ELSEVIER Earth and Planetary Science Letters 163 (1998) 361–379 Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation Max W. Schmidt a,L , Stefano Poli b a CNRS–UMR 6524, Magmas et Volcans, 5, rue Kessler, 63038 Clermont-Ferrand, France b Dip. Scienze della Terra, Sez. Mineralogia, Via Boticelli 23, 20133 Milano, Italy Received 20 February 1997; revised version received 11 May 1998; accepted 4 June 1998 Abstract Phase diagrams of hydrous mid-ocean ridge (MOR) basalts to 330 km depth and of hydrous peridotites to 250 km depth are compiled for conditions characteristic for subduction zones. A synthesis of our experimentally determined phase relations of chlorite, lawsonite, epidote-zoisite, amphibole, paragonite, chloritoid, talc, and phengite in basalts and of phase relations from the literature of serpentine, talc, chlorite, amphibole, and phase A in ultramafics permits calculation of H 2 O contents in hydrous phase assemblages that occur in natural compositions. This yields the information necessary to calculate water budgets for descending slabs. Starting from low-grade blueschist conditions (10–20 km depth) with H 2 O contents between 5 and 6 wt% for hydrated oceanic crust, complete dehydration is achieved between 70 and >300 km depth as a function of individual slab geotherms. Hydrous phases which decompose at depth below volcanic arcs are lawsonite, zoisite, chloritoid, and talc (š phengite) in mafic compositions and chlorite and serpentine in peridotite. Approximately 15–35% of the initially subducted H 2 O are released below volcanic arcs. The contribution of amphibole dehydration to the water budget is small (5–20%) and occurs at relatively shallow depth (65–90 km). In any predicted thermal structure, dehydration is a combination of a stepwise and a continuous process through many different reactions which occur simultaneously in the different portions of the descending slab. Such a dehydration characteristic is incompatible with ‘single phase dehydration models’ which focus fluid flow through a unique major dehydration event in order to explain volcanic fronts. As a consequence of continuously progressing dehydration, water ascending from the slab will be generally available to depth of ca. 150–200 km. The fluid rising from the subducting lithosphere will cause partial melting in the hot portion of the mantle wedge. We propose that the volcanic front simply forms above the mantle wedge isotherm where the extent of melting is sufficient to allow for the mechanical extraction of parental arc magmas. Thermal models show that such an isotherm (ca. 1300ºC) locates below volcanic fronts, slab surface depths below such an isotherm are compatible with the observed depths of the slab surface below volcanic fronts. 1998 Elsevier Science B.V. All rights reserved. Keywords: mid-ocean ridge basalts; magmas; island arcs; dehydration; slabs L Corresponding author. Fax: C33 47334 6744; E-mail: [email protected]0012-821X/98/$ – see front matter 1998 Elsevier Science B.V. All rights reserved. PII:S0012-821X(98)00142-3
Transcript
ELSEVIER Earth and Planetary Science Letters 163 (1998) 361–379
Experimentally based water budgets for dehydrating slabs andconsequences for arc magma generation
Max W. Schmidt a,Ł, Stefano Poli b
a CNRS–UMR 6524, Magmas et Volcans, 5, rue Kessler, 63038 Clermont-Ferrand, Franceb Dip. Scienze della Terra, Sez. Mineralogia, Via Boticelli 23, 20133 Milano, Italy
Received 20 February 1997; revised version received 11 May 1998; accepted 4 June 1998
Abstract
Phase diagrams of hydrous mid-ocean ridge (MOR) basalts to 330 km depth and of hydrous peridotites to 250 kmdepth are compiled for conditions characteristic for subduction zones. A synthesis of our experimentally determinedphase relations of chlorite, lawsonite, epidote-zoisite, amphibole, paragonite, chloritoid, talc, and phengite in basaltsand of phase relations from the literature of serpentine, talc, chlorite, amphibole, and phase A in ultramafics permitscalculation of H2O contents in hydrous phase assemblages that occur in natural compositions. This yields the informationnecessary to calculate water budgets for descending slabs. Starting from low-grade blueschist conditions (10–20 kmdepth) with H2O contents between 5 and 6 wt% for hydrated oceanic crust, complete dehydration is achieved between70 and >300 km depth as a function of individual slab geotherms. Hydrous phases which decompose at depth belowvolcanic arcs are lawsonite, zoisite, chloritoid, and talc (š phengite) in mafic compositions and chlorite and serpentine inperidotite. Approximately 15–35% of the initially subducted H2O are released below volcanic arcs. The contribution ofamphibole dehydration to the water budget is small (5–20%) and occurs at relatively shallow depth (65–90 km). In anypredicted thermal structure, dehydration is a combination of a stepwise and a continuous process through many differentreactions which occur simultaneously in the different portions of the descending slab. Such a dehydration characteristic isincompatible with ‘single phase dehydration models’ which focus fluid flow through a unique major dehydration event inorder to explain volcanic fronts. As a consequence of continuously progressing dehydration, water ascending from the slabwill be generally available to depth of ca. 150–200 km. The fluid rising from the subducting lithosphere will cause partialmelting in the hot portion of the mantle wedge. We propose that the volcanic front simply forms above the mantle wedgeisotherm where the extent of melting is sufficient to allow for the mechanical extraction of parental arc magmas. Thermalmodels show that such an isotherm (ca. 1300ºC) locates below volcanic fronts, slab surface depths below such an isothermare compatible with the observed depths of the slab surface below volcanic fronts. 1998 Elsevier Science B.V. Allrights reserved.
Keywords: mid-ocean ridge basalts; magmas; island arcs; dehydration; slabs
0012-821X/98/$ – see front matter 1998 Elsevier Science B.V. All rights reserved.PII: S 0 0 1 2 - 8 2 1 X ( 9 8 ) 0 0 1 4 2 - 3
362 M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379
1. Introduction
Volcanic arcs are the most spectacular testimonyof destructive plate boundaries on the Earth’s sur-face. Volcanic fronts, i.e. the onset of arc volcanism,form in correspondence to a depth of the subductingslab between 90 and 150 km, with a distributionmaximum of 128 š 38 km [1] or 110 š 38 km (2¦ )[2]. The width of volcanic arcs varies between 50and 200 km.
It is widely accepted that arc volcanism is ini-tiated through partial melting of a mantle wedgemetasomatized by a fluid originating from dehydra-tion of subducted crust [3–6]. The fundamental roleof H2O for island arc magmatism is also documentedby an average H2O content of 1.7 wt% in primi-tive island arc magmas contrasting mid-ocean ridgesettings where primitive magmas have 0.1–0.5 wt%H2O [7].
Recent efforts in determining the stability of hy-drous phases at high pressures allow us to quantifythe progressive dehydration of subducting slabs. Thepressure–temperature stabilities of all major hydrousphases in the principal subducted lithologies (chlo-rite, lawsonite, epidote=zoisite, amphibole, chlori-toid, talc, and phengite in gabbros, basalts, and sedi-ments; serpentine, chlorite, amphibole, phase A, andtalc in mantle compositions) are now experimentallydetermined. Mass balance calculations yield the rela-tive contributions of individual hydrous phases to thetransport of H2O within the subducted lithosphereand between the subducted lithosphere and the over-lying mantle wedge. The aim of this paper is tocharacterize and quantify the dehydration process.We present our experimental data on MOR basalt,discuss H2O-saturated phase diagrams of basalt andperidotite, and compile P–T diagrams of calculatedwt% (H2O) bound in hydrous phases. On the basis ofthese data we draw implications and discuss modelsfor the formation of the volcanic front.
2. Experimental procedure
Three different basaltic compositions (Table 1)were investigated, all of them being mixtures ofsynthetic glasses with Fe3C=Fetot ratios of 0.2–0.25and synthetic and natural crystal seeds [8,9]. LTBC
The first three compositions are basaltic starting materials, theothers averages for MOR basalt, harzburgite (hzb), lherzolite(lherz) [57] and pyrolite (pyrol) [58].a Total iron as FeO.
is close to a Na2O–CaO–FeO–MgO–Al2O3–SiO2–H2O system, KMB contains additional K2O, TBadditional TiO2. Experiments contained 5 wt% H2Oat pressures to 40 kbar and 2 or 1 wt% H2O at higherpressures. All experiments were fluid-saturated andin piston cylinder experiments oxygen fugacity wasbuffered to NNO conditions by a double capsuletechnique.
A total of 63 experiments were performed at pres-sures between 18 and 120 kbar and temperaturesbetween 550 and 1100ºC. Run times varied from 629h (18 GPa, 550ºC) to 30 h (above 800ºC, 50 kbar;run table in EPSL Online Background Dataset 1).Experiments to 36 kbar were performed in an end-loaded piston cylinder apparatus with a half-inchbore. Full salt assemblies with a tapered graphiteheater were employed [8,10]. Experiments above 40kbar were performed in a split sphere multi-anvilapparatus (MA-8) at the Bayerisches Geoinstitut. An18-mm edge length octahedra (18M) with a steppedgraphite furnace was employed on an 11-mm trun-cation edge length (TEL) for experiments to 100kbar and a 14M octahedra with a stepped LaCrO3
furnace was employed on an 8-mm TEL for ex-periments above 100 kbar (for details see [11]).Thermocouples were of Pt–Pt90Rh10 and situated indirect contact to the capsules of 1.5–2.0 mm ini-
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tial length. Calibrations were against quartz–coesite,fayalite–spinel, and coesite–stishovite for the 18Massemblage [11] and against coesite–stishovite andMg2SiO4 forsterite-β-phase at 1200ºC for the 14Massemblage.
3. MOR basalt
3.1. Experimental results
Experiments between 18 and 120 kbar and 550and 1100ºC of this study and experiments of previous
Fig. 1. Experimentally determined phase relationships for water-saturated MOR basalt: amph D amphibole, chl D chlorite, cld Dchloritoid, cpx D jadeitic or omphacitic clinopyroxene, epi D epidote, gar D garnet, law D lawsonite, zo D zoisite.
work [8,10] are compiled in Fig. 1. All experimentsof this study contain garnet and omphacite, thus rep-resent eclogitic assemblages. Several hydrous phasesoccur at pressures above the amphibole stability:lawsonite, zoisite, chloritoid, and occasionally talcor staurolite. In potassium-bearing basalt phengiteappears ubiquitous at subsolidus conditions.
Most experiments represented in Fig. 1 were con-ducted with two or three different basaltic com-positions (TB, KMB, and LTBC). Resulting phaseassemblages for the three bulk compositions wereidentical with the following exceptions. The potas-sium-bearing basalt KMB has an additional potas-
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sium phase (phengite or K-hollandite) at subsolidusconditions. The high temperature stability of chlo-ritoid differs for about 20–40ºC for the differentbasalts. The appearance of some minor phases dif-fers slightly between the different basalts: in LTBCtalc appears in a few experiments (32–26 kbar at650ºC); staurolite occurs in some experiments (pref-erentially in TB) between 25 and 30 kbar, 650–700ºC; and at conditions above 30 kbar, 700ºC mi-nor kyanite occurs in some experiments (LTBC andKMB).
Clinopyroxene forms a matrix of small (3–15µm) euhedral crystals. Jadeite and diopside seedsdisappeared completely and clinopyroxenes werein each experiment. Jadeite fractions are mostlybetween 0.40 and 0.50 but increase to 0.60–0.65above 90 kbar when garnets become majoritic (Ta-ble 2). At about 40 kbar, K2O contents in clinopy-roxene (in KMB) become significant and increasesystematically with pressure to a maximum of
Table 2Average phase compositions
Clinopyroxene Garnet Law Zoisite Cld Stau Phengite
Fe3C calculated from charge balance for lawsonite, chloritoid, and staurolite. Staurolite normalized for 15 cations and 24 oxygensincluding 1(OH) group. Mica formulas on 11 oxygens and 2(OH), all Fe2C. Abbrevations as in Fig. 1.
1.2 wt% at 100 kbar (Fig. 2). K2O contents ofclinopyroxene increase with temperature until thesolidus is reached but at supersolidus conditions theKAlSi2O6-component decreases drastically.
Garnets form texturally equilibrated crystals butoften preserve cores which are intermediate betweenseed and rim compositions. In most experimentsrims around grossular-, pyrope- or almandine-richcores are identical in composition, but in some runsof shorter duration equilibration of rims was incom-plete. Between 25 and 90 kbar the pyrope componentin garnet increases slightly. At about 90 kbar garnetsbecome majoritic and silica contents increase to ca.3.15 Si per formula unit (pfu) at 110 kbar. Simulta-neously, the grossular component decreases by about15 mol% (e.g. in KMB from gros35 to gros20).
Lawsonite crystallizes anew and forms rectangu-lar-shaped euhedral crystals of 7–20 µm in size.Lawsonite compositions are close to endmembercompositions, impurities are 0.3–0.8 wt% FeO, 0.2–
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Fig. 2. Compositions of phengite (in Si per formula unit, givenonly in the phengite stability field) and clinopyroxene (in wt%K2O) in MOR basalt.
0.5 wt% MgO, and maximal 0.6 wt% TiO2. Zoisiteis present in all runs below 32 kbar which do notcontain lawsonite. Zoisite forms large (15–30 µm)crystals which rarely preserve almost Fe-free cores(present in the starting material). Equilibrium com-positions are constant at 0.14–0.18 Fe pfu (Table 2).
Phengite is present in all subsolidus experiments(bulk composition KMB) to 95 kbar. In most experi-ments, phengite tablets are larger than 15 µm whichallows measurement with defocused beam. Gener-ally, the celadonite component increases with pres-sure and decreases slightly with temperature (Fig. 2).Around 90–95 kbar, phengite reaches an endmemberceladonite composition and decomposes with furtherincrease of pressure. In experiments above 50 kbarwith 2–3 wt% of free fluid phase, an enrichment ofpotassium along grain boundaries could be detectedin backscatter images. However, the resolution of themicroprobe did not permit to identify the origin ofthis enrichment. K-hollandite which replaces phen-gite towards higher pressure was extremely rare inthe experimental charges. In most experiments onlya few rectangular-shaped grains of 5 to 20 µm with
almost ideal compositions (KAlSi3O8) have beenfound.
Chloritoid appears as aggregates of 10–40 µm insize. An increase of magnesium number with tem-perature is observed. At pressures between 26 and30 kbar and at temperatures above the chloritoidstability (see Fig. 1), staurolite occurs occasion-ally (mostly bulk composition TB). Staurolites formsmall euhedral crystals of 5–10 µm size. Talc oc-curs in some experiments with LTBC composition. Itforms relatively large crystals of 10–20 µm.
Interstitial liquid forms in small amounts (<5–10%) above the solidus; compositions could notbe measured. When liquid forms, phengite andquartz=coesite=stishovite disappear. In the scanningelectron microscope it appears that quenched liquidscontain frequently small (<2 µm) euhedral garnetsand sometimes liquid is partly replaced by quenchcrystals of submicron size.
3.2. Petrological evolution of subducting basalt
Phase relationships of MOR basalt at water-sat-urated conditions are compiled in Figs. 1 and 3.For the low-temperature region (�550ºC) assem-blages from natural blueschists are employed. Thesubsolidus region below 8 kbar was experimentallyinvestigated by Apted and Liou [12] and from 8 to24 kbar by Poli [8]. The experiments by Pawley andHolloway [13] at pressures from 20 to 30 kbar areconsistent with our experiments.
During subduction, the oceanic crust passesthrough a blueschist facies stage (10–50 km depth).Natural blueschists of basaltic origin are mostlyfully hydrated while coarse-grained gabbros [14]often show domainal equilibria resulting from un-even water distribution. A general sequence ofhydrous phases in basalt with increasing pres-sure might be described as follows. At blueschistconditions, assemblages are typically composedof lawsonite–glaucophane–chlorite–albite=jadeite šphengite (fields G and K, Fig. 3). With increas-ing temperature lawsonite reacts to epidote=zoisite(fields F, I, and J), chlorite decomposes formingmainly garnet (fields E and H), and glaucophanechanges composition to barrositic amphibole. Attemperatures above 600ºC, pressures from 15 to 22kbar, an amphibole–eclogite results (field E). When
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Fig. 3. Maximum H2O contents bound in hydrous phases inH2O-saturated MOR basalt. For assemblages and references seeTable 3.
pressure oversteps 22–24 kbar, amphiboles decom-pose forming chloritoid at temperatures below 660ºC(fields B and D). Blueschists transform to lawsonite-eclogites (fields A and B) and amphibole-eclogitesto zoisite-eclogite (fields C and D). At ca. 30 kbarzoisite reaches its maximum stability in metabasalts(at fO2 D NNO), its breakdown leaves an almost dryeclogite at conditions above 30 kbar, 700ºC (fieldO). At lower temperatures, lawsonite persists reach-ing its maximum temperature stability at 830ºC, 84kbar (field A). If potassium is present in a subductedMOR basalt, phengite forms ubiquitously to pres-sures of 100 kbar [9]. The presence of phengite alsodetermines minimum melting conditions of MORbasalts at high pressures.
Other minor hydrous phases in MOR basalt areparagonite, talc, and staurolite. Paragonite forms at12–16 kbar (at 500–650ºC) and decomposes around22 kbar (500–700ºC, field E and part of field F).The conditions of talc and staurolite occurrence arerelatively restricted, although talc is probably animportant high-pressure phase in Mg-gabbros.
Most of the reactions involved are discussed by[8,10]. We complete the discussion of phase relationswith the lawsonite and phengite breakdown reactionsat high pressures. At pressures above the zoisitestability, lawsonite decomposes through a continuousreaction of the type
lawsoniteC diopsideC garnet1
D garnet2 C coesite=stishoviteC H2O (1)
where garnet2 is richer in grossular and pyrope com-ponent than garnet1. In the coesite stability field,this reaction has a positive 1V and a positivedP=dT slope which steepens with increasing pres-sure (Fig. 1). When stishovite replaces coesite, 1Vof reaction becomes negative. Thus, above 80 kbar,reaction (1) has a negative dP=dT slope and thuslimits the maximum pressure stability of lawsonitein subduction zones to 80–90 kbar.
White mica, the principal potassium host in nat-ural blueschists and eclogites, decomposes througha continuous and a discontinuous reaction. At pres-sures above 40 kbar, a significant solubility of potas-sium in clinopyroxene is observed and phengite de-composes forming KAlSi2O6–cpx C enstatite C co-esiteC K-rich fluid [9]. This reaction is continuouslyprogressing from 40 kbar until the terminal phengitestability near 100 kbar is reached. The discontinuousbreakdown of phengite to K-hollandite (KAlSi3O8)C clinopyroxene C K-rich fluid and the related de-hydration might be of small significance in basalticcompositions. Mass balance calculations on the ba-sis of the experiments reveal that in MOR basalts(0.1–0.5 wt% K2O) phengite may react completelyto K–cpx before the pressure stability of phengiteis reached (Fig. 4). With a potassium solubility inclinopyroxene of about 1 wt% (at 100 kbar) anda modal amount of 40–50 wt% cpx, 0.4–0.5 wt%bulk-K2O can be stored in clinopyroxene.
3.3. Mass balance
Modal amounts of phases in basalt (Table 3)are compiled from natural blueschists with basalticchemistry [15–19], from [12] and from our ownexperiments. For a description of modal amount cal-culations in basalt see [8]. Fig. 3 gives wt% of H2Oin the bulk rock bound in hydrous phases. These
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Fig. 4. Modal amounts of phases in MOR basalt with 0.5 wt%K2O along a typical subduction geotherm. Phengite continu-ously decomposes when potassium solubility in clinopyroxeneincreases.
values are for potassium-free basalt. For potassium-bearing basalt 0.5 wt% H2O (equivalent to 8–9 wt%phengite) must be added per 1 wt% K2O. Our resultsare comparable to those of [20] which estimated wa-ter contents for subducting oceanic crust to 40 kmdepth.
As can be seen from Fig. 3, many different re-actions involving hydrous phases modify the watercontents during the transition from blueschists to dryeclogites. Water contents will continuously decreasebecause most of the reactions are continuous reactionsinvolving solid solutions and thus take place over a
Table 3Assemblages, modal amounts of hydrous phases (wt%) and H2O (wt%) bound in hydrous phases in MOR basalt
Field in Fig. 3 Assemblage Ranges for modal amounts of hydrous phases Range of H2Otot References
a [10]: references in Ref. [10], Ref. [19] is from a coesite bearing gabbro.
temperature=pressure interval [21]. In an experimen-tal study [8] it was shown that continuous amphiboledecomposition in amphibole eclogite (field H and Ein Fig. 3) causes as much dehydration as the finalamphibole breakdown (each1H2Otot D 0.6 wt%).
The phases with large contributions to the wa-ter budget of subducting MOR basalt are lawsonite,chlorite, and amphibole (glaucophane or barroisitein composition). Lawsonite contains 11.2 wt% H2Oand typically amounts to 7–15 wt%. Chlorite inbasalt contains about 12 wt% H2O and amounts to5–30 wt%. Amphibole only contains 2.1 wt% H2Obut forms 20–60 wt% of a basalt. Other signifi-cant contributions arise from epidote=zoisite (1.95wt% H2O) which forms typically 5–20 wt%, andfrom chloritoid (7.5 wt% H2O) amounting to 3–8 wt%. Paragonite (4.6 wt% H2O), talc (4.8 wt%H2O), and staurolite (1.0 wt% H2O) have only minorcontributions due to medium–low water contents incombination with relatively low modal abundances.
The following dehydration characteristics can bedefined. In the shallow portion of the subductingslab, basalts typically contain 5–6 wt% H2O. Withincreasing pressure (but still in the stability fieldof amphibole), water contents decrease to ca. 3wt% at low temperatures where lawsonite C chlo-rite C glaucophane is stable, to 2.0–2.5 wt% atintermediate temperatures where epidote C chloriteC glaucophane=barroisite is stable, and to 1–1.5wt% where epidote=zoisite C barroisite is stable.
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The most significant discontinuous reactions in theamphibole stability field are the formation of om-phacitic clinopyroxene and the breakdown of chlo-rite. At pressures exceeding the stability of amphi-bole (>22–24 kbar), ca. 1 wt% H2O remain boundin lawsonite and chloritoid (<520ºC) and 0.5–0.7wt% H2O bound in zoisite š chloritoid. At higherpressures (>30–40 kbar) when lawsonite constitutesthe only non-potassic hydrous phase, the bulk watercontent is typically 0.2–0.5 wt% H2O.
Apart from the continuity of several reactions,temperature gradients and compositional variationsin the descending oceanic crust will further broadenthe dehydration signal. Temperature gradients withina 7–10 km thick oceanic crust are 100–300ºC [22].Consequently, several dehydration reactions maytake place at the same depth. Strictly, these cal-culations and phase relationships are only valid forMOR basaltic compositions. However, the stabilitiesof lawsonite, epidote=zoisite, amphibole, and chlo-rite are relatively insensitive to bulk rock compo-sition [10]. On the contrary, ferromagnesian phasessuch as chloritoid and talc will have significantly en-larged stability fields and abundances in Mg-gabbrosas can be deduced from field occurrences ([23] andreferences therein).
4. Metapelites and metagreywackes
Metapelites and metagreywackes constitute on av-erage the upper 10% of the oceanic crust. Althoughthey contribute only a small portion to the water bud-get, they are important for hosting the vast amountof certain elements or isotopes characteristic for thesediment-signature in arc magmas (K, Rb, Ba, Sr,10Be, enrichment in 207Pb [3,5,6]). Phase relationsfor sedimentary materials metamorphosed at subduc-tion zone conditions were experimentally studied by[24–26].
Chlorite C kyanite and staurolite C quartzare limited to pressures below 20 kbar formost natural bulk compositions, the latter hav-ing Mg=(Mg C Fe) ratios of 0.3–0.6. At pres-sures above 20 kbar, hydrous assemblages com-prise talc–chloritoid–phengite or talc–phlogopite–phengite. Between 40 and 50 kbar talc C kyan-ite destabilizes and Mg–Al–pumpellyite (7.1 wt%
H2O) and=or topaz–OH (10.7 wt% H2O) might formin the presence of phengite. Metagreywackes arecommonly richer in Ca than metapelites and thuslawsonite or zoisite are more abundant. The stabilityfield of lawsonite in metapelite [26] is very similarto that in mafic systems. H2O storage in sedimentsto more than 200 km depth is ensured by phen-gite, lawsonite, MgAl–pumpellyite, and topaz–OHat temperatures as high as 900ºC.
5. Peridotites
The principal hydrous phases in H2O-saturatedperidotite to 80 kbar are serpentine, phase A, chlo-rite, talc, and amphibole. Stabilities of serpentine,phase A, and talc are expected to be almost iden-tical for depleted or fertile peridotite because thesephases are almost Ca- and Na-free and contain onlyminor amounts of Al2O3. Their stabilities have beenexperimentally determined on natural mineral com-positions [27,28]. The stabilities of chlorite and am-phibole are expected to vary for harzburgitic, lher-zolitic, and pyrolitic bulk compositions. Mysen andBoettcher [29] determined a stability increase of am-phibole from ca. 22 kbar in harzburgite to ca. 25 kbarin enriched lherzolite (at 800–1000ºC).
Serpentine contains 12.3 wt% H2O and domi-nates, together with chlorite (13.0 wt% H2O), thewater budget of hydrous peridotite to ca. 150 kmdepth. In natural peridotites serpentine forms duringlow-grade hydration, and its maximum temperaturestability is 720ºC (21 kbar, invariant point 3 inFig. 5). Serpentine (antigorite, Mg48Si34O85(OH)62)decomposes to talc C olivine C H2O below 21 kbar[27], to orthopyroxeneolivine C H2O from 21 to 62kbar, and to phase A C orthopyroxene C H2O [28]at higher pressures (invariant point 4 in Fig. 5).
In average mantle compositions, talc (4.7 wt%H2O) C olivine have a rather limited stability fieldless than 100ºC wide, and they decompose to en-statite C H2O between 690ºC (10 kbar) and 720ºC(20 kbar) [28]. In relatively silica-rich veins, talcmight locally persist to higher temperatures and pres-sures [30,31].
In subducted peridotite, phase A(Mg7Si2O8(OH)6, 11.8 wt% H2O) replaces serpen-tine at pressures between 60 and 70 kbar through
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Fig. 5. Phase diagram for H2O-saturated average mantle peridotite and maximum H2O contents bound in hydrous phases in averageperidotites. Upper value: harzburgite; middle value: lherzolite; lower value: pyrolite. The italic labels are assemblages in a given stabilityfield. ‘A’ D phase A, amph D amphibole, chl D chlorite, cpx D clinopyroxene, gar D garnet, ol D olivine, opx D orthopyroxene, serp Dserpentine, sp D spinel, tc D talc.
a water-conserving reaction (dotted line emergingfrom invariant point 4 in Fig. 5). Although phase Acould form at lower pressures through olivineC H2OD phase A C serpentine, the free water necessary forthis reaction would not be available in subductingperidotite. Above 62 kbar, phase A C orthopyroxenedecompose with increasing temperature to enstatiteC H2O. This reaction has a moderate positive slopein P–T space [32].
In natural peridotites, chlorites have compositionsclose to clinochlor (Mg5Al2Si3O10(OH)8). Syntheticclinochlor decomposes between 3 and 21 kbar toolivine C orthopyroxene C spinel C H2O with amaximum temperature stability of 870ºC ([31] andreferences therein). Above 21 kbar, the occurrenceof chlorite in peridotite is limited through the reac-tion chlorite C orthopyroxene D olivine C garnet CH2O. This reaction was experimentally determinedto pressures of 40 kbar [33]. Although considerable
uncertainty remains, it can be deduced from experi-ments on the terminal stability of chlorite [34] thatthe stability of chlorite C orthopyroxene do not ex-ceed 50 kbar. The stability of chlorite outlined inFig. 5 represents probably a maximum, and Fe coulddecrease the stability of chlorite in natural peridotitesignificantly.
5.1. Amphibole in peridotite
A compilation of all experimental studies on am-phibole in water-saturated peridotite compositionsyields the following picture. Close to the water-satu-rated solidus (approximately 1000ºC, at 20–30 kbar),amphibole is pargasitic hornblende in compositionand decomposes between 22 and 30 kbar. In harzbur-gite, calcic amphibole decomposes at 22 kbar, inlherzolite at 25–28 kbar and in enriched pyroliteat 28–30 kbar [29,35–37]. Mysen and Boettcher
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Fig. 6. Result of Schreinemaker analysis for breakdown reactions of amphibole in peridotite. The solid lines are amphibole breakdownreactions in a system where excess water is present, the stippled lines represent amphibole breakdown reactions in a fluid-absent system.The highest pressure where amphibole is stable is located at the solidus. Numbers without parentheses refer to invariant points in Fig. 5.The amphibole breakdown reaction between invariant point 1 and 2 is amph C ol C H2O D tc C cpx C chl. Abbreviations as in Fig. 5.
[29] determined experimentally a slightly positiveP–T slope for the amphibole-out reaction at sub-solidus conditions. Calculations of the P–T slope ofthe amphibole-out reaction (2) of Fig. 6 agree with[29] yielding 3 to �1 bar=K as a function of ex-act amphibole-composition and solid solution model(molar volumes and entropies from 1996 update of[38]). Consequently, the highest pressure stability ofamphibole is located at the intersection of the amphi-bole-out curve with the solidus, i.e. between 22 and30 kbar (65–90 km depth). With decreasing tempera-ture, this pressure is reduced by several kbar (Figs. 5and 6). The experimental evidence is in sharp con-trast to Tatsumi [2] who extrapolates the slope ofMillholen et al.’s [35] amphibole melting curve intothe subsolidus in order to reach 35 kbar at 700ºC.The amphibole-out reaction near the solidus does notshift as a function of the bulk H2O content: as longas the volatile species is only water, reaction (2) ofFig. 6 does not change position in P–T space. Whenthe fluid is diluted by components other than H2O,this reaction is shifted to lower pressures, conse-quently, H2O-saturated bulk compositions yield thehighest amphibole stabilities.
At temperatures where chlorite appears insteadof garnet, reaction (2) transforms into reaction (3)(Fig. 6). Calculated P–T slopes for the latter reactionare positive and range from 11 to 13 bar=K. The
amphibole breakdown reactions (3) and (4), latterin the chlorite C serpentine stability field, do notliberate water. On the contrary, in a water-saturatedsystem, the decomposition of amphibole consumeswater with increasing pressure. In the descendingslab, additional water is not necessarily available andamphibole decomposes through water conservingreactions (5) and (6) (Fig. 6). Thus, it is only attemperatures above the chlorite stability that theamphibole breakdown liberates water.
The exact stoichiometries of the amphibole de-composition reactions are complicated by the changeof amphibole composition with temperature. Fromexperiments and natural xenoliths it is known that attemperatures near the solidus amphibole is pargasitichornblende in composition. In low temperature–highpressure metamorphic peridotites, amphiboles arealuminous tremolites. For the pressure breakdownreactions we thus continuously interpolated betweenthese two amphibole compositions. Fig. 6 is valid foramphibole compositions to be expected in naturalperidotite; however, reaction topologies are differentfor endmembers such as tremolite or edenite.
5.2. Water contents in fully hydrated peridotite
Modal abundances in hydrous peridotite as-semblages are calculated employing a least-square
M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379 371
Table 4Modal amounts (wt%) and maximum H2O contents (wt%) bound in hydrous phases in hydrous peridotites
C Cr2O3, and Na2O C K2O were each re-duced to a single chemical component. Serpen-tine and chlorite compositions were held con-stant at the ideal stoichiometries of antigoriteand clinochlor. This introduces a small error forthe calculated abundances of chlorite. By vary-ing the alumina content of clinochlor for š20%,relative differences in the calculated H2Otot con-tents are �11=C16% for the assemblage chlorite–amphibole–orthopyroxene–olivine and �17=C26%for chlorite–clinopyroxene–orthopyroxene–olivine.Amphibole compositions were varied with tem-perature as described above. The relative amountsof clinopyroxene and amphibole depend on theirNa : Ca ratios. Since compositional data at low tem-peratures are scarce, we allowed amphibole to com-pletely replace clinopyroxene. This results in maxi-mum amphibole and minimum chlorite abundances.
The results of the mass balance calculations foraverage peridotite compositions (Table 1) are pre-sented in Table 4 and Fig. 5. In the stability fieldsof serpentine and phase A the total amount of H2Obound in hydrous phases varies between 3.9 and 10.6wt% of the bulk rock. In the stability field of ser-pentine C chlorite C amphibole, chlorite contributesapproximately 30% (lherzolite) of the H2O boundin hydrous phases, whereas amphibole contributesless than 10% (Table 4). When serpentine reacts totalc, H2O contents decrease from 5.1 to 2.5 wt%in lherzolite. At higher temperatures, when chloriteC amphibole constitute the stable hydrous phases,
H2O contents decrease to 1.1 wt% in lherzolite (0.5wt% in harzburgite). In this assemblage chlorite con-tributes about 60% whereas amphibole contains 40%of the H2O bound in hydrous phases. When chloriteis the only hydrous phase, H2O contents are 2.1 wt%in lherzolite (1.0 wt% in harzburgite). When amphi-bole is the only stable hydrous phase, H2O contentsamount to less than 0.5 wt% in lherzolite (0.2 wt%in harzburgite). In chlorite=amphibole assemblages,water contents are higher in fertile than in depletedperidotite, because higher Na2O, CaO, and Al2O3
contents result in larger amounts of amphibole andchlorite.
Serpentine has a predominant role in the waterbudget of peridotite. When serpentine decomposes,chlorite becomes the dominant hydrous phase. Al-though amphibole contains ca. 40% of the H2Obound in the P–T domain of chlorite C amphibole,water is not liberated when amphibole decomposes.It is only between 800–850ºC and the wet solidus(ca. 1000ºC) that fluids result from the breakdownof amphibole. The quantity of this fluid is smallcompared to the fluid resulting from other phases:in lherzolite, the maximum content of amphiboleamounts to 19 wt% (harzburgite 9 wt%) contribut-ing a maximum of 0.5 wt% (harzburgite 0.2 wt%)H2O. Recently, Niida and Green [37] have shownthat half of the 25% amphibole present in a pyro-lite at 10–15 kbar decompose through a continuousreaction before the pressure stability of amphiboleis reached at 28–30 kbar. Obviously, in a natu-ral mantle, the reactions of Fig. 5 will take place
372 M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379
over a temperature–pressure interval. Over a certaintemperature–pressure range reactants and productswill both be stable and a broad rather than a sharpdehydration signal will result from each particularreaction.
Although phase A has no direct bearing on theformation of arc volcanism, the position of invariantpoint 4 (Fig. 5) is essential for transporting H2O togreat depth. When a portion of hydrated peridotiticslab passes a depth of 180 km at temperatures belowca. 600ºC, phase A forms directly from serpentineand the contained water will not directly recycle intothe mantle wedge.
6. The fate of water released from the subductinglithosphere
Only a portion of the H2O initially containedin the subducting lithosphere will escape from theslab in the depth range suitable for arc magmaformation. A large portion of H2O will be dehydratedat relatively shallow levels and a small portion willfade into the deeper mantle. In the following wecalculate estimates for the different final destinies ofsubducted water.
6.1. Oceanic crust
Blueschists of basaltic origin representing shallowportions of subduction zones (10–20 km) have typi-cally 5–6 wt% H2O bound in hydrous phases. Witha 7 km thick mafic oceanic crust, composed by equalparts of basalt and gabbro, 0.5–0:6ð 109 g H2O m�2
are bound in the basaltic layer 2 at shallow depth.Oceanic gabbros are only partially hydrated; with20–30% hydration by volume we estimate 0.11–0:18 ð 109 g H2O m�2 for the gabbroic layer. Anadditional amount of 0.01–0:02 ð 109 g H2O m�2
will be subducted in the sedimentary layer (200–400m thick). Some trace elements characteristics for arcvolcanism originate mostly from this sedimentarylayer; however, its contribution to the H2O budget islargely within the error of the estimates for the otherlayers of the oceanic lithosphere.
2 In the calculations we use a vertical section through the sub-ducting crust with a unit area of 1 m2, thus this number repre-sents the quantity of water in a column of 1 m2.
In a vertical section of 1 m2, typical dehydra-tion rates between 20 and 70 km depth amountto 0:8 š 0:2 ð 107 g H2O per kilometre of depth.When the oceanic crust reaches depths of 70–80 kmthe resulting lawsonite-eclogites (cold subductionzones) or zoisite-eclogites (relatively warm subduc-tion zones, ‘warm’ being intended as just below thewet solidus of the oceanic crust) retain 1 or 0.5 wt%H2O, i.e. 0.2 or 0:1 ð 109 g H2O m�2, respectively.If lawsonite-eclogites persist to the maximum depthof lawsonite stability (cold subduction zones), 0.02–0:08 ð 109 g H2O m�2 will descend to 260–280 kmdepth and may be lost for arc volcanism (Figs. 7 and8). At ca. 200 km depth, typical dehydration ratesin the stability field of lawsonite are in the order of0.05–0:09 ð 107 g H2O per km depth, i.e. one orderof magnitude lower than at shallow levels.
6.2. Mantle wedge peridotite overlying the oceaniccrust
Here, the only hydrous phase which may persistbelow the depth of arc volcanism is chlorite (Fig. 7).Thermal models indicate that at depth below thevolcanic arc a narrow zone of 1–3 km may containchlorite in cold subduction zones yielding a quantityof 0.06 to 0:16ð109 g H2O m�2. The rest (60–90%)of the previously dehydrated 0.4 to 0:6ð 109 g H2Om�2 from the oceanic crust remains mostly in thecold corner of the mantle wedge and a portion of thisH2O will give raise to serpentine diapirs.
6.3. Subducting peridotite underlying the oceaniccrust
With a maximum amphibole content of 19 wt%in lherzolite the amount of H2O bound in amphi-bole may not exceed 0.5 wt% of the bulk rock.To achieve a similar bulk H2O content due to ser-pentinization, only 3 wt% serpentine are necessary.A serpentinization of 3% causes the density of alherzolite to decrease from 3.30 to 3.28 g=cm3, anegligible effect. Even 10% serpentinization (result-ing in 1.3 wt% H2O in the bulk rock) cause only asmall decrease in density from 3.30 to 3.23 g=cm3.Direct information on the extent of serpentinizationbelow the oceanic crust is not available. Serpentinecertainly forms at transform faults and fracture zones
M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379 373
Fig. 7. Stability of hydrous phases for thermal structures with subduction parameters spanning most of the range of moderncircum-Pacific subduction zones. Amphibole in peridotite is stable within the light-grey-shaded area (note that at temperatures below450–500ºC diopside is stable instead of tremolitic amphibole). Thermal model of [22] (A and B), and of [47] (C). Stippled lines areisotherms and bold arrows indicate flow lines in the mantle wedge. The principal difference between these two models is the much lowercorner flow in (C) [47] which leads to a significantly colder peridotite overlying the oceanic crust. The stability limits of hydrous phasesin the subducting peridotite underlying the oceanic crust do not imply that these hydrous phases are present deep into the subductingperidotite. Abbreviations as in Figs. 1 and 5; labels in capital letters mark stability boundaries in peridotite.
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Fig. 8. Model for the formation of the volcanic front. Dehydration from peridotite and oceanic crust occurs at almost any depth to ca.150–200 km, thus water will be generally available above the subducting lithosphere. The grey region in the mantle wedge will have asignificant amount of melt present. The volcanic front forms where the amount of melt is sufficient to mechanically extract and give riseto arc magmatism. Open arrows indicate rise of fluid, short solid arrows indicate rise of melts. Long arrows indicate flow in the mantlewedge. Stippled lines outline stability fields of hydrous phases in peridotite; however, we do not want to imply that peridotite underlyingthe oceanic crust is fully hydrated. The striped region shows the zone where talc occurs in average peridotite compositions. In somethermal structures (Fig. 7b and c) a portion of the peridotitic lithosphere will be colder than 600ºC at 62 kbar, serpentine will react tophase A and thus H2O will be subducted to large depth. In the oceanic crust, temperatures can be low enough to preserve lawsonite andphengite to their maximum pressure stability; however, at somewhat warmer conditions (slower subduction, shallower angle, youngercrust) zoisite (zo*) will be the last potassium-free phase to decompose and the top of the oceanic crust (phengite rich sedimentary layer)might melt. Single phase transitions which cause a potassium-rich fluid pulse could be constituted by the pressure breakdown of phengitein oceanic crust and by the phlogopite to K-richterite reaction in previously K-metasomatized mantle.
which subduct together with undisturbed ocean floor.It was estimated that about 20% of the Pacific Oceanfloor are effected by fracture zones [39] and alter-ation and serpentinization was found to be extensivein ocean basins [40]. We assume a shallow zone ofca. 5 km depth with an average serpentinization of10% which amounts to 0:21ð109 g H2O m�2. As anupper limit, this value might be doubled, but then theserpentinized peridotite becomes increasingly buoy-ant (² D 3:15 g=cm3 for 20% serpentinization) and itbecomes questionable if such a peridotite might stillbe able to descend. Less than half of the above valueappears to be inconsistent with the degree of serpen-tinization observed at modern seafloors. Obviously,
the hydrated peridotite below the oceanic crust can-not be water-saturated (full serpentinization). Thus,in cold subduction zones where the stability fields ofphase A and serpentine overlap, H2O would entirelydescend to a large depth and will not be availablefor arc magmatism. It is only in intermediate towarm subduction zones where fluid will result fromserpentine dehydration (Fig. 7). Temperatures in theperidotite of the subducting oceanic lithosphere arelower than in the subducting crust, thus, significantdehydration of this portion of the subducting litho-sphere does not take place at shallow levels. Atdepths where the oceanic crust contains 0.5–2.0 wt%H2O, the underlying peridotite might well be a reser-
M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379 375
voir of water with a mass of H2O similar or superiorto the H2O bound in oceanic crust.
6.4. The H2O-budget
The above calculations show that the oceaniclithosphere contains about 0.71–1:08 ð 109 g H2Om�2 at 20 km depth. Both the hydration state of theunmetamorphosed oceanic crust and the early diage-netic processes (e.g. expulsion of pore water [41])in and below the accretionary prism are extremelydifficult to quantify. Thus, we prefer to define the‘input’ into subduction zones through the state ofthe oceanic lithosphere at 20 km depth. Our estimateis higher than the 0:3 ð 109 g H2O m�2 of Pea-cock [42] which did not consider serpentinizationbelow the oceanic crust and based his estimate onthe few DSDP drill holes which perforated into thecrystalline oceanic crust. It was pointed out by Be-bout [41] that this latter procedure may significantlyunderestimate the volatile content of subducted crust.
The summation of the above estimates yield thefollowing picture for the amount of H2O availablefor arc volcanism. In cold subduction zones, 0.12to 0:18 ð 109 g H2O m�2 will rise from the oceaniccrust and 0.06 to 0:16ð109 g H2O m�2 will rise fromthe peridotitic layer above the oceanic crust. Thus,the total amounts to 0.18 to 0:34 ð 109 g H2O m�2
since no dehydration occurs below the oceanic crust.In intermediate to warm subduction zones, about0:1ð 109 g H2O m�2 will rise from the oceanic crustand roughly 0.1 to 0:3 ð 109 g H2O m�2 from theperidotite below the oceanic crust. Thus, the totalamounts to 0.2 to 0:4ð 109 g H2O m�2 since no de-hydration occurs above the oceanic crust. It followsthat between 18 and 37% of the 0.7 to 1:1 ð 109 gH2O m�2 present at 10–20 km depth are available forthe generation of arc magmas. This portion of H2Owill rise into the mantle wedge in the depth intervalbetween 80 and 150 km. An interesting feature of theabove estimate is that there is no principal differencefor the recycling rate of water in the volcanic arcbetween a cold and a warm subduction zone (‘warm’always being intended as temperatures of the oceaniccrust just below its solidus). The above result ishigher than estimates that globally 10% to 20% ofthe subducted water is expelled via arc magmatism[42,43]. However, not all of the water dehydrated
below the arc region will appear at the surface, oneportion of the dehydrated water will remain in themantle wedge.
The above estimates also yield that 30–70% or0.25 to 0:7 ð 109 g H2O m�2 of the initially sub-ducted H2O will be released to the cold corner of themantle wedge where serpentine is stable. Accordingto the thermal models of Furukawa [22] this regionhas cross-sections of ca. 3:1 ð 109 m2 for a 30ºsubduction angle and of 0.7 to 0:9 ð 109 m2 for a60º subduction angle (4–8 cm=a subduction rates).Full serpentinization (and chloritization) of the coldcorner requires about 0.2ð106 g H2O m�3, the bulkrock density would be 2.88 g=cm3. For a stationarycold corner, full serpentinization (and chloritization)would be achieved after subduction of 570–790 kmslab for a 30º subduction angle and 130–220 km slabfor a 60º subduction angle.
7. Consequences of the experimentally basedwater budget for the formation of the volcanicfront
In the following we discuss models for the for-mation of volcanic fronts in the light of our waterbudget. We focus exclusively on arcs where parentmagmas form from melting of the mantle wedge anddo not discuss slab-derived melts. Such direct melt-ing of the oceanic crust (for models see Ref. [44])and subsequent adakitic magmatism may only occurwhen very hot young crust is subducted and is rareon the modern Earth [45].
7.1. ‘Single phase dehydration models’ and the roleof amphibole for arc magmatism
Many models for the formation of arc magma-tism have been proposed in the last two decades.Some of these models involve a multistage processwhich evokes dehydration of the slab at a relativelyshallow depth and intermediate storage of H2O (andtrace elements deriving from the subducted slab) in asingle, key hydrous phase in the mantle wedge. Sub-sequently it is either proposed that the down-draggedmantle wedge dehydrates and melting occurs abovethis dehydration zone or that melting of the man-tle wedge occurs on the wet solidus of peridotite.
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A number of hydrous phases have been called onto explain such transient storage of H2O and ofother ‘metasomatizing components’ within the man-tle wedge [30,46,47].
Tatsumi [2,46] and Davies and Stevenson [47]ascribe the distribution maximum of volcanic frontsto the amphibole breakdown in peridotite. Tatsumiassigned the formation of volcanic front to melt-ing within a hydrous column forming directly aboveamphibole dehydration in peridotite at ca. 110 kmdepth. Davies and Stevenson [47] evoked a mech-anism of lateral transport of water in the mantlewedge. They propose that amphibole dehydrates atca. 90 km depth but the rising fluid (or melts) re-hydrates the convecting mantle wedge and againforms amphibole. Through such a stepwise processwater was thought to be transported laterally withinthe hot wedge. Both these amphibole-based modelsare in disagreement with available experimental andgeochemical constraints.
(1) As previously discussed, the breakdown ofcalcic amphibole in peridotite occurs, as a functionof peridotite composition, between 22 and 30 kbar,i.e. between 65 and 90 km depth. A lateral transportmechanism [47] is only possible for a limited hori-zontal distance because amphibole melts between the1000 and 1100ºC isotherm in the mantle wedge [48]to form high Mg-andesites. Although such magmacompositions occur in arcs, they are not candidatesfor primitive arc magmas. The most primitive mag-mas in arcs are basalts to picritic basalts. For theirgeneration, temperatures of 1300–1350ºC are neces-sary ([49–51], fig. 5.6 in [46] and references therein).Thus, even though this mechanism is possible, am-phibole falls short in transporting laterally water intothe source region of basaltic magmas (Fig. 7).
(2) Even when stable, amphibole contributesmoderately to the water budget. It follows fromthe mass balance calculations that amphibole has arelatively small contribution to the water budget ofperidotite (0.2–0.5 wt% H2O of the bulk rock) withrespect to the other hydrous phases. Furthermore itdecomposes through continuous reactions [37] andits final breakdown depth is extremely sensible tobulk composition.
(3) Gill and Williams [52] described that about60% of island arc magmas have U–Th disequilib-rium due to U-excess. This U-excess is believed
to result from dehydration of the oceanic crust,the preservation of U–Th disequilibrium thus con-strains dehydration of the slab to take place less than350,000 years before eruption of arc magmas. Fordifferent arcs, the time interval between U-additionto the magma and eruption was determined to be lessthan 20,000 to 200,000 years ([4,53] and referencestherein). The additional time required for ‘lateraltransport’ [47] or ‘simple downdrag’ [46] of H2Othrough amphibole can be calculated from subduc-tion geometries and velocity fields in the wedge. Athigh subduction rates of 10 cm=a and maintainingthis high velocity between subducted crust and melt-ing region in the mantle wedge, 460,000 years (60ºangle of subduction) to 700,000 years (30º angle ofsubduction) are necessary for H2O transport throughamphibole to a position below the volcanic front.Because the convection velocity in the mantle wedgeis much slower than the velocity of the subductedlithosphere, realistic time spans for such transportmechanisms would be 1.5 to 3 times longer (velocityfields from [22,47]) and thus would largely exceed350,000 years.
We conclude that the role of amphibole is notoutstanding with respect to other hydrous phases.Amphibole cannot constitute ‘the’ fluid source situ-ated below the volcanic front.
7.2. Towards a ‘continuous dehydration model’
The phase petrological data indicate that somefluid rises from the subducting slab and hydratesa portion of the overlying mantle wedge at almostany depth to ca. 150–200 km. Phase relationshipsin mafic and ultramafic lithologies reveal that, ata given depth, several hydrous phases will decom-pose through either discontinuous or continuous re-actions. Since temperature gradients are high (up to50ºC=km) close to the slab surface, most dehydra-tion reactions will be smeared out over a significantdepth interval instead of taking place at a distinctdepth. Further broadening of the dehydration sig-nal results from (1) local=vertical inhomogeneities inbulk composition cause different solid phase compo-sitions and consequently dehydration reactions takeplace at slightly varying P–T conditions, and (2) un-even H2O distribution in coronitic metagabbros andmetaperidotites. As a consequence, sections through
M.W. Schmidt, S. Poli / Earth and Planetary Science Letters 163 (1998) 361–379 377
natural subducting lithosphere will have an evenmore continuous dehydration signal than indicatedby our investigated basalt–H2O and peridotite–H2Osystems. It should be noted that, independent fromphase equilibria data, several recent geochemicalstudies on modern arcs [6,54,55] equally propose acontinuous dehydration of the subducting lithosphereon the basis of trace element and isotopic data.
Dehydration at low pressures is almost entirelydue to potassium-free phases, whereas at pressuresabove 30 kbar, phengite becomes more and more im-portant. The extensive pressure stability of phengite[9,26] indicates that it is more difficult to recyclepotassium from the slab than descending potassiumto the breakdown depth of phengite (270–300 km).Two continuous processes will allow to recover aportion of the potassium (and related elements). (1)A continuous formation of K-clinopyroxene fromphengite liberates water (and thus hydrophile ele-ments such as potassium). This reaction becomes ef-ficient only above 50 kbar. (2) Potassium loss in theoceanic sediments and basalts could be obtained byleaching of hydrophile elements by fluids rising fromthe lower part of the oceanic crust or the hydratedsubducting peridotite through the sedimentary layer.
Consequently H2O-rich fluids with a variable geo-chemical signal will be generally available abovesubducting lithosphere to at least 150–200 km depth(Fig. 8).
The next question to answer is ‘where does thevolcanic front form?’. Kushiro [49] suggested that asignificant amount of melt forms and collects at thetrench-nearside of the partially molten region anddeduced that the volcanic front simply forms abovethe trench-nearest partially molten mantle wedge.Temperatures for attainment of melting in the mantlewedge below volcanic fronts can be deduced fromprimary arc magmas. Primitive arc magmas from theEarth’s surface and experiments on such magmas(which must be in equilibrium with mantle composi-tions) indicate mantle wedge temperatures of 1300–1350ºC [46,49–51]. Although the quantity of prim-itive magmas in arcs is small, the very existence ofthese magmas shows that such temperatures existbelow the volcanic arc. Thus, most likely the temper-ature necessary for mechanical extraction of parentalarc magmas is achieved at 1300–1350ºC. Thereforethe location of the volcanic front is expected to de-
pend mostly on the position of the 1300ºC isothermand (Fig. 8) on the individual subduction parameterswhich determine the mantle wedge structure.
Unfortunately only few models deal in detail withthe concurrent solution of both thermal and mechan-ical properties of the mantle wedge. The geometryof velocity fields in the wedge and therefore theevolution of the thermal structure close to the ‘magiccorner’ mostly depend on the degree of coupling be-tween the subducting slab and the wedge, on mantlewedge viscosity, and on the thickness of the rigidmechanical lithosphere (compare models in [22,47]and [56]). The attainment of mantle wedge temper-atures of 1300ºC below volcanic arcs are obtainedwhen a vigorous, convective return flow is enhancedeither by decoupling and=or progressive thinning ofthe mechanical lithosphere. A number of other pa-rameters such as convergence rate, dip angle of theslab, age of the slab and initial temperature profileobviously play a major role in determining the mag-nitude of such return flow [56]. Without a significantcorner flow below volcanic fronts mantle wedgetemperatures would rapidly cool down and arc mag-matism would cease. Seismic tomography and heatflow distribution in Japan are among the most strik-ing evidences which support that such induced flowdrives hotter isotherms into the corner ([22,56] andreferences therein).
Thus, a model where the location of melting inthe mantle wedge determines the position of thevolcanic front [49] would be consistent with naturalobservations.
Acknowledgements
We would like to thank O. Sigmarsson and P.Ulmer for discussions. The first version of themanuscript was prepared as M.W. Schmidt was aguest at Tokyo Institute of Technology. Many thanksfor enlightening discussions with S. Maruyama. [CL][FA]
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