Shinohara Hedenquist 1997 Constraints on Magma Degassing Beneath the FSE Porphyry Cu-Au Deposit,...

7/27/2019 Shinohara Hedenquist 1997 Constraints on Magma Degassing Beneath the FSE Porphyry Cu-Au Deposit, Philippines.

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JOURNAL OF PETROLOGY VOLUME 38 NUMBER 12 PAGES 17411752 1997

Constraints on Magma Degassing beneath

the Far Southeast Porphyry CuAu Deposit,Philippines

H. SHINOHARA1 AND J. W. HEDENQUIST2

1MINERAL AND FUEL RESOURCES DEPARTMENT, GEOLOGICAL SURVEY OF JAPAN, 1-1-3 HIGASHI, TSUKUBA 305, JAPAN

2INSTITUTE OF GEOLOGICAL AND NUCLEAR SCIENCES, P.O. BOX 31312, LOWER HUTT, NEW ZEALAND

RECEIVED JANUARY 1997; ACCEPTED AUGUST 1997

K-silicate alteration in the Far Southeast (FSE) porphyry CuAu INTRODUCTIONdeposit formed at ~14 Ma, concentric to dikes of quartz diorite

Magmas become saturated with volatile components asporphyry. At ~2 km paleodepth the hydrothermal system consisted

they ascend in the crust, and either erupt or crystallizeof magmatic hypersaline liquid and vapor at 550C and 50

at depth. Saturated magmas that remain at depth exsolveMPa (lithostatic pressure). Advanced argillic alteration formed at

an aqueous fluid that forms a hydrothermal system onthe same time over the deposit at 1 km paleodepth from acidic

separation from the magma chamber (Burnham, 1979).condensates of the vapor. At [13 Ma, K-silicate alteration was

Although part of the exsolved fluid (a low-density vapor)overprinted by 350C magmatic liquid (~5 wt % NaCl equiv.)

may discharge to the surface from passively degassingat hydrostatic pressure. Sericite alteration and much of the CuAu

volcanoes (Hedenquist, 1995), part will remain at depth,mineralization formed at this later stage. Evolution of the magmatic and this dense, residual magmatic liquid may be as-

fluid composition was simulated with a magma-chamber crys-sociated with hydrothermal mineralization in the por-

tallization model. Homogeneous crystallization during early stagephyry ore environment (Hedenquist & Lowenstern,

convection is assumed, whereas at 50 vol. % crystals the chamber1994).

becomes stagnant and crystallizes from rim to core over a narrowIn this paper we use evidence from a porphyry ore

crystallization interval. The model calculation, based on a magmadeposit in the Philippines to determine the composition

chamber with 2 km thickness at 6 km depth (150 MPa) andof the fluid that exsolved from the deeper parent magma.

[800C (saturated melt composition at 30 vol. % crystals, 5The evolving composition of the magmatic fluid at the

wt % H2O, 02 wt % Cl andD of 40), can reproduce thedepth of the ore body can then be used to place constraints

chemical and isotopic compositions of the early and late magmaticon the nature of the evolved magma chamber far below

fluids. The most critical factor controlling the compositional evolutionthe porphyry ore body. This approach differs from that

of the model hydrothermal system is the transition from convectiveof other studies where a portion of the parent magma

to stagnant magma-chamber crystallization. There is also a sharp

chamber is available to study directly (e.g. at Yerington,decrease in the rate of fluid exsolution associated with this transition, Nevada; Dilles, 1987).which can account for the thermal collapse of the FSE porphyry

The Far Southeast ( FSE) porphyry CuAu depositsystem, from K-silicate to sericite alteration.

and overlying Lepanto epithermal CuAu ore body are

located in northern Luzon. These deposits formed at the

same time, from coupled hydrothermal processes, in a

system dominated by magmatic fluid (Arribas et al., 1995;

Hedenquist et al., 1997). We use the isotopic compositionof contemporaneous K-silicate and advanced argillic al-

teration, formed from hypersaline liquid and vapor,KEY WORDS: magma; degassing; evolution; porphyry-copper; isotopes

Corresponding author.

Present address: Mineral and Fuel Resources Department, Geological

Survey of Japan, 1-1-3 Higashi, Tsukuba 305, Japan Oxford University Press 1997


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JOURNAL OF PETROLOGY VOLUME 38 NUMBER 12 DECEMBER 1997

respectively, plus fluid inclusion data, to estimate the hematitesericite. The chlorite alteration is cut by veinscomposition of the initial aqueous fluid that exsolved of euhedral quartz accompanied by anhydritefrom the parent magma at depth. Similar data on the chloritesericitepyritechalcopyritebornite, with halossubsequent stage of mineralization, dominated by a of sericite alteration (dated at 130007 Ma, n=10). Ahydrothermal fluid of lower salinity but also magmatic hydrothermal breccia, post-K-silicate alteration in timing,in origin, constrain the evolution of the crystallizing cut the center of the porphyry and the Imbanguila dacitemagma chamber, despite it lying several kilometers be- (Fig. 1) at the same time as formation of the later veinsneath the ore deposit. with sericite halos. Porphyry ore is concentric to the K-

silicate and sericitic alteration, with much of the gradeassociated with the sericitic stage. The Lepanto epi-thermal ore is supradjacent to porphyry mineralization,

THE FSE DEPOSIT hosted by the quartzalunite alteration (Fig. 1).Geology

The geology of the FSELepanto deposit is dominatedby three main lithologic groups (Fig. 1): (1) a volcanic to

Summary of constraintsepiclastic basement consisting of several units (including

Magma rose from a chamber at depth to form the early-

the Lepanto metavolcanic and Balili volcaniclastic rocks) mineral quartz diorite dikes after ~18 Ma. The parentof Late Cretaceous to middle Miocene age; (2) a Pliocenemagma, well below our depth of observation, exsolveddacitic pyroclastic and porphyry unit (Imbanguila horn-an aqueous fluid on saturation at some stage in itsblende dacite) which pre-dates alteration and CuAucrystallization, and this aqueous fluid ascended to themineralization; (3) an unaltered (post-mineralization)porphyry environment to form K-silicate alteration atPleistocene dacitic porphyry of flowdome complexes~14 Ma, including alteration of the early-mineral dikes.(Bato hornblendebiotite dacite). Near the FSE depositWe do not know the water content of the melt attwo volcanic vents are filled with Imbanguila dacitesaturation, nor the temperature, pressure or depth ofporphyry and pyroclastic rocks (Fig. 1). These vents areexsolution. We assume that saturation was timed to theclearly shown by the structural contours of the base ofemplacement of the early-mineral dikes, as many studiesthe Imbanguila unit (Garcia, 1991), and are probablyfind that porphyry dikes intruded with the same timingthe source of the Imbanguila pyroclastic and porphyryas hydrothermal alteration (e.g. Gustafson & Hunt, 1975).rocks. The Imbanguila dacite that hosts the Lepanto oreThus, the crystal content of the magma at the inception

body erupted at 2002 Ma, and the deposit is capped of saturation is taken as 30 vol. %, based on the minimumby unaltered (post-mineralization) Bato dacite, dated atcrystal content of the early-mineral quartz diorite dikes.~10 Ma (Arribas et al., 1995).

At the maximum depth of our observation (100 mFSE alteration and ore are centered on dikes of me-below sea level), the maximum trapping temperature oflanocratic quartz diorite porphyry that cut the epiclasticinclusions of the early aqueous fluid was 550C. Thisbasement rocks (Fig. 1). Minimum phenocryst content offluid consisted of hypersaline liquid (5055 wt % NaClthese altered dikes is ~30 vol. %, consisting of andesine,equiv.) and a coexisting low-salinity, gas-rich vapor (Hed-quartz, hornblende and biotite in order of decreasingenquist et al., 1997). The pressure of this fluid was 4050abundance (Concepcio

n & Cinco, 1989). Individual in-

MPa, indicating a paleodepth of 162 km in a lithostatictrusive bodies are typically 50150 m wide, and aresystem (i.e. paleosurface near 15001900 m elevation).elongated up to 5:1 in a northwesterly direction, parallel

A single-phase supercritical liquid at 550C intersects itsto the district-wide faulting (Concepcio

n & Cinco, 1989).solvus in the NaClH2O system at ~70 MPa (SourirajanThe intrusions extend up to about 300 m elevation, with

& Kennedy, 1962), i.e. at ~3 km paleodepth. The depthintrusive contacts nearly vertical at sea level. An inter- of immiscibility may have been greater if the fluid wasmineral series of altered leucocratic quartz diorite dikesgas rich, as the 550C solvus extends to a higher pressure.intrude to 500700 m elevation, cut the first set of

The hypersaline liquid formed the K-silicate alteration,intrusions, and contain lower grades of porphyry min-and the hydrothermal biotite indicates that this liquideralization.had a depleted-D isotopic composition (Fig. 2), averagingA core of early K-silicate alteration (biotite

45 D (Table 1). The vapor separated from themagnetiteK feldspar) is associated with veins of vitreous,two-phase fluid and condensed into meteoric water atanhedral quartz (biotite dated at 141005 Ma, n=6;shallower depths to form the acidic water that producedArribas et al., 1995). This is overlain by a shallow, blanket-the contemporaneous advanced argillic alteration. Basedlike zone of contemporaneous (142008 Ma, n=5) toon isotopic analysis of alunite, the end-member vaporlater advanced argillic (quartzalunite) alteration. The(Fig. 2) had a D composition of 20 to 25 (TableK-silicate alteration concentrically envelopes the dikes,

and is overprinted by an assemblage of chlorite 1). The isotopic compositions of the hypersaline liquid

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SHINOHARA AND HEDENQUIST MAGMA DEGASSING BENEATH A PORPHYRY Cu DEPOSIT

Fig. 1. Schematic NWSE longitudinal section through the Far Southeast (FSE) porphyry CuAu and Lepanto epithermal CuAu deposits,showing the geological units and extent of porphyry and high-sulfidation epithermal mineralization (from Concepcio

n & Cinco, 1989; Garcia,

1991). The upper limit of the hypersaline liquid, determined from fluid inclusion study, is shown by the heavy dashed line (Hedenquist et al.,1997).

Table 1: Comparison of constraints from FSELepanto deposits(Hedenquist et al., 1997) and the results of model calculation

Constraints from FSELepanto deposits

Stage Early Late Residual

Age ( Ma) 142008 141005 130007 143021

Mineral Alunite Biotite Illite Hornblende

Temp. (C)


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Fig. 2. Calculated fluid compositions for waters in equilibrium with hydrothermal minerals at temperatures determined from fluid inclusionand sulfur isotope geothermometry, and the estimated composition of local paleometeoric water (Hedenquist et al., 1997). The end-membermagmatic compositions (Table 1) are for fluid phases with 18O >56. The compositional fields were determined for water related to K-silicate alteration (biotite, n=6), advanced argillic alteration (alunite, n=8; samples from SE to NW extent of the Lepanto deposit), and sericiticalteration (illite, n=12; samples from the center of the deposit, labeled SE, to the northwest margin, labeled NW). Boxes show the ranges ofcompositions of water dissolved in silicic melts (Taylor, 1992) and discharged from high-temperature volcanic fumaroles (Giggenbach, 1992).

and vapor are consistent with those expected for water sea level. This is consistent with hydrostatic pressures at[

1600 m depth, i.e. similar to the paleodepth during K-exsolved from felsic magmas (Taylor, 1992), including silicate alteration at 14 Ma, indicating that there was littlethat of separated vapor discharging from volcanoes (Gig-erosion between these two alteration events.genbach, 1992). Hence, the hydrothermal system at 14

Ma was dominated by magmatic fluid.The +20 D difference between the vapor and

hypersaline liquid, estimated from mineral data, is similarTHE MODELto that predicted from extrapolation of experimental data

for fractionation between high-temperature vapor and Shinohara & Kazahaya (1995) modeled the physicalhypersaline liquid (Horita et al., 1995). Thus, FSE isotope processes involved in the release of hydrothermal fluidsdata are consistent with the vapor and hypersaline liquid from a crystallizing magma chamber. We extend thishaving coexisted, as also indicated by the con- model to evaluate the chemical and isotopic evolution oftemporaneity of alunite and biotite. However, we do not the fluid exsolved from such a crystallizing magma,know the vapor:liquid ratio, so we cannot estimate the and apply the results to interpreting the hydrothermal

composition of the bulk fluid exsolved from the magma evolution of the FSE porphyry deposit.at depth. As a stagnant magma chamber cools by conductiveSubsequent to the 14 Ma coupled K-silicate and heat loss to the surrounding crust, the magma chamber

advanced argillic alteration event, the porphyry was cut will crystallize from rim to core with a narrow, inward-by the hydrothermal breccia, typical of those associated moving zone of crystallization ( Jaeger, 1968; Brandeiswith build-up of pressure from a crystallizing magma & Jaupart, 1987). Fluid that exsolves from such a stagnant(Burnham, 1979). Sericite had already begun to form at magma will have a constant chemical and isotopic com-this time. position (Shinohara & Kazahaya, 1995).

The isotopic composition of sericite (illite) in the ore In contrast, if there is a temperature gradient acrossbody center (Fig. 2) indicates formation from liquid with a the magma chamber, convection may occur (Sparks,magmatic signature (i.e. 18O >5 and an average D 1990), although there is as yet no agreement on this topic

value of 43; Table 1). This liquid, with a salinity of ~5 (Marsh, 1990). If such convection occurs and mixes thechamber well, crystallization will occur homogeneously,wt % NaCl equiv., boiled at a temperature of 350C near

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and exsolution may proceed in an open-system manner. magma crystallizes and reaches saturation at 30 vol. %crystals (the minimum crystal content of early-mineralAs convective crystallization of a saturated magma cham-dikes at the FSE deposit) and 800900C. Magma isber proceeds, the chemical and isotopic composition ofassumed to convect until an abrupt transition to stagnantthe exsolving fluid will vary (Shinohara & Kazahaya,conditions occurs at a crystal content of 50 vol. %.1995). The pattern of compositional change will be

The composition of the initial aphyric magma (=melt)a function of several variables, particularly pressureis assumed to be 35 wt % H2O, 014 wt % Cl, and D(Shinohara, 1994; Cline, 1995).of 40, with the water and chlorine contents increasingHowever, homogeneous crystallization of a well-mixedto 50 and 02 wt %, respectively, at 30 vol. % crystals.magma chamber is only possible during the early stages of

As water solubility is also about 50 wt % at 150 MPamagma chamber crystallization, because magma viscosityand 800C (Silver et al., 1990), magma with 30 vol. %increases with increasing crystal content to the pointcrystals should be just saturated with water at its top atwhere movement in a magma chamber slows and even-6 km depth. The partition coefficient of Cl, i.e. the ratiotually stops, probably at 2555 vol. % crystals, a valueof the molal concentration of Cl in the aqueous phasethat depends on the rheological properties of the magmato that in the melt, is ~27 at the crystallization condition(Marsh, 1981). Therefore, if a magma chamber is con-(Shinohara et al., 1989). Although the chamber has avecting in its early stages of crystallization, there mustthickness of 2 km, the pressure of the crystallizing magmaeventually be a transition to stagnant magma-chamberis assumed to be 150 MPa throughout the chamber forcrystallization. This transition will be associated with asimplicity in this initial model.marked change in the nature of the fluid being exsolved

The deuterium fractionation factor between water in(Shinohara & Kazahaya, 1995), a change that may bean aqueous fluid and a melt containing 5 wt % dissolvedrecorded by alteration minerals in an overlying mag-water is ~15 (Taylor, 1991). Hornblende is assumedmatichydrothermal (i.e. porphyry) system. It is thisto form throughout the crystallization history, fixing achange that we examine by modeling data from the FSEmaximum of 10% of the water present in the saturatedporphyry system, in an attempt to constrain, remotelymelt (therefore 05 wt % water in the final pluton),and in a preliminary fashion, the conditions of magma-whereas Cl fixation is assumed to be negligible. Deu-chamber crystallization.terium fractionation between an aqueous fluid and horn-blende is 15 at 800C (i.e. there is no fractionationbetween melt and hornblende; Suzuoki & Epstein, 1976).

Model conditions The complete hornblende crystals are assumed to remainThe chemical and isotopic variations of fluid discharging in isotopic equilibrium with the surrounding magma

throughout crystallization.from a crystallizing chamber are calculated under the

following conditions. A variety of assumptions have beenmade, such as melt composition and pressure of crys-

tallization. These assumptions are based in part on typicalRESULTS OF CALCULATIONcompositions for silicic magmas deduced from melt in-Fluid discharge rateclusion studies (Lowenstern, 1995). The volume of the

magma chamber is that required to supply copper to a The thickness of the subsolidus portion of a stagnantmagma chamber (x, in m) increases in proportion to theporphyry deposit the size of the FSE (see below), andsquare root of time (t, in seconds),the thickness chosen controls the period of crystallization.

Pressure is a major control on the model results, andx(t)=at1/2 (1)

was chosen to replicate the observed fluid evolution. As

we demonstrate below, the model is relatively insensitive where a is a constant that depends on crystal growthto some of the geological variables, such as the crystal rate, nucleation rate, Stefan number, and thermal dif-content of the magma at saturation, and the crystal fusivity. The constant is estimated to be 8104 for acontent at which convection ceases, as long as saturation silicic magma chamber in the crust ( Jaeger, 1968; Bran-occurs before stagnation. deis & Jaupart, 1987). It will take 151012 s (~50 000

A flat-topped aphyric magma chamber 2 km thick and years) for complete crystallization of a stagnant chamber50 km2 in area (100 km3) is assumed to be emplaced to with a 1 km half-thickness (Fig. 3a).a depth of 6 km (150 MPa) and 1100C (its liquidus The maximum rate of crystallization during convectiontemperature, based on that of granodiorite; Naney, 1983). is assumed to be twice that of the stagnant chamberThe magma convects initially, and cools by conductive (Marsh, 1989), and the rate of crystallization once con-loss of heat to the crust (initially at ambient temperature) vection ceases is taken to be equal to that in a stagnantfrom the top and bottom of the chamber only, ignoring chamber (Fig. 3a). Intrusion of an aphyric magma to 6

km depth into a host rock at ambient temperature willfor simplicity any side-wall crystallization. The silicic

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Fig. 3. Variation of the average crystallized fraction (a) and fluid discharge rate (b) of the crystallizing magma chamber. (a) The variation forthe case without convection (dashed curve) is calculated with Eq. (1), where the average crystallized fraction is equal to the thickness of thesubsolidus magma (x) normalized by the half-thickness of the chamber (1000 m). Convection increases the crystallization rate by a factor of two(continuous curve). The convection stops at a crystal fraction of ~050; the subsequent rate of crystallization will be similar to that in a stagnantchamber. (b) The rate of fluid discharge is calculated by multiplying the crystallization rate by the area of the chamber (50 km2) and watercontent (45 wt %, excluding water in hornblende). The curve is the calculated result for a chamber that convects while the crystallized fractionis 050. The fluid discharge rate from the stagnant chamber is smaller by a factor of five compared with the average for the convection stage.The hatchured and shaded areas show the duration of pre- and post-saturation convection (1000 and 2000 years), respectively.

crystallize during convection to 30 vol % crystals in 1 Compositional evolutionkyr, at which point the melt becomes saturated. The

Open-system fluid exsolution and discharge occurtransition from convecting to stagnant magma conditions during convection because the bubbles formed in theis set here at 50 vol. % crystallization, which will be chamber can be removed rapidly from the homo-achieved 2 kyr after saturation; the stagnant magma geneously mixed chamber (Shinohara & Kazahaya,chamber will completely crystallize in a further 25 kyr. 1995). In contrast, in a stagnant magma chamber,Thus the time required from melt saturation (at 30 vol. crystallization and fluid exsolution only occur over a% crystals) to complete crystallization of the initially narrow zone of crystallization (700C) vapor discharge from passively stagnant chamber does not vary over the lifetime ofdegassing volcanoes (5106 ton/year; Hed- crystallization of the whole chamber (Shinohara & Ka-enquist, 1995). After convection ceases, the fluid discharge zahaya, 1995). Most previous modeling of the variationrate quickly decreases to an average of about 500 ton/ of chemical and isotopic compositions of the dischargingday (2105 ton/year, equal to the lower rate of low- fluid has assumed open-system conditions (Taylor, 1991;temperature fumarolic discharge from passively degassing Shinohara, 1994; Cline, 1995). Such open-system dis-

charge, however, is limited to the early stage ofvolcanoes).

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Fig. 4. Variation of NaCl (a) and D (b) content in the exsolving fluid (continuous curves) resulting from the crystallization of an early convectingand late stagnant magma chamber. Initial conditions at saturation (030 crystallized fraction) for the calculations are 5 wt % H2O, 02 wt % Cland 40 D. Open-system degassing is expected during the convection stage (030050 crystallized fraction), whereas complete, closed-systemdegassing occurs from the narrow crystallization interval in the later stage; variation under open-system conditions is shown by the light curve.Hornblende (dashdot curve) is assumed to incorporate 10% of the water at saturation (05 wt %) with a D 15 less than the exsolved fluid.

crystallization, during magma convection, or possibly model isotopic composition of this vapor and liquid canduring eruption. be determined by a mass balance calculation, assuming

The variation of the NaCl concentration in the ex- a simple NaClH2O fluid (Table 1). NaCl concentrationssolving fluid depends on the relative exsolving tendencies of the 550C, 50 MPa vapor and liquid are 07 and 50

of Cl and water (Shinohara, 1994). As Cl tends to wt %, respectively (Sourirajan & Kennedy, 1962). Theexsolve more readily than water at 150 MPa, the NaCl deuterium fractionation factor between vapor and liquidconcentration decreases with time under open-system is taken to be 20, estimated by extrapolation of 4 molaldegassing of the convecting chamber. The NaCl con- NaCl experimental data (Horita et al., 1995).centration of the exsolving fluid decreases from 86 wt The fluid exsolved during initial crystallization (from 30% at the beginning of saturation (30 vol. % crystals) to 50 vol. % crystals, taking ~2000 years in a convectingto 79 wt % at 50 vol. % crystallization; the average magma chamber) has a bulk composition of 83 wt %concentration is 83 wt % (Fig. 4a). After convection NaCl and 27 D. Cooling of this fluid to 550Cceases at 50 vol. % crystallization, the NaCl concentration during ascent along permeable channelways (an apophysewill remain constant at 62 wt % (Fig. 4a). During extending upward from the parent magma chamber andconvection, the D values of the discharging water de- splitting into individual dikes) results in the solvus beingcrease from 25 to 30 (bulk 274), then become reached at [3 km depth (70 MPa). Further ascent to 2constant at 43 on transition from the convecting km depth (50 MPa) forms hypersaline liquid of 50 wt %to stagnant state (Fig. 4b). The D value of igneous NaCl and 45 D, and vapor of 07 wt % NaCl andhornblende, which is equal to that of the melt, also 25 D. A mass balance based on the salinity indicatesdecreases with time in the early stage. However, if that the vapor accounts for 916 wt % of the bulk waterhornblende continually reequilibrates with the exsolving exsolved from the magma, whereas the remaining 84fluid by diffusion through to the later stage (Zaluski et wt % comprises the hypersaline liquid, i.e. there is 11al., 1994), the hornblende will finally have a model times more vapor, by mass, than liquid in the earlycomposition ofD 58.

hydrothermal system. These model results of the exsolved

bulk fluid, present as liquid and vapor phases near the

apices of the dikes, match well the isotopic compositionsDISCUSSION

Fluid evolution of the FSE system of the aqueous phases that formed the K-silicate andadvanced argillic alteration, respectively, plus the salinityThe early fluid in the FSE hydrothermal system was a

mixture of immiscible vapor and hypersaline liquid. The of the associated inclusion fluids ( Table 1).

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Once the saturated magma chamber contains >50 to stagnant transition, which can account qualitatively forvol. % crystals, convection ceases, and the remaining the temperature decrease, from 550 to 350C, recorded incrystallization (from 50 to 100 vol. %) will take the veins associated with the change from the early K-silicatebalance of the ~25 000 years under stagnant conditions. to later sericite stage of alteration.During this period, an evolved fluid of constant com- Igneous hornblende collected from a relatively freshposition exsolves from the magma with a composition dacite breccia at surface (possibly inter-mineral with an62 wt % NaCl and 43 D, as exsolution now occurs age of 143021 Ma) has a D value of 75, similarfrom only the relatively narrow rind that crystallizes to hornblende from pre- and post-mineral volcanic rocksfrom the margin inward (Candela, 1991; Shinohara & in the district (ranging from 70 to 80). This DKazahaya, 1995). Ascent and cooling of this fluid to value is lighter than that calculated here (58), probably350C also matches well the salinity and isotopic com- because this rock erupted to the surface. During eruption,position of water associated with the later-stage sericite shallow, open-system degassing (Taylor, 1991) typicallythat overprinted the early K-silicate alteration (Table 1). results in D-depleted hornblende. The hornblende crys-

In addition to the model results fitting the observed tals may also have been affected by deuteric alterationcompositional variation, the lifetime of the hydrothermal (Nabelek et al., 1983).system, 100 kyr, is of the same order of magnitude asthe ~27 kyr duration of saturated crystallization for themodel magma chamber. Doubling the half-thickness of

Porphyry constraints on magma degassingthe parent magma chamber to 2 km will increase thebeneath the FSE systemtotal crystallization period after saturation by a factor ofThe model of crystallization of a 2 km thick magma atfour, i.e. ~108 kyr.6 km depth (150 MPa) can reproduce the chemical andBut why does the later magmatic fluid, that associatedisotopic evolution of the magmatic fluid in the FSEwith sericite alteration, not separate into hypersalineporphyry system. The parameters used for the calculationliquid and vapor during ascent? In the early porphyry(i.e. magma chamber thickness, depth and pressure,system, inception of immiscibility and formation of vaportiming of saturation and crystallization history, and con-and hypersaline liquid are deduced to occur at about 3centrations of water, Cl and D) are not constrainedkm depth, followed by vapor removal near the top ofdirectly from field evidence, as the source magma cham-the zone of K-silicate alteration. During saturated magmaber beneath the deposit is undocumented, unlike that ofconvection (~2000 years), there is a huge advective fluxthe Yerington deposit (Dilles, 1987). However, our model

of volatiles (Fig. 3b) and heat that generates a plume of is relatively insensitive to several variables, such as timingmagmatic fluid (Fig. 5a). Once the crystal proportionof saturation and crystallization history. For example,increases to about 50 vol. %, magma convection ceasessaturation of an aphyric melt at 0 vol. % crystals, andand the mass (Fig. 3b) and heat flux decrease markedlystagnation at 25 vol. % crystals, results in the model[i.e. 09(50/70)=643% of the total water is exsolvedresults changing by 36106 tonnes (Garcia, 1991). As-during stagnant crystallization never intersects its solvus

suming that the quartz diorite magma contained a mini-during ascent, and hypersaline liquid cannot form. A mum of 50 mg/kg Cu (Cline, 1995; Dilles & Proffett,similar argument was used by Fournier (1987) to predict1995), the minimum volume of such a source magmaunder what conditions a magmatic intrusion would cause(density 2000 kg/m3) would be ~36 km3 if there werea dilute meteoric water to exceed its critical point.complete extraction of copper from the magma andIn summary, the transition from a convecting to a100% transport and deposition efficiency in the orestagnant magma chamber is required to match the ob-bodies. Therefore, the actual size of the source magmaserved compositional evolution of the FSE hydrothermalchamber is likely to be larger, of the order of 100 km3,system. Simple fluid exsolution from a crystallizing stag-similar to that observed at Yerington (Dilles & Proffett,nant magma chamber will result in a constant com-1995), and predicted by Cline (1995).position of magmatic fluid that does not match the FSE

The calculated duration of magmatic fluid dischargeobservations. Equally convincing, when coupled with thefrom a 2 km thick, 100 km3 chamber is 27 kyr, of theevidence for a compositional change, is the sharp decrease

in advective fluid (and hence heat) flux at the convective same order of magnitude as the 100 kyr lifetime of

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Fig. 5. Schematic representation of the parent magma chamber and hydrothermal system associated with the FSE porphyry deposit, during(a) K-silicate and (b) sericite alteration, at ~14 Ma and [13 Ma, respectively. The high-temperature magmatic plume in the early stage isrelated to the high discharge rate from the convecting magma chamber. The collapse of isotherms in the later stage reflects the factor of fivedecrease in the rate of fluid exsolution from the stagnant magma chamber (Fig. 3b). The shaded area near sea level (at the apex of the dikes)

shows the extent of the K-silicate and sericite alteration and CuAu mineralization in the FSE porphyry system (lower limit is a guess). Thelithostatic to hydrostatic pressure transition occurs at a temperature 4wt % water at ~10 vol. % crystallization (Dilles, 1987;deposits of the Yerington district (Dilles & Proffett, 1995).

A combination of the crystallization pressure and initial Dilles & Proffett, 1995).

The partition coefficient of chlorine has a strong de-water content of the magma controls the timing of water

saturation during crystallization. When the initial water pendence on pressure, which results in a pressure-de-

pendent pattern of Cl variation during open-systemcontent of the magma is far below its saturation level

(


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of fluid discharge, resulting in a very low salinity of late- of early-mineral dikes. Nevertheless, we believe that ourexsolving fluid. Such a high crystallization pressure is not results are qualitatively valid, as the model is relativelyconsistent with the observation of a 5 wt % NaCl fluid insensitive to several of the variables (e.g. degree ofduring the later, sericite stage at the FSE system. Likewise, crystallization at saturation, and the timing of the trans-if crystallization had occurred at a pressure


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porphyry and epithermal CuAu deposits, Philippines. EconomicREFERENCESGeology (in press).

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