of thermodynamic modeling (Fig. 69b)reveal the same tendency for Fe and Zn.39
This correlation is related to a dramatic change inthe physicochemical properties of water (primarily, itsdielectric constant) in the near-critical region. Thedecrease of the dielectric constant toward Pcr stabilizescomplex compounds in the solution (Ryzhenko, 1981).As a result, the solubility of heavy metal sulfidesincreases strongly. Far from the critical point, the influ-ence of P on the properties of water is negligible.
Influence of the rock/water ratio. The R/W values ofthe hydrothermal systems studied vary over a relatively
39 The pressure dependency of the thermodynamic properties ofCu and Pb compounds is not well constrained at high tempera-tures.
150 200–10
–2
250 320 350 400 450
0
–4
–6
–8
Temperature, °C
log(Me) [mol/kg]
(b)
Fe
S
Cu
Zn
Pb
340 360–6
–5
380 400 420
–4
–3
–2
Cu
Zn
Mn
Fe
(a)
Fig. 68. Influence of temperature on the concentration ofore elements in solutions from the interior part of a hydro-thermal system. (a) The experiments of Seyfried and Jan-ecky (1985) at P = 400 bar and R/W = 1 and (b) the resultsof calculations at P = 500 bar and R/W = 0.3.
log(Me) [mol/kg]
S308
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
narrow range from 0.5 to 2.0 (Tables 16, 17). The R/Westimates for individual systems differ, but the uncer-tainties of these estimates are comparable with the vari-ations between the systems. The influence of this factor
on the metal content of hydrothermal systems cannotyet be estimated from natural observations. Accordingto Bowers et al. (1988), the compositions of solutionsin the systems studied are independent of R/W, becausethey are buffered by the mineral associations of thecountry rocks in high-temperature zones (rock-domi-nated systems).
In an early experimental study, Seyfried and Mottl(1982) demonstrated that the concentrations of ore met-als depend on the regime of water–basalt interaction.Low R/W values in water-dominated systems result in ahigher metal content compared with those generated inrock-dominated systems with high R/W. According totheir results, the boundary between the two regimes liesat an R/W value between 0.02 and 0.1. Most of the sub-sequent experimental studies were performed at higherR/W values (>0.1), which are closer to the conditions inactive hydrothermal systems. The studies that includeda series of kinetic experiments showed that the concen-trations of heavy metals were high at the beginning ofexperiments (when the effective R/W was small) anddecreased toward the steady-state stage. However, it isstill unclear what the role is of sorption and incongruentdissolution of solid phases during the initial stage, andwhat can these results tell us about the conditions ofnatural processes?
The thermodynamic modeling supported the con-clusions of Seyfried and Mottl (1982). Figure 70 showsthe calculated dependency of the concentrations of oremetals on the cumulative R/W ratio during water–basaltinteraction. In the model with low R/W, the metals arecompletely mobilized owing to oxidation processes.Their concentrations in the solution show a positivecorrelation with the amount of rock participating in thereaction. When the system reaches a certain R/W value,it becomes reduced, and H2S appears in the solutionand precipitates heavy metal sulfides. This results in anegative correlation between metal concentrations andR/W at high R/W values. Thus, the concentration ofheavy metals passes through a maximum at R/W =0.03–0.1. The concentration of iron in the solution alsoreaches a maximum value when assemblage I (hematite +chlorite + …) is changed to assemblage II (epidote +Fe-chlorite + actinolite + …). In contrast, the bulk sul-fur content shows a minimum corresponding to thechange of its migration forms from SVI at low ΣR/W toSII at high ΣR/W. The results of our simulation suggestthat each chalcophile element shows a change in thecharacter of R/W dependency when its own sulfideappears in the phase association. Thus, the concentra-tions of these elements in solution are indirectly relatedto the type of the system (fluid-dominated or rock-dom-inated), through H2S accumulation in the solution.
Influence of the lifetime of a hydrothermal system.As was noted in Section 3.1, the investigation of oce-anic hydrothermal ores (Krasnov, 1990; HydrothermalSulfide…, 1992) showed that small, short-lived sulfideedifices are usually significantly enriched in Zn,
200 4000
2
100 2000
2
(a)
600 800 1000 1200
4
6
8
10
12Me, mmol/kg
Fe
Mn
Zn × 50
300 400 500 600
4
6
8
Pressure, bar
FeZn × 50
(b)
Fig. 69. Effect of pressure on the concentration of ore ele-ments in hydrothermal solutions. (a) The experiments ofMottl et al. (1979), Rosenbauer and Bischoff (1983), andSeyfried and Janecky (1985) at T = 400°C and R/W = 1; and(b) the results of calculations at T = 400°C and R/W = 0.6.
–3 –2–8
–7
–1 0 1
–6
–5
–4
–3
–2
–1
log(R/W)
log(Me)
Bn SphGn
Po
Fe
S
Cu
Zn
Pb
Fig. 70. Concentrations of ore elements as functions of thebulk rock/water ratio. Thermodynamic calculations at T =350°C and P = 500 bar. The dashed line shows the changefrom metasomatic assemblage I (oxidized, Mg-Chl + Hem +Anh + Qtz) to assemblage II (reduced, Ep + Fe-Chl + Act +Ab +…), and the arrows mark the appearance of ore metalsulfides in the assemblage.
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S309
whereas large long-lived bodies comprising megatonsof sulfides are distinguished by high copper concentra-tions. However, the available natural observations arenot sufficient to reliably estimate the influence of thelifetime of a hydrothermal system on the compositionof its solutions. For instance, the major-element com-position of solutions from a long-lived hydrothermalsystem of the TAG field is similar to that of solutionsfrom the EPR systems (Table 6).
There is no experimental evidence for the influenceof the lifetime of systems on the composition of theirsolutions.
Our thermodynamic modeling showed that heavymetals are fractionated during prolonged convectivecirculation owing to metasomatic phenomena in theinterior parts of hydrothermal systems (Section 4.1.2).The more mobile Zn and Pb are intensely removed bythe first hydrothermal solutions (Fig. 27), whereas Cu isretained in the rock in relatively less soluble sulfides(chalcopyrite and bornite). In the course of prolongedsystem evolution, the mobility of Cu begins to increasein response to SII removal from the rock, and the lateportions of hydrothermal solutions are copper-rich.This model result is in agreement with the aforemen-tioned correlation between the geochemical signatureof sulfide bodies and the lifetime of hydrothermal sys-tems. It is also supported by the relative Cu enrichmentin the solutions from the long-lived TAG system(Edmond et al., 1995).
Influence of boiling. As was shown in Section 6.2,the available natural data indicate that boiling withinhydrothermal systems increases the concentrations ofore metals in hydrothermal solutions (Fig. 52). Low-salinity hydrothermal vents are formed simultaneously.They are enriched in hydrogen sulfide and other vola-tiles and depleted in metals. These phenomena arerelated to phase separation in the interior parts of thesystem. The metal-rich smokers discharge residualbrine, and the freshened solutions contain an admixtureof the condensed vapor phase.
Boiling in contact with a rock was experimentallystudied by Bischoff and Rosenbauer (1987) under P–Tconditions similar to those of oceanic hydrothermalsystems (Table 25). The results of their experimentsshowed that boiling promotes extensive mobilization ofchalcophile elements from rocks.
Our thermodynamic modeling of boiling hydrother-mal solutions in the focus of a hydrothermal system isin agreement with these experiments. As can be seenfrom the results of the simulation of isothermal boilingin contact with a rock (Section 6.4.1), an increase in thefraction of the vapor phase is accompanied by sulfidedissolution in the rock and metal accumulation in thesolution. The analysis of our model results suggests thatthe main reason for the mobilization of chalcophile ele-ments is H2S partitioning from solution into the vaporphase and the corresponding shift of sulfide equilibriato dissolution.
An important feature of the thermodynamic modelincluding boiling is the change of the role of the R/Wratio. If there is no boiling, the equilibrium concentra-tions of heavy metals decrease in the solution withincreasing R/W (at R/W > 0.1) owing to the accumula-tion of H2S in it (Fig. 70). This effect is completelyeliminated by boiling owing to H2S partitioning into thevapor phase. The larger the amount of rock interactingwith heterogeneous fluid (vapor + water), the higher thedegree of ore metal mobilization. Therefore, the metalcontent of hydrothermal solutions from boiling systemsmust increase with increasing R/W value.
Thus, the comparative analysis of natural and exper-imental data and computer simulation allowed us toevaluate the influence of T, P, R/W, τ, and η on the metalconcentrations in oceanic hydrothermal systems. Theresults of this analysis are summarized in Table 34.
As can be seen in this table, all the aforementionedsources of information give identical results. Since allthe methods are mutually independent and the possibleuncertainties of these methods are uncorrelated, suchan agreement supports the reliability of the results. Thisallows us to state the following.
(1) An increase in temperature within the system isfavorable for the accumulation of ore metals in thehydrothermal solutions.
(2) If near critical conditions are maintained in theinterior parts, the metal content is higher in lower pres-sure systems (and other parameters are equal), i.e.,when the roof of the magma chamber is situated at shal-lower levels.
(3) The maximum removal of metals from basalts isattained at R/W ≈ 0.03. An increase in R/W above0.1 reduces the mobility of ore metals.
Table 34. Influence of various factors on the metal concen-trations in oceanic hydrothermal systems
FactorNatural observa-
tionsExperiments
Thermo-dynamicmodeling
Temperature + + +
Pressure (outsidethe boiling zone)
i.i. – –
Boiling Fe, Zn+ Cu, Zn+ +
Rock/water ratio
—outside theboiling zone
i.i. maximum between 0.02 and 0.1
maximumbetween 0.025 and 0.1
—during boiling i.i. i.i. +
Lifetime of the hy-drothermal system
(+) i.i. Cu+Zn, Pb, S–
Note: The plus and minus signs denote an increase and a decreasein metal content in response to the increase of the factor,respectively; i.i. means insufficient information; and valuesin parentheses show estimates.
S310
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
(4) Boiling in the interior parts of hydrothermal sys-tems enhances the removal of heavy metals.
(5) An increase in the duration of hydrothermalactivity must result in a relative copper enrichment inthe solutions, which is reflected in the geochemical sig-nature of the ore bodies.
It should be noted that the method of thermody-namic modeling appeared to be the most informativeamong the approaches used to distinguish the factors ofmetal accumulation in hydrothermal systems(Table 34).
The following general conclusions were drawnfrom the results of this study.
(1) A complex method for thermodynamic modelingwas developed in this study on the basis of the principleof a step flow reactor. This method allows for thedescription of all major geochemical processes in sub-marine hydrothermal systems. A comprehensivegeochemical model was constructed for an ore-formingoceanic system from the formation of hydrothermalsolution in the downwelling limb of the convection cellto ore deposition during solution discharge on theocean floor. The results of computer simulation are con-sistent with the natural prototypes in a variety of char-acteristics, including metasomatic and ore mineralassemblages and the chemical and isotopic composi-tions of solutions. This enables the use of our model forthe prediction of the behavior of elements in naturalsystems.
(2) The alteration of seawater in the downwellinglimb of the convection cell occurs when water passesthrough the regions of two mineral assemblages: (I) oxi-dized (chlorite + hematite + anhydrite + quartz + …) and(II) reduced (chlorite + epidote + actinolite + albite +sulfides + …), which show fundamentally differentbehaviors of major and trace elements. The position ofthe boundary between these regions depends on theinput–output budget of the components. A prolongedinflux of seawater into the system changes its bulkchemical composition and shifts the oxidation front tohigher temperatures, which results in the time evolutionof hydrothermal solutions produced by the system.
(3) The ore metals are extensively extracted fromthe region of assemblage I and are weakly mobile in theregion of assemblage II because of sulfide precipita-tion. The sequence of element mobility in oceanichydrothermal environments is SII > Pb > Zn > Cu > Fe.The metallogenic signature of solutions changes duringtheir evolution from Zn > Cu to Cu > Zn. The reason forthis inversion is the different rates of Zn and Curemoval from the interior parts of the system. Thisexplains the observed discrimination of sulfide edificeson the ocean floor into zinc- and copper-dominatedones and the confinement of the latter to long-lived sys-tems in slow-spreading ridges and volcanic seamounts.
(4) The physicochemical mechanism of formationof large sulfide bodies includes (a) the evolution of par-ent solutions, (b) ore deposition during mixing with
seawater, and (c) the metasomatic redeposition of mat-ter within the edifice. Such a scenario produces a zonedore body, whose structure and metallogenic signaturechange with time. The early stages of this evolutioncorrespond to the known active vents on the oceanfloor. The mature stages of evolution yield massive sul-fide ore bodies, whose structures and compositions aresimilar to those of ancient ore occurrences on conti-nents.
(5) A thermodynamic model was constructed on thebasis of the method of an isotopic chemical system forsulfur isotopes. It showed that the sulfur isotopic sys-tematics of convective hydrothermal systems in theocean are controlled by the mixing of sulfur from twoisotopically different sources, seawater and basalt. Themain characteristic is the regular increase in the δ34Svalue of a solution over the lifetime of the system. Thisexplains the observed variations in the isotopic compo-sitions of oceanic sulfides, including the different isoto-pic signatures of small and large ore bodies and the 34Senrichment of dissolved hydrogen sulfide comparedwith sulfides from the ore body.
(6) Boiling within oceanic hydrothermal systems isfavorable for the concentration of metals in hydrother-mal solutions. Phase separation in solution results inthe partitioning of metals and hydrogen sulfide into dif-ferent phases and an increase in the solubility of sul-fides in heterogeneous systems (mechanism of precipi-tant loss). This effect must evidently be manifested forany elements with a sulfide form of precipitation andnonsulfide forms of solution transfer. Hence, boilingincreases the ore-generating potentials of hydrothermalsystems. The cooling and condensation of heteroge-neous fluids during their submarine discharge can pro-duce considerable amounts of sulfide deposits. Theydiffer from the ores formed by cooling of homogeneousfluids in having a higher fraction of heavy metals.
In accordance with the classification of simulationproblems proposed in Section 1.1, conclusion (1) cor-responds to type II problems (internal with respect tothe method), and conclusions (2)–(6) are related to typeII problems (external with respect to the method). Con-clusions (3)–(5) provide solutions to the problems offunctional relationships between the known propertiesof the object studied, and conclusions (2) and (6) pre-dict the properties of natural objects. We believe thatfurther progress in the thermodynamic modeling of nat-ural geochemical processes will result in a higher frac-tion of predictive results, which are most valuable forthe development of geochemistry.
REFERENCES
1. E. E. Abramova and D. V. Grichuk, “ComputationalThermodynamic Model of the Recycled HydrothermalSystem,” Rudy Metally, No. 2, 36–44 (1994).
2. V. A. Akimtsev, V. N. Sharapov, V. Yu. Kolobov, et al.,“Hydrothermal Mineralization of Mid-Atlantic Ridge
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S311
in the Zone of Its Junction with the Green Cape Trans-form Fault,” Geol. Geofiz., No. 8, 32–37 (1991).
3. N. N. Akinfiev, “A Method for the Computation of OreDeposition from Boiling Fluid: Accounting for Dielec-tric Constant,” Geokhimiya, No. 10, 1465–1478 (1994).
4. F. Albarede, A. Michard, J. F. Minster, and G. Michard,“87Sr/86Sr Ratios in Hydrothermal Waters and Depositsfrom the East Pacific Rise at 21° N,” Earth Planet. Sci.Lett. 55, 229–236 (1981).
5. V. A. Alexeyev, “Dynamic Model of Infiltration Meta-somatism Based on the Calculation of Local Equilib-ria,” Geokhimiya, No. 9, 1311–1320 (1985).
6. A. I. Almukhamedov, G. L. Kashintsev, andV. V. Matveenkov, Evolution of Basaltic Volcanism inthe Red Sea Region (Nauka, Novosibirsk, 1985) [inRussian].
7. J. C. Alt, “The Chemistry and Sulfur Isotopic Composi-tion of Massive Sulfide and Associated Deposits onGreen Seamount, Eastern Pacific,” Econ. Geol. 83,1026–1033 (1988).
8. J. C. Alt and W. C. Shanks, III, “Serpentinization ofAbyssal Peridotites from the MARK Area, Mid-Atlan-tic Ridge: Sulfur Geochemistry and Reaction Modeling,”Geochim. Cosmochim. Acta 67, 641–653 (2003).
9. J. C. Alt, J. Honnorez, H.-W. Hubberten, and E. Saltz-man, “Occurrence and Origin of Anhydrite from DeepSea Drilling Project Leg 70, Hole 504B, Costa RicaRift,” Deep Sea Drill. Project: Initial Rep. 69, 547–550(1983).
10. J. C. Alt, P. Lonsdale, R. Haymon, and K. Muechlenb-achs, “Hydrothermal Sulfide and Oxide Deposits onSeamounts near 21° N, East Pacific Rise,” Bull. Geol.Soc. Am. 98, 157–168 (1987).
11. J. C. Alt, T. E. Anderson, and L. Bonell, “The Geochem-istry of Sulfur in a 1.3 km Section of HydrothermallyAltered Oceanic Crust, DSDP Hole 504B,” Geochim.Cosmochim. Acta 53, 1011–1023 (1989).
12. J. Alvarez, R. Crovetto, and R. Fernandez-Prini, “TheDissolution of N2 and H2 in Water from Room Temper-ature to 640 K,” Ber. Bunsen-Ges. Phys. Chem. 92,935–940 (1988).
13. R. N. Anderson, J. C. Alt, J. Malpas, et al., “Geochem-ical Well Logging in Basalts: The Palisades Sill and theOceanic Crust of Hole 504B,” J. Geophys. Res. 95,9265–9292 (1990).
14. M. Arnold and S. M. F. Sheppard, “East Pacific Rise atLatitude 21° N: Isotopic Composition and Origin of theHydrothermal Sulfur,” Earth Planet. Sci. Lett. 56, 148–156 (1981).
15. Yu. A. Averkin, “Dynamics of Sulfide Deposition fromHydrothermal Solution in the Process of CO2 Boilingoff,” Geokhimiya, No. 11, 1580–1586 (1987).
16. R. D. Ballard and J. Francheteau, “The Relationshipbetween Active Sulfide Deposition and the Axial Pro-cesses of the Mid-Ocean Ridge,” Mar. Technol. Soc.Japan 16 (3), 8–22 (1982).
17. L. A. Bannikova, Organic Matter in Hydrothermal OreFormation (Nauka, Moscow, 1990) [in Russian].
18. L. A. Bannikova and B. N. Ryzhenko, “Isotopic Ratiosof Carbon and Sulfur in the Products of Redox Reac-tions inunder Hydrothermal Conditions: The System
19. L. A. Bannikova, D. V. Grichuk, and B. N. Ryzhenko,“Calculation of Chemical and Isotopic Equilibria in theC–H–O System and Their Application to a Study ofRedox Reactions in Hydrothermal Systems,”Geokhimiya, No. 3, 416–428 (1987).
20. Vikt. L. Barsukov and M. V. Borisov, “Simulation ofGeochemical Consequences of Selfmixing Hydrother-mal Solutions: 1. Mass Transfer at the Spreading Areasof Flows of Hydrothermal Solutions,” Geokhimiya,No. 8, 1108–1123 (1982a).
21. Vikt. L. Barsukov and M. V. Borisov, “Simulation ofGeochemical Consequences of Selfmixing Hydrother-mal Solutions: 2. Mass Transfer at the Areas of Con-striction of Flows of Hydrothermal Solutions,”Geochem. Int. 19 (5), 26–36 (1982b).
22. R. Becker, M. G. Langseth, R. P. Von Herzen, andR. N. Anderson, “Deep Crustal Geothermal Measure-ments, Hole 504B, Costa Rica Rift,” J. Geophys. Res.88, 3447–3457 (1983).
23. V. M. Belyi, A. A. Migdisov, N. V. Barskaya, andV. A. Grinenko, “Redistribution of Sulfur and Its Iso-topes in Hydrothermally Altered Oceanic Basalts, CostaRica Rift, Hole 504B,” Geokhimiya, No 3, 390–402(1984).
24. M. E. Berndt and W. E. Seyfried, Jr., “Boron, Bromine,and Other Trace Elements as Clues to the Fate of Chlo-rine in Mid-Ocean Ridge Vent Fluids,” Geochim. Cos-mochim. Acta 54, 2235–2245 (1990).
25. M. E. Berndt and W. E. Seyfried, Jr., “Calibration ofBr/Cl Fractionation during Subcritical Phase Separationof Seawater: Possible Halite at 9 to 10° N East PacificRise,” Geochim. Cosmochim. Acta 61, 2849–2854(1997).
26. D. Bideau, Y. Fouquet, and R. Hekinian, “Decouvertede basaltes metamorphises dans le graben axial de ladorsale Est-Pacifique a 12°50′ N,” Soc. Geol. Fr. Bull.1, 905–913 (1985).
27. R. A. Binns and S. D. Scott, “Actively Forming Polyme-tallic Sulfide Deposits Associated with Felsic VolcanicRocks in the Eastern Manus Back-Arc Basin, PapuaNew Guinea,” Econ. Geol. 88, 2226–2236 (1993).
28. D. K. Bird and H. C. Helgeson, “Chemical Interactionof Aqueous Solutions with Epidote–Feldspar MineralAssemblages in Geological Systems: I. Thermody-namic Analysis of Phase Relations in the System CaO–FeO–Fe2O3–Al2O3–SiO2–H2O–CO2,” Am. J. Sci. 280,907–910 (1980).
29. J. L. Bischoff and F. W. Dickson, “Seawater–BasaltInteraction at 200°C and 500 bars: Implication for Ori-gin of Sea-Floor Heavy-Metal Deposits and Regulationof Seawater Chemistry,” Earth Planet. Sci. Lett. 25,385–397 (1975).
30. J. L. Bischoff and K. S. Pitzer, “Phase Relations andAdiabats in Boiling Seafloor Geothermal Systems”Earth Planet. Sci. Lett. 75, 327–338 (1985).
31. J. L. Bischoff and K. S. Pitzer, “Liquid–Vapor Relationsfor the System NaCl–H2O: Summary of the P–T–x Sur-face from 300° to 500°C,” Am. J. Sci. 289, 217–248(1989).
S312
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
32. J. L. Bischoff and R. J. Rosenbauer, “Phase Separationin Seafloor Geothermal Systems: An ExperimentalStudy of the Effects on Metal Transport,” Am. J. Sci.287, 953–978 (1987).
33. J. L. Bischoff and R. J. Rosenbauer, “Liquid–VaporRelation in the Critical Region of the System NaCl–H2O from 380° to 415°C: A Refined Determination ofthe Critical Point and Two-Phase Boundary of Seawa-ter,” Geochim. Cosmochim. Acta 52, 2121–2126(1988).
34. J. L. Bischoff and R. J. Rosenbauer, “Phase Separationin Seafloor Geothermal Systems by Layered Double-Diffusive Convection,” J. Geol. 97, 613–623 (1989).
35. J. L. Bischoff and W. E. Seyfried, Jr., HydrothermalChemistry of Seawater from 25° to 350°C”, Am. J. Sci.278, 838–860 (1978).
36. J. L. Bischoff, A. S. Radtke, and R. J. Rosenbauer,“Hydrothermal Alteration of Graywacke by Brine andSeawater: Roles of Alteration and Chloride Complexingon Metal Solubilization at 200° and 350°C,” Econ.Geol. 76, 659–676 (1981).
37. J. L. Bischoff, R. J. Rosenbauer, and K. S. Pitzer, “Thesystem NaCl–H2O: Relation of Vapor–Liquid near theCritical Temperature of Water and of Vapor–Liquid–Halite from 300° to 500°,” Geochim. Cosmochim. Acta50, 1437–1444 (1986).
38. J. L. Bischoff, R. J. Rosenbauer, and R. O. Fournier,“The Generation of HCl in the System CaCl2–H2O:Vapor–Liquid Relations from 380° to 500°C,” Geochim.Cosmochim. Acta 60, 7–16 (1996).
39. C. W. Blownt and F. W. Dickson, “The Solubility ofAnhydrite (CaSO4) in NaCl–H2O from 100 to 450°Cand 1 to 1000 bars,” Geochim. Cosmochim. Acta 33,227–245 (1969).
40. G. J. Bluth, and H. Ohmoto, “Sulfide–Sulfate Chimneyson the East Pacific Rise, 11° and 13° N Latitudes:II. Sulfur Isotopes,” Can. Mineral. 26, 505–515 (1988).
41. Yu. A. Bogdanov, Hydrothermal Deposits of the Mid-Atlantic Ridge Rifts (Nauchnyi Mir, Moscow, 1997) [inRussian].
42. Yu. A. Bogdanov, N. S. Bortnikov, I. V. Vikent’ev, et al.,“A New Type of Modern Mineral-Forming System:Black Smokers of the Hydrothermal Field at 14°45′ NLatitude, Mid-Atlantic Ridge,” Geol. Rudn. Mestor-ozhd. 39, 58–78 (1997a) [Geol. Ore Depos. 39, 68–90(1997a)].
43. Yu. A. Bogdanov, N. S. Bortnikov, and A. P. Lisitsin,“The Origin of Hydrothermal Sulfide Ores in the AxialParts of the Mid-Atlantic Ridge,” Geol. Rudn. Mestor-ozhd. 39, 351–370 (1997b) [Geol. Ore Depos. 39, 409–429 (1997b)].
44. M. V. Borisov, “Geochemical and ThermodynamicModels for the Genesis of Low and Medium Tempera-ture Vein Mineralization and Metasomatism in WallRocks,” Geochem. Int. 41 (Suppl. 2), S145–S312(2003).
45. M. V. Borisov and I. L. Khodakovskii, “On the Problemof Error Identification in Physicochemical Modeling,”Geokhimiya, No. 6, 907–908 (1989).
46. M. V. Borisov and Yu. V. Shvarov, Thermodynamics ofGeochemical Processes (Mosk. Gos. Univ., Moscow,1992) [in Russian].
47. N. S. Bortnikov, D. T. Fedorov, and K. G. Muraviov,“Mineral Composition and Formation Conditions ofSulfide Deposits in the Lau Basin, SouthwesternPacific,” Geol. Rudn. Mestorozhd. 35, 528–544 (1993).
48. Y. Bottinga and M. Javoy, “MORB Degassing: Evolu-tion of CO2,” Earth Planet. Sci. Lett. 95, 215–225(1989).
49. T. S. Bowers, “Stable Isotope Signatures of Water–Rock Interaction in Mid-Ocean Ridge HydrothermalSystems: Sulfur, Oxygen, and Hydrogen,” J. Geophys.Res. 94, 5775–5786 (1989).
50. T. S. Bowers, “The Deposition of Gold and Other Met-als: Pressure-Induced Fluid Immiscibility and Associ-ated Stable Isotope Signatures,” Geochim. Cosmochim.Acta 55, 2417–2434 (1991).
51. T. S. Bowers and H. P. Taylor, “An Integrated Chemicaland Stable-Isotope Model of the Origin of Mid-OceanRidge Hot Spring Systems,” J. Geophys. Res. 90,12 583–12 606 (1985).
52. T. S. Bowers, K. L. Von Damm, and J. M. Edmond,“Chemical Evolution of Mid-Ocean Ridge HotSprings,” Geochim. Cosmochim. Acta 49, 2239–2252(1985).
53. T. S. Bowers, A. C. Campbell, C. I. Measures, et al.,“Chemical Controls on the Composition of Vent Fluidsat 13°–11° N and 21° N, East Pacific Rise,” J. Geophys.Res. 93, 4522–4536 (1988).
54. T. Brikovski and D. Norton, “Influence of MagmaChamber Geometry on Hydrothermal Activity at Mid-Ocean Ridges,” Earth Planet. Sci. Lett. 93, 241–255(1989).
55. O. V. Bryzgalin, “Some Strong Electrolytes in theSupercritical Field: Estimation of Dissociation Con-stants Based of the Electrostatic Model,” Geokhimiya,No. 8, 1184–1195 (1985).
56. D. A. Butterfield and G. J. Massoth, “Geochemistry ofNorth Cleft Segment Vent Fluids: Temporal Changes inChlorinity and Their Possible Relation to Recent Volca-nism,” J. Geophys. Res. 99, 4951–4968 (1994).
57. D. A. Butterfield, G. J. Massoth, R. E. McDuff, et al.,“Geochemistry of Hydrothermal Fluids from Axial Sea-mount Hydrothermal Emission Study Vent Field, JuanDe Fuca Ridge: Subseafloor Boiling and SubsequentFluid–Rock Interaction,” J. Geophys. Res. 95, 12 895–12 921 (1990).
58. D. A. Butterfield, R. E. McDuff, M. J. Mottl, et al.,“Gradients in the Composition of Hydrothermal Fluidsfrom the Endeavour Segment Vent Field: Phase Separa-tion and Brine Loss,” J. Geophys. Res. 99, 9561–9583(1994).
59. G. Yu. Butuzova, “On the Sources of Substances in theHydrothermal–Sedimentary Ore Formation: I. Sourcesof Water, Gases, Sulfur, and the Generation of MajorSalt Composition of Ore-Forming Solutions,” Litol.Polezn. Iskop., No. 5, 3–18 (1986a).
60. G. Yu. Butuzova, “On the Sources of Substances in theHydrothermal–Sedimentary Ore Formation: II. Sourcesof Ore-Forming Elements,” Litol. Polezn. Iskop., No. 6,3–18 (1986b).
61. A. Yu. Bychkov, Candidate’s Dissertation in Geologyand Mineralogy (MGU, Moscow, 1995).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S313
62. A. Yu. Bychkov and D. V. Grichuk, “A ThermodynamicModel for the Sb–As Mineralization in the UzonCaldera,” Geokhimiya, No. 4, 527–538 (1991).
63. C. D. Byers, D. W. Muenow, and M. D. Garcia, “Vola-tiles in Basalts and Andesites from the GalapagosSpreading Center, 85 to 86° W,” Geochim. Cosmochim.Acta 47, 1551–1558 (1983).
64. C. D. Byers, D. M. Christie, D. W. Muenow, andJ. M. Sinton, “Volatile Contents and Ferric–FerrousRatios of Basalt, Ferrobasalt, Andesite, and RhyodaciteGlasses from the Galapagos 95.5° W Propagating Rift,”Geochim. Cosmochim. Acta 48, 2239–2245 (1984).
65. C. D. Byers, M. D. Garcia, and D. W. Muenow, “Vola-tiles in Basaltic Glasses from the East Pacific Rise at21° N: Implications for MORB Sources and SubmarineLava Flow Morphology,” Earth Planet. Sci. Lett. 79,9−20 (1986).
66. A. C. Campbell and J. M. Edmond, “Halide Systematicsof Submarine Hydrothermal Vents,” Nature 342, 168–170 (1989).
67. A. C. Campbell, T. S. Bowers, C. I. Measures, et al.,“A Time Series of Vent Fluid Compositions From 21° N,East Pacific Rise (1979, 1981, 1985), and the GuaymasBasin, Gulf of California (1982, 1985),” J. Geophys.Res. 93, 4537–4549 (1988a).
68. A. C. Campbell, M. R. Palmer, G. P. Klinkhammer,et al., “Chemistry of Hot Springs on the Mid-AtlanticRidge,” Nature 335, 514–519 (1988b).
69. J. R. Cann and M. R. Strens, “Black Smokers Fueled byFreezing Magma,” Nature 298, 147–149 (1982).
70. J. R. Cann, M. R. Strens, and A. Rise, “A SimpleMagma-Driven Thermal Model for Formation of Volca-nogenic Massive Sulfides,” Earth Planet. Sci. Lett. 76,123–134 (1985/1986).
71. J. J. Carroll and A. E. Mather, “Phase Equilibrium in theSystem Water–Hydrogen Sulfide: Modeling the PhaseBehavior with an Equation of State,” Can. J. Chem.Eng. 67, 999–1003 (1989).
72. L. M. Cathles, “An Analysis of the Cooling of Intrusivesby Groundwater Convection which Includes Boiling,”Econ. Geol. 72, 804–826 (1977).
73. J.-P. Charlou, Y. Fouquet, J. P. Donval, et al., “Mineraland Gas Chemistry of Hydrothermal Fluids on anUltrafast Spreading Ridge: East Pacific Rise, 17–19° S(Naudir Cruise, 1993) Phase Separation Processes Con-trolled by Volcanic and Tectonic Activity,” J. Geophys.Res. 101, 15 889–15 919 (1996).
74. J. L. Charlou, J. P. Donval, E. Douville, et al., “Com-pared Geochemical Signatures and the Evolution ofMenez Gwen (37°50′ N) and Lucky Strike (37°17′ N)Hydrothermal Fluids, South of the Azores Triple Junc-tion on the Mid-Atlantic Ridge,” Chem. Geol. 171, 49–75 (2000).
75. J. L. Charlou, J. P. Donval, Y. Fouquet, et al.,“Geochemistry of High H2 and CH4 Vent Fluids Issuingfrom Ultramafic Rocks at the Rainbow HydrothermalField (36°14′ N, MAR),” Chem. Geol. 191, 345–359(2002).
76. J. H. Chen, “U, Th and Pb Isotopes in Hot Springs onthe Juan de Fuca Ridge,” J. Geophys. Res. 92, 11 411–11 415 (1987).
77. J. H. Chen, G. J. Wasserburg, K. L. Von Damm, andJ. M. Edmond, “The U–Th–Pb Systematics in HotSprings on the East Pacific Rise at 21° N and GuaymasBasin,” Geochim. Cosmochim. Acta 50, 2467–2479(1986).
78. J. S. Cline, R. J. Bodnar, and J. D. Rimstidt, “NumericalSimulation of Fluid Flow and Silica Transport and Dep-osition in Boiling Hydrothermal Solution: Applicationto Epithermal Gold Deposits,” J. Geophys. Res. 97,9085–9103 (1992).
79. R. G. Coleman, Ophiolites (Springer-Verlag, Berlin,1977).
80. D. R. Converse, H. D. Holland, and J. M. Edmond,“Flow Rates in the Axial Hot Springs of the East PacificRise (21° N): Implication for the Heat Budget and theFormation of Massive Sulfide Deposits,” Earth Planet.Sci. Lett. 69, 159–175 (1984).
81. H. Craig, Y. Horibe, K. A. Farley, et al., “HydrothermalVents in the Mariana Trough: Results of the First AlvinDives,” EOS, Trans. Am. Geophys. Union 68, 1531(1987).
82. K. Crane, “The Distribution of Geothermal Fields alongthe Mid-Ocean Ridge: An Overview,” in HydrothermalVents in the East Pacific, Bull. Biol. Soc. Washington,No. 30, 3–18 (1985).
83. C. G. Cunningham, “Characteristics of Boiling-Water-Table and Carbon Dioxide Models for Epithermal GoldDeposition,” US Geol. Surv. Bull., No. 1646, 43–46(1985).
84. A. S. Davis and D. A. Clague, “Geochemistry, Mineral-ogy, and Petrogenesis of Basalt from the Gorda Ridge,”J. Geophys. Res. 92, 10 467–10 483 (1987).
85. E. E. Davis, W. D. Goodfellow, B. D. Bornhold, et al.,“Massive Sulfides in a Sediment Rift Valley, NorthernJuan de Fuca Ridge,” Earth Planet. Sci. Lett. 82, 49–61(1987).
86. E. E. Davis, M. J. Mottl, A. T. Fisher, et al., Proceedingsof the Ocean Drilling Program. Initial Reports (1992),Vol. 139.
87. J. R. Delaney and B. A. Cosens, “Boiling and MetalDeposition in Submarine Hydrothermal Systems,”J. Mar. Technol. Soc. 16 (3), 62–66 (1982).
88. J. R. Delaney, R. E. McDuff, M. K. Tivey, and J. E. Lup-ton, “Measurements of 400°C Hydrothermal Fluids andTemporal Variability in the Endeavour Vent Field,”EOS, Trans. Am. Geophys. Union 70, 1163 (1989).
89. C. G. Denbig, Theory of Chemical Reactors (Nauka,Moscow, 1968) [in Russian].
90. R. S. Detrick, J. Honnorez, et al., “Mid-Atlantic Bare-Rock Drilling and Hydrothermal Vents,” Nature 321,14–15 (1986).
91. F. W. Dickson, “The Role of Rhyolite–Seawater Reac-tion in the Genesis of Kuroko Ore Deposits,” in Pro-ceedings of the 2nd International Symposium on Water–Rock Interaction, Sec. 4 (Strasbourg, 1977), pp. 181–190.
92. B. R. Doe, “Zinc, Copper, and Lead in Mid-OceanRidge Basalts and the Source Rock Control on Zn/Pb inOcean-Ridge Hydrothermal Deposits,” Geochim. Cos-mochim. Acta 58, 2215–2223 (1994).
S314
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
93. P. I. Dorogokupetz and I. K. Karpov, Thermodynamicsof Minerals and Mineral Equilibria (Nauka, Novosi-birsk, 1984) [in Russian].
94. E. Douville, J. L. Charlou, E. H. Oelkers, et al., “TheRainbow Fluids (36°14′ N, MAR): The Influence ofUltramafic Rocks and Phase Separation on Trace MetalContent in Mid-Atlantic Ridge Hydrothermal Fluids,”Chem. Geol. 184, 37–48 (2002).
95. S. E. Drummond and H. Ohmoto, “Chemical Evolutionand Mineral Deposition in Boiling Hydrothermal Sys-tems,” Econ. Geol. 80, 126–147 (1985).
96. F. Duennebier and G. Blackington, “A Man-Made HotSpring on the Ocean Floor,” Deep Sea Drill. Project:Initial Rep. 65, 357–359 (1983).
97. J. Dymond, R. Cobler, L. Gordon, et al., “226Ra and222Rn Contents of Galapagos Rift HydrothermalWaters: The Importance of Low-Temperature Interac-tions with Crustal Rocks,” Earth Planet. Sci. Lett. 64,417–429 (1983).
98. J. M. Edmond, C. Measures, B. Mangum, et al., “TheFormation of Metal-Rich Deposits at Ridge Crests,”Earth Planet. Sci. Lett. 46, 19–30 (1979).
99. J. M. Edmond, K. L. Von Damm, R. E. McDuff, andC. J. Measures, “Chemistry of Hot Springs on the EastPacific Rise and Their Effluent Dispersal,” Nature 297,187–191 (1982).
100. J. M. Edmond, A. C. Campbell, M. R. Palmer, et al.,“Time Series Studies of Vent Fluids from the TAG andMARK Sites (1986, 1990), Mid-Atlantic Ridge: A NewSolution Chemistry Model and a Mechanism for Cu/ZnZonation in Massive Sulfide Orebodies,” in Hydrother-mal Vents and Processes, Ed. by L. M. Parson,C. L. Walker, and D. P. Dixon (Geol. Soc. Spec. Publ.,London, 1995), No. 87, pp. 77–86.
101. H. Edmonds and J. M. Edmond, “A Three-ComponentMixing Model for Ridge-Crest Hydrothermal Fluids,”Earth Planet. Sci. Lett. 134, 53–67 (1995).
102. A. J. Ellis, “Explored Geothermal Systems,” inGeochemistry of Hydrothermal Ore Deposits, Ed. byH. Barnes, 2nd ed. (Wiley, New York, 1979),pp. 632–737.
103. R. W. Embley, I. R. Jonasson, M. R. Perfit, et al., “Sub-mersible Investigation of an Extinct Hydrothermal Sys-tem on the Galapagos Ridge: Sulfide Mounds, Stock-work Zone, and Differentiated Lavas,” Can. Mineral.26, 517–539 (1988).
104. T. Erdey-Gru, Transport Phenomena in Aqueous Solu-tions (Akadémiai Kiadó, Budapest, 1974).
105. W. C. Evans, L. D. White, and J. B. Rapp, “Geochemis-try of Some Gases in Hydrothermal Fluids from theSouthern Juan de Fuca Ridge,” J. Geophys. Res. 93,15 305–15 313 (1988).
106. G. Faure, Principles of Isotope Geology, 2nd ed. (Wiley,New York, 1986; Mir, Moscow, 1989).
107. R. A. Feely, T. L. Geiselman, E. T. Baker, and G. J. Mas-soth, “Distribution and Composition of HydrothermalPlume Particles from the ASHES Vent Field of AxialVolcano, Juan de Fuca Ridge,” J. Geophys. Res. 95,12 855–12 873 (1990).
108. U. Fehn and L. M. Cathles, “Hydrothermal Convectionat Slow-Spreading Mid-Ocean Ridges,” Tectonophysics55, 239–260 (1979).
109. R. Fernandez-Prini and R. Crovetto, “A Critical Evalu-ation of the Solubility of Simple Inorganic Gases inWater at High Temperature,” AIChE J. 31, 513–516(1985).
110. R. Fernandez-Prini and R. Crovetto, “Evaluation ofData on Solubility of Simple Apolar Gases in Light andHeavy Water at High Temperature,” J. Phys. Chem.Refer. Data 18, 1231–1243 (1989).
111. T. Finlow-Bates and D. E. Large, “Water Depth as aMajor Control on the Formation of Submarine Exhala-tive Ore Deposits,” Geol. Jahrb. 30, 27–39 (1978).
112. A. T. Fisher and T. N. Narasimhan, “Numerical Simula-tion of Hydrothermal Circulation Resulting from BasaltIntrusion in a Buried Spreading Centers,” Earth Planet.Sci. Lett. 103, 100–115 (1991).
113. D. E. Fisher and M. R. Perfit, “Evidence from RareGases for Magma-Chamber Degassing of HighlyEvolved Mid-Ocean-Ridge Basalt,” Nature 343, 450–452 (1990).
114. D. J. Fornari, T. Shank, K. L. Von Damm, et al., “Time-Series Temperature Measurements at High-TemperatureHydrothermal Vents, East Pacific Rise 9°49′–51′ N: Evi-dence for Monitoring a Crustal Cracking Event,” EarthPlanet. Sci. Lett. 160, 419–431 (1998).
115. A. M. Fouillac and M. Javoy, “Oxygen and HydrogenIsotopes in the Volcano–Sedimentary Complex ofHuelva (Iberian Pyrite Belt): Example of Water Circu-lation through a Volcano–Sedimentary Sequence,”Earth Planet. Sci. Lett. 87, 473–484 (1988).
116. C. Fouillac, G. Michard, and G. Bocquier, “Une Meth-ode de Simulation de l’Evolution des Profils d’Alter-ation,” Geochim. Cosmochim. Acta 41, 207–213(1977).
117. Y. Fouquet, G. Auclair, P. Cambon, and J. Etoubleau,“Geological Setting and Geochemical Investigations onSulfide Deposits near 13° N on the East Pacific Rise,”Mar. Geol. 84, 143–178 (1988).
118. Y. Fouquet, U. Von Stackelberg, J. L. Charlou, et al.,“Metallogenesis in Back-Arc Environments: The LauBasin Example,” Econ. Geol. 88, 2154–2181 (1993).
119. Y. Fouquet, R. Knott, P. Cambon, et al., “Formation ofLarge Sulfide Mineral Deposits along Fast SpreadingRidges: Examples from Off-Axial Deposits at 12°43′ Non the East Pacific Rise,” Earth Planet. Sci. Lett. 144,147–162 (1996).
120. C. G. Fox, “Consequences of Phase Separation on theDistribution of Hydrothermal Fluids at ASHES VentField, Axial Volcano, Juan de Fuca Ridge,” J. Geophys.Res. 95, 12 923–12 926 (1990).
121. R. D. Francis, “Sulfide Globules in Mid-Ocean RidgeBasalts (MORB), and the Effect of Oxygen Abundancein Fe–S–O Liquids on the Ability of These Liquids toPartition Metals from MORB and Komatiite Magmas,”Chem. Geol. 85, 199–213 (1990).
122. J. M. Franklin, J. W. Lydon, and D. F. Sangster, “Volca-nic-Associated Massive Sulfide Deposits,” Econ. Geol.75, 485–627 (1981).
123. I. Friedman and J. R. O’Neil, “Compilation of StableIsotope Fractionation Factors of Geochemical Interest,”in Data of Geochemistry, 6th ed. US Geol. Surv. Prof.Pap., No. 440 (1977).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S315
124. B. Fritz, “Etude thermodynamique et simulation desreactions entre mineraux et solutions: application a lageochimie des alterations et des eaux continentales,”Memoire Sci. Geol., No. 41 (1975).
125. B. Fritz, “Etude thermodinamique et modelisation desreactions hydrothermales et diagenetiques,” MemoirSci. Geol., No. 65 (1981).
126. W. S. Fyfe, N. J. Price, and A. B. Thompson, Fluids inthe Earth Crust (Elsevier, Amsterdam, 1978).
127. T. Gamo, H. Sakai, E.-S. Kim, et al., “High AlkalinityDue to Sulfate Reduction in the CLAM HydrothermalField, Okinawa Trough,” Earth Planet. Sci. Lett. 107,328–338 (1991).
128. T. Gamo, K. Okamura, J.-L. Charlou, et al., “Acidic andSulfate-Rich Hydrothermal Fluids from the ManusBack-Arc Basin, Papua New Guinea,” Geology 25,139–142 (1997).
129. T. Gamo, H. Chiba, T. Yamanaka, et al., “ChemicalCharacteristics of Newly Discovered Black SmokerFluids and Associated Hydrothermal Plumes at theRodrigues Triple Junction, Central Indian Ridge,” EarthPlanet. Sci. Lett. 193, 371–379 (2001).
130. R. M. Garrels and C. L. Christ, Solutions, Minerals, andEquilibria (Freeman Cooper, San Francisco, 1965).
131. M. O. Garcia, D. W. Muenow, and K. E. Aggrey, “MajorElement, Volatile, and Stable Isotope Geochemistry ofHawaiian Submarine Tholeiitic Glasses,” J. Geophys.Res. 94, 10525–10538 (1989).
132. J. B. Gemmell and R. Sharpe, “Detailed Sulfur-IsotopeInvestigation of the TAG Hydrothermal Mound andStockwork Zone, 26° N, Mid-Atlantic Ridge,” Proc.Ocean Drill. Program: Sci. Results 158, 71–84 (1998).
133. General Chemical Technology, Ed. by A. G. Amelin(Khimiya, Moscow, 1977) [in Russian].
134. Geophysics of the Ocean, Vol. 2: Geodynamics, Ed. byO. G. Sorokhtin (Nauka, Moscow, 1979) [in Russian].
135. T. M. Gerlach, “Exsolution of H2O, CO2, and S duringEruptive Episodes at Kilauea Volcano, Hawaii,” J. Geo-phys. Res. 91, 12177–12185 (1986).
136. T. M. Gerlach, “CO2 from Magma-Chamber Degas-sing,” Nature 337, 124 (1989).
137. H. A. Gild and S. M. F. Sheppard, “Hydrogen IsotopeFractionation between Kaolinite and Water Revisited,”Geochim. Cosmochim. Acta 60, 529–533 (1996).
138. K. M. Gillis, “Controls on Hydrothermal Alteration in aSection of Fast-Spreading Oceanic Crust,” Earth Planet.Sci. Lett. 134, 473–489 (1995).
139. K. M. Gillis and P. T. Robinson, “Pattern and Processesof Alteration in the Lavas and Dykes of the TroodosOphiolite, Cyprus,” J. Geophys. Res. 95, 21 523–21 548(1990).
140. U. Ginster, M. J. Mottl, and R. P. Von Herzen, “HeatFlux from Black Smokers on the Endeavour and CleftSegments, Juan de Fuca Ridge,” J. Geophys. Res. 99,4937–4950 (1994).
141. E. Gitlin, “Sulfide Remobilization during Low Temper-ature Alteration of Seafloor Basalt,” Geochim. Cosmo-chim. Acta 49, 1567–1579 (1985).
142. G. P. Glasby and K. Notsu, “Submarine HydrothermalMineralization in the Okinawa Trough, SW of Japan:An Overview,” Ore Geol. Rev. 23, 299–339 (2003).
143. S. I. Golyshev, V. A. Grinenko, and N. L. Padalko,“Thermodynamic Isotope Effects for Sulfur and Oxy-gen in the System Solution–Sulfate Mineral,”Geokhimiya, No. 10, 1379–1390 (1983).
144. W. D. Goodfellow and B. Blaise, “Sulfide Formationand Hydrothermal Alteration of Hemipelagic Sedimentin Middle Valley, Northern Juan de Fuca,” Can. Min-eral. 26, 675–696 (1988).
145. W. D. Goodfellow and J. M. Franklin, “Geology, Min-eralogy, and Chemistry of Sediment-Hosted ClasticMassive Sulfides in Shallow Cores, Middle Valley,Northern Juan de Fuca Ridge,” Econ. Geol. 88, 2037–2068 (1993).
146. N. S. Gorbachev and G. A. Kashirtseva, “Fluid–Mag-matic Differentiation of Basaltic Magmas and Mag-matic Sulfide Formation,” in Experiment in the Solutionof Important Geological Problems (Nauka, Moscow,1986), pp. 98–119 [in Russian].
147. S. Goto, M. Kinoshita, O. Matsubayashi, and R. P. VonHerzen, “Geothermal Constraints on the HydrologicalRegime of the TAG Active Hydrothermal Mound,Inferred from Long-Term Monitoring,” Earth Planet.Sci. Lett. 203, 149–163 (2002).
148. U. M. Graham, G. J. Bluth, and H. Omoto, “Sulfide–Sulfate Chimneys on the East Pacific Rise, 11 and 13° NLatitudes: I. Mineralogy and Paragenesis,” Can. Min-eral. 26, 487–504 (1988).
149. D. V. Grichuk, “Estimation of Gibbs Free Energies forIsotopic Forms of Compounds,” Geokhimiya, No. 2,178–191 (1987).
150. D. V. Grichuk, “An Isotopic–Chemical ThermodynamicModel of a Hydrothermal System,” Dokl. Akad. NaukSSSR 298, 1222–1225 (1988).
151. D. V. Grichuk, “Ore Elements in a Mid-Ocean RidgeHydrothermal System,” Geokhimiya, No. 7, 650–672(1996) [Geochem. Int. 34, 586–606 (1996)].
152. D. V. Grichuk, “Estimation of the Isotopic–ChemicalCross Effect,” in Abstracts of XV Symposium on IsotopeGeochemistry (Moscow, 1998), p. 81.
153. D. V. Grichuk and M. V. Borisov, “ThermodynamicModel of a Hydrothermal System in the Oceanic Crust,”Dokl. Akad. Nauk SSSR 270, 424–427 (1983).
154. D. V. Grichuk and S. G. Krasnov, “Sources of Ore Ele-ments in Present-Day Hydrothermal Systems of theOcean Floor,” Litol. Polezn. Iskop., No. 1, 105–113(1989).
155. D. V. Grichuk and A. Yu. Lein, “Evolution of an OceanicHydrothermal System and Sulfur Isotope Compositionof Sulfides,” Dokl. Akad. Nauk SSSR 318, 422–425(1991).
156. D. V. Grichuk and Yu. V. Shvarov, “Boiling as a Factorof Metal Mobilization in Hydrothermal Systems,” inAbstracts of II International Symposium on Thermody-namics of Natural Processes (Novosibirsk, 1992), p.24.
157. D. V. Grichuk and Yu. V. Shvarov, “Comparative Anal-ysis of Techniques Used in the Equilibrium–DynamicSimulations of Infiltration Metasomatic Zoning,”Petrologiya 10, 654–667 (2002) [Petrology 10, 580–592 (2002)].
158. D. V. Grichuk, M. V. Borisov, and G. L. Mel’nikova,“Thermodynamic Modeling of Seawater–Basalt andSeawater–Peridotite Interactions under Hydrothermal
S316
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Conditions,” in Abstracts of V All-Union School onMarine Geology: Geology of Seas and Oceans (1982),Vol. 3, pp. 166–168.
159. D. V. Grichuk, M. V. Borisov, and G. L. Mel’nikova,“Thermodynamic Modeling of a Hydrothermal Systemin Oceanic Crust: Evolution of Fluid Composition,” Int.Geol. Rev. 27, 1097–1115 (1985).
160. D. V. Grichuk, N. M. Sushchevskaya, Yu. V. Vasyuta,and N. N. Kononkova, “On the Role of Magmatic Flu-ids in the Formation of the Chemical Composition ofHydrothermal Waters of Mid-Ocean Ridges,”Geokhimiya, No. 12, 1741–1750 (1988).
161. D. V. Grichuk, E. E. Abramova, and A. V. Tutubalin, “AThermodynamic Model of Submarine Massive SulfideFormation in a Convecting Hydrothermal System,”Geol. Rudn. Mestorozhd., No. 1, 1–15 (1998) [Geol.Ore Depos. 40, 1–15 (1998)].
162. D. Grimaud, J. Ishibashi, J. M. Auzende, and T. Urabe,“The Chemistry of the Hot Springs at 17° S on theNorth Fiji Basin Axis (SW Pacific),” EOS, Trans. Am.Geophys. Union 70, 1398 (1989).
163. L. N. Grinenko, V. A. Grinenko, G. D. Zagryaiskaya,and Yu. M. Stolyarov, “Isotopic Composition of Sulfurin Sulfides and Sulfates from the Levikha Massive Sul-fide Deposits with Applications to Their Genesis,” Geol.Rudn. Mestorozhd., No. 3, 26–38 (1969).
164. V. A. Grinenko and H. G. Thode, “Sulfur IsotopeEffects in Volcanic Gas Mixtures,” Can. J. Earth Sci. 7,1402–1409 (1970).
165. A. A. Gurenko, A. V. Sobolev, and N. N. Kononkova,“Petrology of the Primary Melt of Rift Tholeiites fromthe Reikjanes Peninsula, Iceland,” Geokhimiya, No. 8,1137–1150 (1990).
166. J. L. Haas, Jr., “Physical Properties of the CoexistingPhases and Thermochemical Properties of the H2OComponent in Boiling NaCl Solutions: PreliminarySteam Tables for NaCl Solutions,” Geol. Surv. Bull.,No. 1421-A (1976a).
167. J. L. Haas, Jr., “Thermodynamic Properties of the Coex-isting Phases and Thermochemical Properties of theNaCl Component in Boiling NaCl Solutions: Prelimi-nary Steam Tables for NaCl Solutions,” Geol. Surv.Bull., No. 1421-B (1976b).
168. A. Hajash, “Hydrothermal Processes along Mid-OceanRidges: An Experimental Investigation,” Contrib. Min-eral. Petrol. 53, 205–226 (1975).
169. A. Hajash and G. W. Chandler, “An Experimental Inves-tigation of High-Temperature Interactions between Sea-water and Rhyolite, Andesite, Basalt and Peridotite,”Contrib. Mineral. Petrol. 78, 240–254 (1981).
170. P. Halbach, K. Nakamura, M. Washner, et al., “ProbableModern Analogue of Kuroko-Type Massive SulfideDeposits in the Okinawa Trough Back-Arc Basin,”Nature 338, 496–499 (1989).
171. P. Halbach, B. Pracejus, and A. Marten, “Geology andMineralogy of Massive Sulfide Ores from the CentralOkinawa Trough, Japan,” Econ. Geol. 88, 2210–2225(1993).
172. Handbook on Heat Transfer, Ed. by B. S. Petukhov andV. K. Shikov (Energoatomizdat, Moscow, 1987) [inRussian].
173. M. D. Hannington and S. D. Scott, “Mineralogy andGeochemistry of a Hydrothermal Silica–Sulfide–Sul-fate Spire in the Caldera of Axial Seamount, Juan deFuca Ridge,” Can. Mineral. 26, 603–625 (1988).
174. M. Hannington, P. Herzig, S. Scott, et al., “ComparativeMineralogy and Geochemistry of Gold-Bearing SulfideDeposits on the Mid-Ocean Ridges,” Mar. Geol. 101,217–248 (1991).
175. M. D. Hannington, A. G. Galley, P. M. Herzig, andS. Petersen, “Comparison of the TAG Mound andStockwork Complex with Cyprus-Type Massive SulfideDeposits,” Proc. Ocean Drill. Program: Sci. Results158, 389–416 (1998).
176. M. Hannington, P. Herzig, P. Stoffers, et al., “FirstObservations of High-Temperature Submarine Hydro-thermal Vents and Massive Anhydrite Deposits off theNorth Coast of Iceland,” Mar. Geol. 177, 199–220(2001).
177. R. Haymon, “The Growth History of Hydrothermal‘Black Smoker’ Chimneys,” Nature 301, 695–698(1983).
178. R. M. Haymon, R. A. Koski, and M. J. Abrams, “Hydro-thermal Discharge Zones beneath Massive SulfideDeposits Mapped in the Oman Ophiolite,” Geology 17,531–535 (1989).
179. C. A. Heinrich, “The Chemistry of HydrothermalTin (-Tungsten) Ore Deposition,” Econ. Geol. 85, 457–481 (1990).
180. C. A. Heinrich, J. H. C. Bain, T. P. Mernagh, et al.,“Fluid and Mass Transfer during Metabasalt Alterationand Copper Mineralization at Mount Isa, Australia,”Econ. Geol. 90, 705–730 (1995).
181. H. Heinrichs, B. Schulz-Dobrick, and K. H. Wedepohl,“Terrestrial Geochemistry of Cd, Bi, Tl, Pb, Zn, andRb,” Geochim. Cosmochim. Acta 44, 1519–1533(1980).
182. R. Hekinian and Y. Fouquet, “Volcanism and Metallo-genesis of Axial and Off-Axial Structures on the EastPacific Rise Near 13° N,” Econ. Geol. 80, 221–249(1985).
183. R. Hekinian and D. Walker, “Diversity and SpatialZonation of Volcanic Rocks from the East Pacific Risenear 21° N,” Contrib. Mineral. Petrol. 96, 265–280(1987).
184. R. Hekinian, J. Francheteau, V. Renard, et al., “IntenseHydrothermal Activity at the Axis of the East PacificRise Near 13° N: Submersible Witness of Growth ofSulfide Chimney,” Mar. Geophys. Res. 6, 1–14 (1983).
185. R. Hekinian, G. Thompson, and D. Bideau, “Axial andOff-Axial Heterogeneity of Basaltic Rocks from theEast Pacific Rise at 12°35′ N–12°51′ N and 11°26′ N–11°30′ N,” J. Geophys. Res. 94, 17 437–17 463 (1989).
186. H. C. Helgeson, “Evaluation of Irreversible Reactionsin Geochemical Processes Involving Minerals andAqueous Solutions: I. Thermodynamic Relations,”Geochim. Cosmochim. Acta 32, 853–877 (1968).
187. H. C. Helgeson, “Mass Transfer among Minerals andHydrothermal Solutions,” in Geochemistry of Hydro-thermal Ore Deposits, 2nd ed., Ed. by H. Barnes (Wiley,New York, 1979), pp. 568–610.
188. H. C. Helgeson and D. H. Kirkham, “Theoretical Pre-diction of the Thermodynamic Behavior of Aqueous
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S317
Electrolytes at High Pressures and Temperatures: IIDebye–Hückel Parameters for Activity Coefficients andRelative Partial Molal Properties,” Am. J. Sci. 274,1199–1261 (1974).
189. H. C. Helgeson, R. M. Garrels, and F. T. McKenzie,“Evaluation of Irreversible Reactions in GeochemicalProcesses Involving Minerals and Aqueous Solutions:II. Applications,” Geochim. Cosmochim. Acta 33, 455–481 (1969).
190. H. C. Helgeson, T. H. Brown, A. Nigrini, andT. A. Jones, “Calculation of Mass Transfer in Geochem-ical Processes Involving Aqueous Solutions,” Geochim.Cosmochim. Acta 34, 569–592 (1970).
191. H. C. Helgeson, J. M. Delany, H. W. Nesbitt, andD. K. Bird, “Summary and Critique of the Thermody-namic Properties of Rock-Forming Minerals,” Am. J. Sci.278-a, 1–228 (1978).
192. H. C. Helgeson, D. H. Kirkham, and G. C. Flowers,“Theoretical Prediction of the Thermodynamic Behav-ior of Aqueous Electrolytes at High Pressures and Tem-peratures: IV. Calculation of Activity Coefficients,Osmotic Coefficients, and Apparent Molal and Stan-dard and Relative Partial Molal Properties to 600°C and5 Kb,” Am. J. Sci. 281, 1249–1516 (1981).
193. J. J. Hemley, G. L. Cygan, J. B. Fein, et al., “Hydrother-mal Ore-Forming Processes in the Light of Studies inRock-Buffered Systems: I. Iron–Copper–Zinc–LeadSulfide Solubility Relations,” Econ. Geol. 87, 1–22(1992).
194. P. M. Herzig, M. D. Hannington, and A. J. Arribas,“Sulfur Isotopic Composition of Hydrothermal Precipi-tates from the Lau Back Arc: Implication for MagmaticContributions to Seafloor Hydrothermal Systems,”Miner. Deposita 33, 226–237 (1998a).
195. P. M. Herzig, S. E. Humphris, D. J. Miller, andR. A. Zierenberg, Proc. Ocean Drill. Program: Sci.Results 158, 389–416 (1998).
196. P. M. Herzig, S. Petersen, and M. D. Hannington,“Geochemistry and Sulfur-Isotopic Composition of theTAG Hydrothermal Mound, Mid-Atlantic Ridge, 26° N,”Proc. Ocean Drill. Program: Sci. Results 158, 47–70(1998b).
197. H. D. Holland and S. D. Malinin, “The Solubility andOccurrence of Non-Ore Minerals,” in Geochemistry ofHydrothermal Ore Deposits, 2nd ed., Ed. by H. L. Bar-nes (Wiley, New York, 1979), pp. 461–508.
198. J. R. Holloway, “Magmatic Fluids,” in ThermodynamicModeling of Geological Materials: Minerals, Fluids,and Melts, Vol. 17 of Reviews in Mineralogy, Ed. byI. S. E. Carmichael and H. P. Eugster (1987).
199. J. Honnorez, L. C. Alt, B.-M. Honnorez-Guerstein,et al., “Stockwork-Like Sulfide Mineralization inYoung Oceanic Crust: Deep Sea Drilling Project Hole504B,” Deep Sea Drill. Project: Initial Rep. 83, 263–282 (1985).
200. Hot Brines and Recent Heavy Metal Deposits in the RedSea, Ed. by E. T. Degens and D. A. Ross (Springer-Ver-lag, Berlin, 1969).
201. S. E. Humphris, P. M. Herzig, D. J. Miller, et al., “TheInternal Structure of an Active Sea-Floor Massive Sul-fide Deposit,” Nature 377, 713–716 (1995).
202. S. E. Humphris, J. C. Alt, D. A. H. Teagle, and J. J. Hon-norez, “Geochemical Changes during HydrothermalAlteration of Basement in the Stockwork beneath theActive TAG Hydrothermal Mound,” Proc. Ocean Drill.Program: Sci. Results 158, 255–276 (1998).
203. S. F. Humphris and G. Thompson, “HydrothermalAlteration of Oceanic Basalts by Seawater,” Geochim.Cosmochim. Acta 42, 107–125 (1978).
204. Hydrothermal Sulfide Ores and Metalliferous Sedi-ments in the Ocean, Ed. by I. S. Gramberg andA. I. Ainemer (Nedra, St. Petersburg, 1992) [in Rus-sian].
205. R. D. Hyndman and M. J. Drury, “The Physical Proper-ties of Oceanic Basement Rocks from Deep Drilling onthe Mid-Atlantic Ridge,” J. Geophys. Res. 81, 4042–4052 (1976).
206. S. A. Igumnov, V. A. Grinenko, and N. B. Poner, “Tem-perature Dependence of Sulfur Isotope Partition Coeffi-cients between Hydrogen Sulfide and Dissolved Sulfateat 260–400°C,” Geokhimiya, No. 7, 1085–1087 (1977).
207. R. P. Ilchik and M. D. Barton, “An Amagmatic Originof Carlin-Type Gold Deposits,” Econ. Geol. 92, 269–288 (1997).
208. J. I. Ishibashi, Y. Sano, H. Wakita, et al., “Helium andCarbon Geochemistry of Hydrothermal Fluids from theMid-Okinawa Trough Back Arc Basin, Southwest ofJapan,” Chem. Geol. 123, 1–15 (1995).
209. E. Ito, D. M. Harris, and A. T. Anderson, “Alteration ofOceanic Crust and Geological Cycling of Chlorine andSeawater,” Geochim. Cosmochim. Acta 47, 1613–1624(1983).
210. I. P. Ivanov and M. V. Borisov, “Estimation of the Com-position of Initial Solution for the Process of Metaso-matic Replacement of Rocks,” Geokhimiya, No. 12,1797–1806 (1980).
211. R. H. James, H. Elderfield, and M. R. Palmer, “TheChemistry of Hydrothermal Fluids from the BrokenSpur Site, 29° N Mid-Atlantic Ridge,” Geochim. Cos-mochim. Acta 59, 651–659 (1995).
212. D. R. Janecky and W. E. Seyfried, Jr., “Formation ofMassive Sulfide Deposits on Oceanic Ridge Crests:Incremental Reaction Models for Mixing betweenHydrothermal Solutions and Seawater,” Geochim. Cos-mochim. Acta 48, 2723–2738 (1984).
213. D. R. Janecky and W. E. Seyfried, Jr., “HydrothermalSerpentinization of Peridotite within the Oceanic Crust:Experimental Investigation of Mineralogy and MajorElement Chemistry,” Geochim. Cosmochim. Acta 50,1357–1378 (1986).
214. D. R. Janecky and W. C. Shanks, III, “ComputationalModeling of Chemical and Sulfur Isotopic ReactionProcesses in Seafloor Hydrothermal Systems: Chim-neys, Massive Sulfides, and Subjacent AlterationZones,” Can. Mineral. 26, 603–625 (1988).
215. M. L. Japas and J. M. H. Levelt Sengers, “Gas Solubil-ity and Henry’s Law Near the Solvent’s Critical Point,”AIChE J. 35, 705–713 (1989).
216. P. Jean-Baptiste, H. Bougault, A. Vangriesheim, et al.,“Mantle 3He in Hydrothermal Vents and Plume of theLucky Strike Site (MAR 37°17′ N) and AssociatedGeothermal Heat Flux,” Earth Planet. Sci. Lett. 157,69–77 (1998).
S318
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
217. Ph. Jean-Baptiste, A. Dapoigny, M. Stievenard, et al.,“Helium and Oxygen Isotope Analyses of Hydrother-mal Fluids from the East Pacific Rise between 17° S and19° S,” Geo-Mar. Lett. 17, 213–219 (1997a).
218. Ph. Jean-Baptiste, J. L. Charlou, and M. Stievenadr,“Oxygen Isotope Study of Mid-Ocean Ridge Hydro-thermal Fluids: Implication for the Oxygen-18 Budgetof the Ocean,” Geochim. Cosmochim. Acta 61, 2669–2677 (1997b).
219. Ph. Jean-Baptiste, J. L. Charlou, M. Stievenard, et al.,“Helium and Methane Measurements in HydrothermalFluids from the Mid-Atlantic Ridge: The Snake Pit Siteat 23° N,” Earth Planet. Sci. Lett. 106, 17–28 (1991).
220. J. Jedwab and J. Boulegue, “Graphite Crystals inHydrothermal Vents,” Nature 310, 41–43 (1984).
221. J. W. Johnson, E. H. Oelkers, and H. C. Helgeson,“SUPCRT 92: A Software Package for Calculating theStandard Molal Thermodynamic Properties of Miner-als, Gases, Aqueous Species, and Reactions from 1 to5000 Bars and 0 to 1000°C,” Comput. Geosci. 18, 899–947 (1992).
222. J. Jonasson, R. Embley, T. Francis, et al., “Exposure ofthe Alteration Zone beneath the Galapagos MassiveSulfide Deposit,” EOS, Trans. Am. Geophys. Union 67,86 (1986).
223. D. Kadko and W. Moore, “Radiochemical Constraintson the Crustal Residence Time of Submarine Hydro-thermal Fluids: Endeavour Ridge,” Geochim. Cosmo-chim. Acta 52, 659–668 (1988).
224. Y. Kajiwara, “Sulfur Isotope Study of the Kuroko Oresof Shakanai No. 1 Deposits, Akita Prefecture, Japan,”Geochem. J. 4, 157–181 (1971).
225. V. S. Kamenetsky, P. Davidson, T. P. Mernagh, et al.,“Fluid Bubbles in Melt Inclusions and Pillow-RimGlasses: High-Temperature Precursors to Hydrother-mal Fluids?” Chem. Geol. 183, 349–364 (2002).
226. E. S. Kappel and J. M. Franklin, “Relationship betweenGeologic Development of Ridge Crests and SulfideDeposits in the Northeast Pacific Ocean,” Econ. Geol.84, 485–505 (1988).
227. I. K. Karpov, Physicochemical Computer Modeling inGeochemistry (Nauka, Novosibirsk, 1981) [in Russian].
228. I. K. Karpov, A. I. Kiselev, and F. A. Letnikov, Com-puter Modeling of Natural Mineral Formation (Nedra,Moscow, 1976) [in Russian].
229. K. Kase, M. Yamamoto, and T. Shibata, “Copper-RichSulfide Deposit Near 23° N, Mid-Atlantic Ridge:Chemical Composition, Mineral Chemistry, and SulfurIsotopes,” Proc. Ocean Drill. Program: Sci. Results106/109, 163–177 (1990).
230. S. A. Kashik, Formation of Mineral Zoning in Alter-ation Residues (Nauka, Novosibirsk, 1989) [in Rus-sian].
231. S. A. Kashik and I. K. Karpov, Physicochemical Theoryof Zoning Formation in the Alteration Residues (Nauka,Novosibirsk, 1978) [in Russian].
232. M. Kastner, H. Craig, and A. Sturz, “HydrothermalDeposition in the Mariana Trough: Preliminary Miner-alogical Investigation,” EOS, Trans. Am. Geophys.Union 68, 1531 (1987).
233. H. Kawahata and S. D. Scott, “Strontium Isotopes andWater–Rock Interactions of the Agrokipia “B” Stock-
work Deposit in the Troodos Ophiolite, Cyprus: A Fos-sil Subseafloor Ore Body,” Geochem. J. 24, 349–356(1990).
234. H. Kawahata and N. Shikazono, “Sulfur Isotope andTotal Sulfur Studies of Basalts and Greenstones fromDSDP Hole 504B, Costa Rica Rift: Implications forHydrothermal Alteration,” Can. Mineral. 26, 555–565(1988).
235. D. S. Kelley and J. R. Delaney, “Two Phase Separationand Fracturing in Mid-Ocean Ridge Gabbros at Tem-peratures Greater than 700°C,” Earth Planet. Sci. Lett.83, 53–66 (1987).
236. D. S. Kelley, J. A. Karson, D. K. Blackman, et al., “Dis-covery of Lost City: An Off-Axis Hydrothermal VentField Near the Mid-Atlantic Ridge at 30° N,” Nature412, 145–149 (2001).
237. J. F. Kerridge, R. M. Haymon, and M. Kastner, “SulfurIsotope Systematics at the 21° N Site, East PacificRise,” Earth Planet. Sci. Lett. 66, 91–100 (1983).
238. I. L. Khodakovsky, V. P. Volkov, Yu. I. Sidorov, andM. V. Borisov, “Preliminary Prediction of the MineralComposition of Surface Rocks and the Hydration andOxidation Processes of the Venus Outer Shell,”Geokhimiya, No. 12, 1821–1835 (1978).
239. H. C. Kim and G. M. McMurtry, “Radial Growth Ratesand 210Pb Ages of Hydrothermal Massive Sulfides fromthe Juan de Fuca Ridge,” Earth Planet. Sci. Lett. 104,299–314 (1991).
240. K.-R. Kim, J. A. Welhan, and H. Craig, “The Hydro-thermal Vent Fields at 13° and 11° N on the East PacificRise: Alvin Results,” EOS, Trans. Am. Geophys. Union65, 973 (1984).
241. R. Knott, Y. Fouquet, J. Honnorez, et al., “Petrology ofHydrothermal Mineralization: A Vertical Sectionthrough the TAG Mound,” Proc. Ocean Drill. Program:Sci. Results 158, 5–26 (1998).
242. G. R. Kolonin, G. A. Palyanova, G. P. Shironosova, andK. G. Morgunov, “Thermodynamic Model of PotentialGold-Bearing High-Temperature Water–Carbon Diox-ide Fluid,” Geokhimiya, No. 12, 1725–1734 (1994).
243. M. Yu. Korotaev, A. A. Pek, and M. N. Kim, “InfiltrationMetasomatism in Gradient Field: Modeling of Metaso-matic Vein Formation,” Geokhimiya, No. 1, 20–35(1992).
244. D. S. Korzhinskii, “General Properties of the InfiltrationMetasomatic Zoning,” Dokl. Akad. Nauk SSSR 78 (1),95–98 (1951).
245. D. S. Korzhinskii, Theory of Metasomatic Zoning(Nauka, Moscow, 1969; Clarendon, Oxford, UK, 1970).
246. D. S. Korzhinskii, “Conventional Stationary Systems,”Zap. Vses. Mineral. O–va 108, 522–523 (1979).
247. R. Koski, P. Lonsdale, W. C. Shanks, III, et al., “Miner-alogy and Geochemistry of a Sediment-Hosted Hydro-thermal Sulfide Deposit from the Southern Trough ofGuaymas Basin, Gulf of California,” J. Geophys. Res.90, 6695–6707 (1985).
248. R. Koski, W. C. Shanks, III, W. A. Bohrson, andR. L. Oscarson, “The Composition of Massive SulfideDeposits from the Sediment-Covered Floor of EscanabaTrough, Gorda Ridge: Implications for DepositionalProcesses,” Can. Mineral. 26, 655–673 (1988).
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S319
249. S. G. Krasnov, “On the Minimum Depth of MassiveSulfide Ore Formation on the Ocean Floor,” Dokl.Akad. Nauk SSSR 296, 1188–1191 (1987).
250. S. G. Krasnov, “Geochemical Specialization of OceanicSulfide Ores,” Dokl. Akad. Nauk SSSR 313, 962–966(1990).
251. S. G. Krasnov, Doctoral Dissertation in Geology andMineralogy (St. Petersburg, 1993).
252. S. G. Krasnov and T. V. Stepanova, “The Formation ofa Large Oceanic Massive Sulfide Ore Body (MiddleValley, Endeavour Ridge): Evidence from GeochemicalStudy of the Cores,” Geochem. Int., No. 9, 768–789(1996).
253. S. G. Krasnov, A. I. Ainemer, and T. V. Stepanova,“Geochemistry and Genesis of Sulfide Ores in thePacific,” in Geology of Seas and Oceans (Nauka, Mos-cow, 1988), pp. 129–140 [in Russian].
254. S. G. Krasnov, D. V. Grichuk, and T. V. Stepanova,“Hydrothermal Mineral Formation in the Ocean,” Zap.Vses. Mineral. O–va 116, 23–32 (1990).
255. S. Krasnov, T. Stepanova, and M. Stepanov, “ChemicalComposition and Formation of a Massive SulfideDeposit, Middle Valley, Northern Juan de Fuca Ridge(Site 856),” Proc. Ocean Drill. Program: Sci. Results139, 353–372 (1994).
256. A. I. Krivtsov, “Qualitative Aspects of the Problem ofOre Matter Sources,” Geol. Rudn. Mestorozhd. 23 (5),3–18 (1981).
257. A. I. Krivtsov, “Environments and Conditions ofAncient and Modern Massive Sulfide Ore Formation,”Geol. Rudn. Mestorozhd., No. 3, 3–17 (1987).
258. V. B. Kurnosov, Hydrothermal Alteration of Basalts inthe Pacific Ocean and Metalliferous Sediments: Evi-dence from Deep-Sea Drilling Data (Nauka, Moscow,1986) [in Russian].
259. M. Kusakabe, S. Mayeda, and E. S. Nakamura, “O andSr Isotope Systematics of Active Vent Materials fromthe Mariana Backarc Basin Spreading Axis at 18° N,”Earth Planet. Sci. Lett. 100, 275–282 (1990).
260. T. K. Kyser and J. R. O’Neil, “Hydrogen Isotope Sys-tematics of Submarine Basalts,” Geochim. Cosmochim.Acta 48, 2123–2133 (1984).
261. C. Lalou, J.-L. Reyss, and E. Brichet, “Actinide-SeriesDisequilibrium as a Tool to Establish the Chronology ofDeep-Sea Hydrothermal Activity,” Geochim. Cosmo-chim. Acta 57, 1221–1231 (1993).
262. C. Lalou, J.-L. Reyss, E. Brichet, et al., “HydrothermalActivity on a 105-Year Scale at a Slow Spreading Ridge,TAG Hydrothermal Field, Mid-Atlantic Ridge 26° N,”J. Geophys. Res. 100, 17 855–17 862 (1995).
263. R. R. Large, “Chemical Evolution and Zonation of Mas-sive Sulphide Deposits in Volcanic Terrains,” Econ.Geol. 72, 549–572 (1977).
264. A. C. Lasaga, “Rate Laws of Chemical Reactions,” inKinetics of Geochemical Processes, Vol. 8 of Reviews inMineralogy (1981), pp. 1–68.
265. A. Yu. Lein, “Sulfur and Carbon Isotopes in Hydrother-mal Fields of the Mid-Atlantic Ridge,” Geokhimiya,No. 11, 1162–1173 (2001) [Geochem. Int. 39, 1066–1077 (2001)].
266. A. Yu. Lein, V. F. Gal’chenko, V. A. Grinenko, et al.,“Mineral Composition and Geochemistry of Rockswith Bacterial Overgrowths from Submarine Hydro-thermal Structures,” Geokhimiya, No. 9, 1235–1248(1988).
267. A. Yu. Lein, N. V. Ul’yanova, V. A. Grinenko, andA. P. Lisitsin, “Geochemistry of the Hydrothermal Sul-fide Ores of the Mid-Atlantic Ridge (26° N),”Geokhimiya, No. 3, 307–319 (1991).
268. A. Yu. Lein, N. V. Ul’yanova, V. A. Grinenko, et al.,“Mineralogical and Geochemical Peculiarities ofHydrothermal Sulfide Ores from the Manus Basin (Bis-mark Sea),” Geokhimiya, No. 4, 524–537 (1993).
269. A. Yu. Lein, D. V. Grichuk, E. G. Gurvich, andYu. A. Bogdanov, “New Type of the HydrothermalFluid Enriched in H2 and CH4 at the Rift Zone of theMid-Atlantic Ridge,” Dokl. Akad. Nauk 375, 380–383(2000) [Trans. Russ. Acad. Sci., Earth Sci. Sec. 375,1391–1394 (2000)].
270. A. Yu. Lein, G. A. Cherkashev, A. A. Ul’yanov, et al.,“Mineralogy and Geochemistry of Sulfide Ores fromthe Logatchev-2 and Rainbow Fields: Similar and Dis-tinctive Features,” Geokhimiya, No. 3, 304–328 (2003)[Geochem. Int. 41, 271–294 (2003)].
271. A. Yu. Lein, Yu. A. Bogdanov, A. M. Sagalevich, et al.,“A New Type of Hydrothermal Field in the Mid-Atlan-tic Ridge (Lost City Field, 30° N),” Dokl. Akad Nauk.394, 380–383 (2004) [Trans. Russ. Acad. Sci., EarthSci. Sec. 394, 92–95 (2004)].
272. P. C. Lichtner and V. N. Balashov, “Metasomatic Zon-ing: Appearance of Ghost Zones in the Limit of PureAdvective Mass Transport,” Geochim. Cosmochim.Acta 57, 369–387 (1993).
273. A. P. Lisitsin, Yu. A. Bogdanov, and E. G. Gurvich,Hydrothermal Deposits in Oceanic Rift Zones (Nauka,Moscow, 1990) [in Russian].
274. A. P. Lisitsin, R. A. Binns, Yu. A. Bogdanov, et al.,“Recent Hydrothermal Activity of Franklin Seamount,Western Woodlark Sea (Papua New Guinea),” Izv.Akad. Nauk SSSR, Ser. Geol., No. 8, 125–140 (1991).
275. A. P. Lisitsin, K. Crook, Yu. A. Bogdanov, et al.,“Hydrothermal Field of a Rift Zone in the ManusBasin,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 10, 34–55 (1992).
276. C. R. B. Lister, “On the Thermal Balance of a Mid-Ocean Ridge,” Geophys. J. R. Astron. Soc. 26, 515–535(1972).
277. C. R. B. Lister, “On the Penetration of Water into HotRock,” Geophys. J. R. Astron. Soc. 39, 465–509 (1974).
278. C. R. B. Lister, “Heat Flow and Hydrothermal Circula-tion,” Annu. Rev. Earth Planet. Sci. 8, 95–117 (1980).
279. S. A. Little, K. D. Stolzenbach, and R. P. Von Herzen,“Measurements of Plume Flow from HydrothermalVent Field,” J. Geophys. Res. 92, 2587–2596 (1987).
280. R. P. Lowell, “Circulation in Fractures, Hot Springs,and Convective Heat Transport on Mid-Ocean RidgeCrests,” Geophys. J. R. Astron. Soc. 40, 351–369(1975).
281. K. C. Macdonald and P. J. Fox, “The Axial Summit Gra-ben and Cross-Sectional Shape of the East Pacific Riseas Indicators of Axial Magma Chambers and Recent
S320
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
Volcanic Eruptions,” Earth Planet. Sci. Lett. 88, 119–131 (1988).
282. K. C. Macdonald, K. Becker, F. N. Spiess, andR. D. Ballard, “Hydrothermal Heat Flux of the ‘BlackSmoker’ Vents on the East Pacific Rise,” Earth Planet.Sci. Lett. 48, 1–7 (1980).
283. V. I. Mal’kovskii and A. A. Pek, “The influence of aHighly Permeable Fault on the Structure of ThermalConvection Flow in Ocean Floor Spreading Zones,”Dokl. Akad. Nauk 354, 787–789 (1997) [Trans. Russ.Acad. Sci., Earth Sci. Sec. 355, 672–675 (1997)].
284. K. Marumo and K. H. Hattori, “Seafloor HydrothermalClay Alteration at Jade in the Back-Arc OkinawaTrough: Mineralogy, Geochemistry and Isotope Char-acteristics,” Geochim. Cosmochim. Acta 63, 2785–2804 (1999).
285. V. V. Maslennikov, Sedimentogenesis, Halmyrolysis,and Ecology of Massive Sulfide-Bearing Paleohydro-thermal Fields: Evidence from the Southern Urals(Geotur, Miass, 1999) [in Russian].
286. J. J. Massoth, D. A. Butterfield, J. E. Lupton, et al.,“Submarine Venting of Phase-Separated HydrothermalFluids at Axial Volcano, Juan de Fuca Ridge,” Nature340, 702–705 (1989).
287. J. S. McClain, J. A. Orcutt, and M. Burnett, “The EastPacific Rise in Cross Section: A Seismic Model,” J.Geophys. Res. 90, 8627–8639 (1985).
288. M. T. McCulloch, R. T. Gregory, G. J. Wasserburg, andH. P. Taylor, “A Neodymium, Strontium, and OxygenIsotopic Study of the Cretaceous Samail Ophiolite andImplications for the Petrogenesis and Seawater–Hydro-thermal Alteration of Oceanic Crust,” Earth Planet. Sci.Lett. 46, 201–211 (1980).
289. I. A. Menyailov, L. P. Nikitina, and V. N. Shapar,“Geochemical Pecularities of Volcanic Gases,” in GreatFissure Tolbachik Eruption (Nauka, Moscow, 1984),pp. 285–309 [in Russian].
290. L. Merlivat, C. Andrie, and P. Jean-Baptiste, “Distribu-tion des isotopes de l’hydrogene, de l’oxygene et del’helium dans les sources hydrothermales sous-marinesde la ride Est-Pacifique, 13° N,” C. R. Acad. Sci., Ser. II299, 1191–1196 (1984).
291. L. Merlivat, F. Pineau, and M. Javoy, “HydrothermalVent Waters at 13° N on the East Pacific Rise: IsotopicComposition and Gas Concentration,” Earth Planet. Sci.Lett. 84, 100–108 (1987).
292. Metasomatism and Metasomatic Rocks, Ed. byV. A. Zharikov and V. L. Rusinov (Nauchniy Mir, Mos-cow, 1998) [in Russian].
293. Methods of Geochemical Modelling and Forecasting inHydrogeology, Ed. by S. R. Krainov (Nedra, Moscow,1988) [in Russian].
294. S. Metz and J. H. Trefry, “Chemical and MineralogicalInfluences on Concentrations of Trace Metals in Hydro-thermal Fluids,” Geochim. Cosmochim. Acta 64, 2267–2279 (2000).
295. G. Michard, F. Albarede, A. Michard, et al., “Chemistryof Solutions from the 13° N East Pacific Rise Hydro-thermal Site,” Earth Planet. Sci. Lett. 67, 297–307(1984).
296. M. V. Mironenko and A. A. Kosorukov, “Calculation ofthe Phase Composition of Fluid Systems,” Geokhimiya,No. 8, 1195–1200 (1990).
297. M. J. Mottl, “Metabasalts, Axial Hot Springs, and theStructure of Hydrothermal Systems at Mid-OceanRidges,” Geol. Soc. Am. Bull. 94, 161–180 (1983).
298. M. J. Mottl and H. D. Holland, “Chemical Exchangeduring Hydrothermal Alteration of Basalt by Seawater:I. Experimental Results for Major and Minor Compo-nents of Seawater,” Geochim. Cosmochim. Acta 42,1103–1115 (1978).
299. M. J. Mottl and W. E. Seyfried, Jr., “Sub-SeafloorHydrothermal Systems: Rock vs. Seawater Domi-nated,” in Seafloor Spreading Centers of HydrothermalSystems, Benchmark Pap. Geol. 56, 66–82 (1980).
300. M. J. Mottl, H. D. Holland, and R. F. Corr, “ChemicalExchange during Hydrothermal Alteration of Basalt bySeawater: II. Experimental Results for Fe, Mn, and Sul-fur Species,” Geochim. Cosmochim. Acta 43, 869–884(1979).
301. N. N. Mozgova, Yu. S. Borodaev, A. V. Efimov, et al.,“Silver Minerals in Oceanic Hydrothermal Ore Depos-its (Manus and Woodlark Basins, Papua New GuineaRegion),” Geol. Rudn. Mestorozhd. 35, 333–343(1993).
302. A. Yu. Namiot, Solubilities of Gases in Water: A Hand-book (Nedra, Moscow) [in Russian].
303. P. Nehlig, T. Juteau, V. Bendel, and J. Cotteu, “The RootZones of Oceanic Hydrothermal Systems: Constraintsfrom the Samail Ophiolite (Oman),” J. Geophys. Res.99, 4703–4713 (1994).
304. G. V. Nesterenko and A. I. Almukhamedov, Geochemis-try of Differentiated Trapps (Nauka, Moscow, 1973) [inRussian].
305. D. Norton and J. Knight, “Transport Phenomena inHydrothermal Systems: Cooling Plutons,” Am. J. Sci.277, 937–981 (1977).
306. H. Ohmoto, “Systematics of Sulfur and Carbon Iso-topes in Hydrothermal Ore Deposits,” Econ. Geol. 65,551–578 (1972).
307. H. Ohmoto and A. Lasaga, “Kinetics of Reactionsbetween Aqueous Sulfates and Sulfides in Hydrother-mal Systems,” Geochim. Cosmochim. Acta 46, 1727–1746 (1982).
308. H. Ohmoto and R. Rye, “Hydrogen and Oxygen Isoto-pic Composition of Fluid Inclusion in the KurokoDeposits,” Econ. Geol. 69, 947–953 (1974).
309. H. Ohmoto and R. O. Rye, “Isotopes of Sulfur and Car-bon,” in Geochemistry of Hydrothermal Ore Deposits,2nd ed., Ed. by H. Barnes (Wiley, New York, 1979),pp. 509–567.
310. H. Ohmoto, M. Mizukami, S. E. Drummond, et al.,“Chemical Processes of Kuroko Formation,” in TheKuroko and Related Volcanogenic Massive SulfideDeposits, Econ. Geol. Monogr. Ser. 5, 570–604 (1983).
311. S. E. Oosting and K. L. Von Damm, “Bromide–HalideFractionation in Seafloor Hydrothermal Fluids from 9–10° N East Pacific Rise,” Earth Planet. Sci. Lett. 144,133–145 (1996).
312. H. Paulick, D. A. Vanko, and C. J. Yeats, “Drill Core-Based Facies Reconstruction of a Deep-Marine FelsicVolcano Hosting an Active Hydrothermal System (Pual
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S321
Ridge, Papua New Guinea, ODP Leg 193),” J. Volcanol.Geotherm. Res. 130, 31–50 (2004).
313. M. R. Palmer and J. M. Edmond, “Cesium and Rubid-ium in Submarine Hydrothermal Fluids: Evidence forRecycling of Alkali Elements,” Earth Planet. Sci. Lett.95, 8–14 (1989).
314. D. Y. Peng and D. B. Robinson, “A New Two-ConstantEquation of State,” Ind. Eng. Chem. Res. 15, 59–64(1976).
315. M. R. Perfit, D. J. Fornari, A. Malahoff, and R. W. Emb-ley, “Geochemical Studies of Abyssal Lavas Recoveredby DSRV Alvin from Eastern Galapagos Rift, IncaTransform, and Ecuador Rift: 3. Trace Element Abun-dances and Petrogenesis,” J. Geophys. Res. 88, 551–572 (1983).
316. J. M. Peter and W. C. Shanks, III, “Sulfur, Carbon, andOxygen Isotope Variations in Submarine HydrothermalDeposits of Guaymas Basin, Gulf of California, USA,”Geochim. Cosmochim. Acta 56, 2025–2040 (1992).
317. S. Petersen, P. M. Herzig, and M. D. Hannington, “FluidInclusion Studies as a Guide to the Temperature Regimewithin the TAG Hydrothermal Mound, 26° N, Mid-Atlantic Ridge,” Proc. Ocean Drill. Program: Sci.Results 158, 163–178 (1998).
318. R. A. Phinney and R. I. Odom, “Seismic Studies ofCrustal Structure,” Rev. Geophys. Space Phys. 21,1318–1332 (1983).
319. Physical Properties of Rocks and Minerals: Handbookfor Geophysics, 2nd ed., Ed. by N. B. Dortman (Nedra,Moscow, 1984) [in Russian].
320. L. P. Plyusnina, Experimental Studies of the Metamor-phism of Basic Rocks (Nedra, Moscow, 1983) [in Rus-sian].
321. A. V. Plyasunov and I. P. Ivanov, “Solubility of ZincOxide in Sodium Chloride Solutions up to 600°C and1000 bars,” Geokhimiya, No. 11, 1605–1617 (1990).
322. Principles of Chemical Technology, 4th ed., Ed. byI. P. Mukhlenov, A. E. Gorenshtein, and E. S. Tumarkina(Vysshaya Shkola, Moscow, 1991) [in Russian].
323. Prospecting for Copper Ore Deposits, Ed. byM. B. Borodaevskaya, R. N. Volodin, A. I. Krivtsov, et al.(Nedra, Moscow, 1985) [in Russian].
324. R. P. Rafal’sky, Water–Rock Interaction under Hydro-thermal Conditions (Nauka, Moscow, 1993) [in Rus-sian].
325. M. H. Reed, “Calculation of Multicomponent ChemicalEquilibria and Reaction Processes in Systems InvolvingMinerals, Gases, and an Aqueous Phase,” Geochim.Cosmochim. Acta 46, 513–528 (1982).
326. M. H. Reed, “Seawater–Basalt Reaction and the Originof Greenstones and Related Ore Deposits,” Econ. Geol.78, 446–485 (1983).
327. R. J. Ribando, K. E. Torrance, and D. L. Turcotte,“Numerical Models for Hydrothermal Circulation inthe Oceanic Crust,” J. Geophys. Res. 81, 3007–3012(1976).
328. C. J. Richardson, J. R. Cann, H. G. Richards, andJ. G. Cowan, “Metal-Depleted Root Zones of the Troo-dos Ore-Forming Hydrothermal Systems, Cyprus,”Earth Planet. Sci. Lett. 84, 243–253 (1987).
329. J. D. Ridge, “Volcanic Exhalations and Ore Depositionin the Vicinity of the Seafloor,” Miner. Deposita 8, 332–348 (1973).
330. W. I. Ridley, M. R. Perfit, I. R. Jonesson, andM. F. Smith, “Hydrothermal Alteration in OceanicRidge Volcanoes: A Detailed Study at the GalapagosFossil Hydrothermal Field,” Geochim. Cosmochim.Acta 58, 2477–2494 (1994).
331. H. Riedesel, J. A. Orcutt, K. C. Macdonald, andJ. S. McClain, “Microearthquakes in the Black SmokerHydrothermal Field, East Pacific Rise at 21° N,” J. Geo-phys. Res. 87, 613–623 (1982).
332. RISE Project Group: F. N. Spiess, K. C. Macdonald,T. Atwater, et al., “East Pacific Rise: Hot Springs andGeophysical Experiments,” Science 207, 1421–1433(1980).
333. B. E. Roberts and P. R. Tremaine, “Vapour Liquid Equi-librium Calculations for Dilute Aqueous Solutions ofCO2, H2S, NH3, and NaOH to 300°C,” Can. J. Chem.Eng. 63, 294–300 (1985).
334. R. A. Robie, B. S. Hemingway, and J. R. Fisher, “Ther-modynamic Properties of Minerals and Related Sub-stances at 298.15 K and 1 Bar (105 Pascals) Pressureand at High Temperatures,” Geol. Surv. Bull., No. 1452(1978).
335. P. A. Rona, “Hydrothermal Mineralization at SeafloorSpreading Centers,” Earth Sci. Rev. 20, 1–104 (1984).
336. P. Rona and S. D. Scott, “A Special Issue on Sea-FloorHydrothermal Mineralization: New Perspectives. Pref-ace,” Econ. Geol. 88, 1935–1976 (1993).
337. P. A. Rona, M. D. Hannington, C. V. Raman, et al.,“Active and Relict Sea-Floor Hydrothermal Mineraliza-tion at the TAG Hydrothermal Field, Mid-AtlanticRidge,” Econ. Geol. 88, 1989–2017 (1993).
338. R. J. Rosenbauer and J. L. Bischoff, “Uptake and Trans-port of Heavy Metals by Heated Seawater: A Summaryof the Experimental Results,” in Hydrothermal Pro-cesses at Seafloor Spreading Centers, Ed. by P. A. Ronaet al. (Springer-Verlag, New York, 1983), pp. 177–197.
339. K. Rubin, “Degassing of Metals and Metalloids fromErupting Seamount and Mid-Ocean Ridge Volcanoes:Observations and Predictions,” Geochim. Cosmochim.Acta 61, 3525–3542 (1997).
340. B. N. Ryzhenko, Thermodynamics of Equilibria inHydrothermal Solutions (Nauka, Moscow, 1981) [inRussian].
341. H. Sakai, “Isotopic Properties of Sulfur Compounds inHydrothermal Processes,” Geochem. J. 2, 29–49(1968).
342. H. Sakai and O. Matsubaya, “Isotopic Geochemistry ofthe Thermal Waters of Japan and Its Bearing on theKuroko Ore Solutions,” Econ. Geol. 69, 974–991(1974).
343. H. Sakai, D. J. Des Marias, A. Ueda, and J. G. Moore,“Concentrations and Isotope Ratios of Carbon, Nitro-gen, and Sulfur in Ocean Floor Basalts,” Geochim. Cos-mochim. Acta 48, 3282–3294 (1984).
344. P. Sarda and D. Graham, “Mid-Ocean Ridge PoppingRocks: Implication for Degassing at Ridge Crests,”Earth Planet. Sci. Lett. 97, 268–289 (1990).
S322
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
345. A. Sasaki, “Seawater Sulfate as a Possible Determinantfor Sulfur Isotopic Compounds in Some StrataboundSulfide Ores,” Geochem. J. 4, 41–51 (1970).
346. J. Seewald, A. Cruse, and P. Saccocia, “Aqueous Vola-tiles in Hydrothermal Fluids from the Main EndeavourField, Northern Juan de Fuca Ridge: Temporal Variabil-ity Following Earthquake Activity,” Earth Planet. Sci.Lett. 216, 575–590 (2003).
347. J. C. Seitz, J. G. Blencoe, D. B. Joyce, and R. J. Bodnar,“Volumetric Properties of CO2–CH4–N2 Fluids at200°C and 1000 Bars: A Comparison of Equations ofState and Experimental Data,” Geochim. Cosmochim.Acta 58, 1065–1071 (1994).
348. T. M. Seward, “The Formation of Lead (II) ChlorideComplexes to 300°C: A Spectrophotometric Study,”Geochim. Cosmochim. Acta 48, 121–134 (1984).
349. W. Seyfried and J. L. Bischoff, “Hydrothermal Trans-port of Heavy Metals by Seawater: The Role of Seawa-ter/Basalt Ratio,” Earth Planet. Sci. Lett. 34, 71–77(1977).
350. W. E. Seyfried, Jr. and J. L. Bischoff, “Low Tempera-ture Basalt Alteration by Seawater: An ExperimentalStudy at 70°C and 150°C,” Geochim. Cosmochim. Acta43, 1937–1947 (1979).
351. W. E. Seyfried, Jr. and W. E. Dibble, Jr., “Seawater–Peridotite Interaction at 300°C and 500 Bar: Implica-tion for the Origin of Oceanic Serpentinites,” Geochim.Cosmochim. Acta 44, 309–321 (1980).
352. W. E. Seyfried, Jr. and D. R. Janecky, “Heavy Metal andSulfur Transport during Subcritical and SupercriticalHydrothermal Alteration of Basalt: Influence of FluidPressure and Basalt Composition and Crystallinity,”Geochim. Cosmochim. Acta 49, 2545–2560 (1985).
353. W. E. Seyfried, Jr. and M. J. Mottl, “HydrothermalAlteration of Basalt by Seawater under Seawater-Dom-inated Conditions,” Geochim. Cosmochim. Acta 46,985–1002 (1982).
354. W. E. Seyfried, Jr., P. C. Gordon, and F. W. Dickson, “ANew Reaction Cell for Hydrothermal Solution Equip-ment,” Am. Mineral. 64, 646–649 (1979).
355. W. E. Seyfried, Jr., D. R. Janecky, and M. J. Mottl,“Alteration of the Oceanic Crust: Implication forGeochemical Cycles of Lithium and Boron,” Geochim.Cosmochim. Acta 48, 557–569 (1984).
356. W. E. Seyfried, Jr., M. E. Berndt, and D. R. Janecky,“Chloride Depletions and Enrichments in SeafloorHydrothermal Fluids: Constraints from ExperimentalBasalt Alteration Studies,” Geochim. Cosmochim. Acta50, 469–475 (1986).
357. T. N. Shadlun, N. S. Bortnikov, Yu. A. Bogdanov, et al.,“Mineral Composition, Textures, and Formational Con-ditions of Modern Sulfide Ores in the Rift Zone ofManus Basin,” Geol. Rudn. Mestorozhd. 34 (5), 3–21(1992).
358. W. C. Shanks, III and W. E. Seyfried, Jr., “Stable Iso-tope Studies of Vent Fluids and Chimney Minerals.Southern Juan de Fuca Ridge: Sodium Metasomatismand Seawater Sulfate Reduction,” J. Geophys. Res. 92,11 387–11 399 (1987).
359. V. N. Sharapov and V. A. Akimtsev, “Ore-FormingMagmatic Systems in the Mid-Oceanic Ridges,” Geol.Geofiz. 38, 1289–1304 (1997).
360. R. Shiraki, H. Sakai, M. Endoh, and N. Kishima,“Experimental Studies on Rhyolite– and Andesite–Sea-water Interactions at 300°C and 1000 Bars,” Geochem.J. 21, 139–148 (1987).
361. E. L. Shock and H. C. Helgeson, “Calculation of theThermodynamic and Transport Properties of AqueousSpecies at High Pressures and Temperatures: Correla-tion Algorithms for Ionic Species and Equation of StatePrediction to 5 Kb and 1000°C,” Geochim. Cosmo-chim. Acta 52, 2009–2036 (1988).
362. Yu. V. Shvarov, V. A. Zharikov, and T. B. Zhandarova,“Calculation of an Infiltrational Metasomatic Columnon the Basis of the Local Equilibrium Principle,”Geochem. Int., No. 11, 1041–1049 (2000).
363. G. Yu. Shvedenkov and A. V. Savinov, “Thermody-namic Calculation of the Critical Parameters of Hydro-gen Sulfide–Water Mixtures,” Geol. Geofiz. 35 (11), 3–12 (1994).
364. S. A. Silantyev, Metamorphic Rocks of the AtlanticOcean Floor (Nauka, Moscow, 1984) [in Russian].
365. S. A. Silantyev, M. V. Mironenko, B. A. Bazylev, andYu. V. Semenov, “Metamorphism Associated with Mid-Ocean Ridge Hydrothermal Systems: ThermodynamicModeling,” Geokhimiya, No. 7, 1015–1034 (1992).
366. R. Skirrow and M. L. Coleman, “Origin of Sulphur andGeothermometry of Hydrothermal Sulphides from theGalapagos Rift, 86° W,” Nature 299, 142–144 (1982).
367. N. S. Skripchenko, Hydrothermal–Sedimentary SulfideOres of Basaltoid Associations (Nedra, Moscow, 1972)[in Russian].
368. N. H. Sleep, “Hydrothermal Convection at RidgeAxes,” in Hydrothermal Processes of Seafloor Spread-ing Centers, Ed. by P. A. Rona et al. (Springer-Verlag,New York, 1983), pp. 71–82.
369. V. I. Smirnov, Geology of Mineral Deposits, 4th ed.(Nedra, Moscow, 1982) [in Russian].
370. M. Solomon and D. Walshe, “The Formation of Mas-sive Sulfide Deposits on the Sea Floor,” Econ. Geol. 74,797–813 (1979).
371. M. A. Sommer and E. K. Gibson, “Volatile Concentra-tions of Submarine Glasses from the Juan de FucaRidge,” EOS, Trans. Am. Geophys. Union 66, 926(1985).
372. E. T. S. Spooner, R. D. Beckinsale, P. C. England, andA. Senior, “Hydration, 18O Enrichment and Oxidationduring Ocean Floor Hydrothermal Metamorphism ofOphiolitic Metabasic Rocks from E. Liguria, Italy,”Geochim. Cosmochim. Acta 41, 875–871 (1977).
373. N. F. Spycher and M. H. Reed, “Fugacity Coefficientsof H2, CO2, CH4, H2O, and of H2O–CO2–CH4 Mix-tures: A Virial Equation Treatment for Moderate Pres-sures Applicable to Calculations of Hydrothermal Boil-ing,” Geochim. Cosmochim. Acta 52, 739–749 (1988).
374. N. F. Spycher and M. H. Reed, “Evolution of a Broad-lands-Type Epithermal Ore Fluid along Alternative P–TPaths: Implications for the Transport and Deposition ofBase, Precious, and Volatile Metals,” Econ. Geol. 84,328–359 (1989).
375. D. S. Stakes, J. W. Shervais, and C. A. Hopson, “TheVolcanic–Tectonic Cycle of the FAMOUS and AMARValleys, Mid-Atlantic Ridge (36°47′ N): Evidence fromBasalt Glass and Phenocryst Compositional Variations
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
THERMODYNAMIC MODELS OF SUBMARINE HYDROTHERMAL SYSTEMS S323
for a Steady State Magma Chamber Beneath the ValleyMidsections, AMAR3,” J. Geophys. Res. 89, 8995–7028 (1984).
376. C. I. Steefel and A. C. A. Lasaga, “Coupled Model forTransport of Multiple Chemical Species and KineticPrecipitation/Dissolution Reactions with Application toReactive Flow in Single Phase Systems,” Am. J. Sci.294, 529–592 (1994).
377. Yu. M. Stolyarov, “Anhydrite and Gypsum in MassiveSulfide Deposits of the Urals and the Problem of TheirGenesis,” Geol. Rudn. Mestorozhd., No. 2, 107–113(1965).
378. Yu. M. Stolyarov, “On the Sulfate Mineralization inMassive Sulfide Deposits of the Southern Urals,” Zap.Vses. Mineral. O–va 101, 352–355 (1972).
379. R. Stryjek and J. H. Vera, “PRSV: An Improved Peng–Robinson Equation of State for Pure Compounds andMixtures,” Can. J. Chem. Eng. 64, 323–333 (1986).
380. M. M. Styrt, A. J. Brockman, H. D. Holland, et al.,“Mineralogy and the Isotopic Composition of Sulfur inHydrothermal Sulfide–Sulfate Deposits on the EastPacific Rise, 21° N Latitude,” Earth Planet. Sci. Lett. 53,382–390 (1981).
381. T. N. Surin, Metasomatism and Massive Sulfide Forma-tion. High-Ural Region (Nauka, Yekaterinburg, 1993)[in Russian].
382. R. B. Symonds and M. H. Reed, “Calculation of Multi-component Chemical Equilibria in Gas–Solid–LiquidSystems: Calculation Methods, Thermochemical Data,and Applications to Studies of High-Temperature Vol-canic Gases with Examples from Mount St. Helens,”Am. J. Sci. 293, 758–864 (1993).
383. R. B. Symonds, M. H. Reed, and W. I. Rose, “Origin,Speciation and Fluxes of Trace Element Gases atAugustine Volcano, Alaska: Insights into MagmaDegassing and Fumarolic Processes,” Geochim. Cos-mochim. Acta 56, 633–657 (1992).
384. J. C. Tanger, IV and K. S. Pitzer, “Thermodynamics ofNaCl–H2O: A New Equation of State for the Near-Crit-ical Region and Comparisons with Other Equations forAdjoining Regions,” Geochim. Cosmochim. Acta 53,973–987 (1989).
385. Yu. A. Taran, J. W. Hedenquist, M. A. Korzhinsky, et al.,“Geochemistry of Magmatic Gases from KudryavyVolcano, Iturup, Kuril Islands,” Geochim. Cosmochim.Acta 59, 1749–1761 (1995).
386. B. E. Taylor, “Carbon Dioxide and Methane in Hydro-thermal Vent Fluids from Middle Valley, a Sediment-Covered Ridge Segment,” EOS, Trans. Am. Geophys.Union 71, 1569 (1990).
387. Thermal Constants of Substances, Ed. by V. P. Glushko(VINITI, Moscow, 1965–1979) [in Russian].
388. Thermodynamic Properties of Pure Substances, Ed. byV. P. Glushko (Nauka, Moscow, 1978–1982), Vols. 1–4[in Russian].
389. E. C. Thornton and W. E. Seyfried, Jr., “Sediment–Sea-water Interaction at 200 and 300°C, 500 Bars Pressure:The Role of Sediment Composition in Diagenesis andLow-Grade Metamorphism of Marine Clay,” Bull.Geol. Soc. Am. 96, 1287–1295 (1985).
390. E. C. Thornton and W. E. Seyfried, Jr., “Reactivity ofOrganic-Rich Sediment in Seawater at 350°C, 500 Bars:
Experimental and Theoretical Constraints and Implica-tions for Guaymas Basin Hydrothermal System,”Geochim. Cosmochim. Acta 51, 1997–2001 (1987).
391. P. P. Timofeev, V. N. Kholodov, and V. P. Zverev, “TheBalance of Water in the Recent Sedimentation Process,”Dokl. Akad. Nauk SSSR 287, 1465–1469 (1986).
392. M. K. Tivey, “The Influence of Hydrothermal FluidComposition and Advection Rates on Black SmokerChimney Mineralogy: Insights from Modeling Trans-port and Reaction,” Geochim. Cosmochim. Acta 59,1933–1949 (1995).
393. M. K. Tivey and R. E. McDuff, “Mineral Precipitationin the Walls of Black Smoker Chimneys: A QuantitativeModel of Transport and Chemical Reaction,” J. Geo-phys. Res. 95, 12 617–12 637 (1990).
394. M. K. Tivey, L. O. Olson, V. W. Miller, and R. D. Light,“Temperature Measurements during Initiation andGrowth of a Black Smoker Chimney,” Nature 346, 5–54(1990).
395. M. K. Tivey, A. M. Bradley, T. M. Joyce, and D. Kadko,“Insights into Tide-Related Variability at SeafloorHydrothermal Vents from Time-Series TemperatureMeasurements,” Earth Planet. Sci. Lett. 202, 693–707(2002).
396. J. H. Trefry, D. B. Butterfield, S. Metz, et al., “TraceMetals in Hydrothermal Solutions from Cleft Segmenton the Southern Juan de Fuca Ridge,” J. Geophys. Res.99, 4925–4935 (1994).
397. V. Tunnicliffe, M. Botros, M. E. de Burgh, et al.,“Hydrothermal Vents of Explorer Ridge, NorthernPacific,” Deep-Sea Res. 33, 401–412 (1986).
398. A. V. Tutubalin and D. V. Grichuk, “Coupled Hydrody-namic and Thermodynamic Model for a ConvectingHydrothermal System: 1. The Marker Method of Mod-eling,” Geochem. Int., No. 11, 972–984 (1997a).
399. A. V. Tutubalin and D. V. Grichuk, “Coupled Hydrody-namic and Thermodynamic Model for a ConvectingHydrothermal System: 2. Self-Mixing of Solutions,”Geochem. Int., No. 12, 1071–1082 (1997b).
400. A. V. Tutubalin, D. V. Grichuk, and V. I. Mal’kovskii,“A Complex Hydrodynamic and ThermodynamicModel of a Convective Hydrothermal System,” in Pro-ceedings of 8th International Symposium on Water–RockInteraction (Balkema, Rotterdam, 1995), pp. 763–766.
401. D. A. Vanko, “Temperature, Pressure, and Compositionof Hydrothermal Fluids, with Their Bearing on theMagnitude of Tectonic Uplift at Mid-Oceanic Ridges,Inferred from Fluid Inclusions in Oceanic Layer 3,”J. Geophys. Res. 93, 4595–4611 (1988).
402. L. N. Var’yash, “Experimental Study of Cu(I) ComplexFormation in NaCl Solutions at 300 and 350°C,”Geokhimiya, No. 8, 1166–1174 (1991).
403. A. P. Vinogradov, L. N. Grinenko, V. A. Grinenko, andYu. M. Stolyarov, “Isotopic Composition of Sulfur inSulfides and Sulfates from the Alekseevskii CopperDeposit (Middle Urals) and Some Problems of Its Gen-esis,” Geokhimiya, No. 8, 915–926 (1968).
404. V. P. Volkov and G. I. Ruzaikin, “On Gas Equilibria andMethods of Their Computation with Application toPetrogenetic Problems,” Geokhimiya, No. 8, 976–991(1969).
S324
GEOCHEMISTRY INTERNATIONAL Vol. 42 Suppl. 2 2004
GRICHUK
405. K. L. Von Damm, “Seafloor Hydrothermal Activity:Black Smoker Chemistry and Chimneys,” Annu. Rev.Earth Planet. Sci. 18, 173–204 (1990).
406. K. L. Von Damm, “Chemistry of Hydrothermal VentFluids from 9°–10° N, East Pacific Rise: ‘Time Zero’,the Immediate Posteruptive Period,” J. Geophys. Res.105, 11 203–11 222 (2000).
407. K. L. Von Damm and J. L. Bischoff, “Chemistry ofHydrothermal Solutions from the Southern Juan deFuca Ridge,” J. Geophys. Res. 92, 11 334–11 346(1987).
408. K. L. Von Damm, J. M. Edmond, B. Grant, et al.,“Chemistry of Submarine Hydrothermal Solutions at21° N, East Pacific Rise,” Geochim. Cosmochim. Acta49, 2197–2220 (1985).
409. K. L. Von Damm, S. E. Oosting, R. Kozlowski, et al.,“Evolution of East Pacific Rise Hydrothermal Vent Flu-ids Following a Volcanic Eruption,” Nature 375, 47–50(1995).
410. K. L. Von Damm, G. Buttermore, S. E. Oosting, et al.,“Direct Observation of the Evolution of a Seafloor‘Black Smoker’ from Vapor to Brine,” Earth Planet. Sci.Lett. 149, 101–111 (1997).
411. K. L. Von Damm, A. M. Bray, G. Buttermore, andS. E. Oosting, “The Geochemical Controls on Vent Flu-ids from the Lacky Strike Vent Field, Mid-AtlanticRidge,” Earth Planet. Sci. Lett. 160, 521–536 (1998).
412. K. L. Von Damm, M. D. Lilley, W. C. Shanks III, et al.,“Extraordinary Phase Separation and Segregation inVent Fluids from the Southern East Pacific Rise,” EarthPlanet. Sci. Lett. 206, 365–378 (2003).
413. P. J. Wallace and I. S. E. Carmichael, “S Speciation inSubmarine Basaltic Glasses as Determinated by Mea-surements of SKα X-ray Wavelength Shifts,” Am. Min-eral. 79, 161–167 (1994).
414. J. V. Walther and B. J. Wood, “Mineral–Fluid ReactionRates,” in Fluid–Rock Interactions during Metamor-phism, Vol. 5 of Advances in Physical Geochemistry,Ed. by J. V. Walther and B. J. Wood (Springer-Verlag,New York, 1986), pp. 194–213.
415. J. A. Welhan and H. Craig, “Methane, Hydrogen, andHelium in Hydrothermal Fluids at 21° N on the EastPacific Rise,” in Hydrothermal Processes at SeafloorSpreading Centers, Ed. by P. A. Rona, et al. (Springer-Verlag, New York, 1983), pp. 391–409.
416. J. T. Wells and M. S. Ghiorso, “Coupled Fluid Flow andReaction in Mid-Ocean Ridge Hydrothermal Systems:The Behavior of Silica,” Geochim. Cosmochim. Acta55, 2467–2481 (1991).
417. W. S. D. Wilcock and J. R. Delaney, “Mid-Ocean RidgeSulfide Deposits: Evidence for Heat Extraction fromMagma Chambers Or Cracking Fronts?,” Earth Planet.Sci. Lett. 145, 49–64 (1996).
418. T. J. Wolery and N. H. Sleep, “Hydrothermal Circula-tion and Geochemical Flux at Mid-Ocean Ridges,”J. Geol. 84, 249–275 (1976).
419. B. J. Wood and J. V. Walther, “Rates of HydrothermalReactions,” Science 222, 413–415 (1983).
420. L. G. Woodruff and W. C. Shanks, III, “Sulfur IsotopeStudy of Chimney Minerals and Vent Fluids from 21° N,East Pacific Rise: Hydrothermal Sulfur Sources andDisequilibrium Sulfate Reduction,” J. Geophys. Res.93, 4562–4572 (1988).
421. A. A. Yaroshevsky and T. I. Tsekhonya, “PetrochemicalTypes of Ocean Floor Magmatic Rocks: Pecularities inTheir Associations and Distribution in StructuralZones,” in Magmatism in the Ocean: Evolution andGeologic Correlation (Nauka, Moscow, 1986), pp. 95–103 [in Russian].
422. G. P. Zaraiskii, V. N. Balashov, and M. I. Lebedeva,“Macrokinetic Model of Metasomatic Zoning,”Geokhimiya, No. 10, 1386–1395 (1989).
423. A. N. Zavaritskii, “On the Genesis of VolcanogenicMassive Deposits,” Izv. Akad. Nauk SSSR, Ser. Geol.,No. 3, 3–18 (1943).
424. V. V. Zaikov, Volcanism and Sulfide Mounds of Paleo-oceanic Margins: Evidence from Massive SulfideDeposits of the Urals and Siberia (Nauka, Moscow,1991) [in Russian].
425. R. A. Zierenberg, W. C. Shanks, III, and J. L. Bischoff,“Massive Sulfide Deposits at 21° N, East Pacific Rise:Chemical Composition, Stable Isotopes, and PhaseEquilibria,” Geol. Soc. Am. Bull. 95, 922–929 (1984).
426. R. A. Zierenberg, R. A. Koski, J. L. Morton, et al.,“Genesis of Massive Sulfide Deposits on a Sediment-Covered Spreading Center, Escanaba Trough, SouthernGorda Ridge,” Econ. Geol. 88, 2069–2098 (1993).
427. A. Zindler and S. Hart, “Helium: Problematic Primor-dial Signals,” Earth Planet. Sci. Lett. 79, 1–8 (1986).
428. A. G. Zlotnik-Khotkevitch, “Diagenetic Transforma-tion of Massive Sulphide Ores,” Geol. Rudn. Mestor-ozhd. 34 (2), 83–98 (1992).
429. A. G. Zlotnik-Khotkevitch and N. A. Andriyanova,“Formation of Mineral Zoning in Volcanogenic MassiveSulfide Ores,” Izv. Akad. Nauk SSSR, Ser. Geol., No. 8,125–132 (1987).
