Carbonate stability and £uid composition in subducted oceanic crust: an experimental ...hera.ugr.es › doi › 1499737x.pdf · 2005-04-12 · Carbonate stability and £uid composition

Carbonate stability and £uid composition in subductedoceanic crust: an experimental study on H2O^CO2-bearing

basalts

Jose© F. Molina a;1, Stefano Poli b;*a Departamento de Mineralog|̈a y Petrolog|̈a, Campus de Fuentenueva, University of Granada, 18002 Granada, Spain

b Dipartimento Scienze della Terra, Universita© degli Studi di Milano, Via Botticelli 23, 20133 Milan, Italy

Received 26 August 1999; accepted 21 January 2000

Abstract

Carbonates and hydrates are common products of the alteration of the upper basaltic crust in modern oceans.However, phase relationships and devolatilization reactions in altered CO2-bearing metabasalts during the subductionprocess are still poorly known. A series of fO2-buffered piston cylinder experiments were performed on three basalticmodel compositions in the presence of a H2O^CO2 mixed fluid, at pressures from 1.0 to 2.0 GPa and temperatures from665 to 730³C. Experimental results on a tholeiite composition demonstrate that amphibole coexists with calcite atP91.4 GPa, with dolomite at 1.49P91.8 GPa, and with dolomite+magnesite at pressures higher than 1.8 GPa. Thestability of calcite increases with pressure with increasing Fe/(Fe+Mg) of the bulk composition. Omphacite was foundin tholeiite only at 2.0 GPa, 730³C. Garnet, plagioclase, paragonite, epidote and kyanite further complicate phaserelationships in the pressure range investigated. Estimates of the coexisting fluid compositions, on the basis of mass-balance and thermodynamic calculations, demonstrate the continuous H2O enrichment with increasing pressure anddecreasing temperature. An almost purely aqueous fluid (X CO2 6 0.05) is obtained at 2.0 GPa, 665³C. Hydrous fluidsand relatively high modal proportions of carbonates at high pressure and low temperature conditions are responsiblefor the displacement of the appearance of omphacite at higher pressures than in H2O-saturated, CO2-free systems.Modeling of devolatilization reactions along subduction zone geotherms reveals that significant decarbonation isfeasible only at low pressures (in the forearc region) and at relatively high temperatures, once young oceanic crust issubducted at slow convergent rates. When the subduction process approaches steady-state conditions, CO2 isfractionated in the solid and deep recycling of CO2 is expected to account for the global-scale imbalance at convergentmargins. ß 2000 Published by Elsevier Science B.V. All rights reserved.

Keywords: carbonates; £uid phase; experimental studies; P-T conditions

1. Introduction

Hydrous £uids released from the altered ocean-ic crust during subduction play a key role in theevolution of convergent margins, as demonstrated(a) by the high H2O content in melt inclusions of

0012-821X / 00 / $ ^ see front matter ß 2000 Published by Elsevier Science B.V. All rights reserved.PII: S 0 0 1 2 - 8 2 1 X ( 0 0 ) 0 0 0 2 1 - 2

* Corresponding author. Tel. : +39-2-23698324;Fax: +39-2-70638681; E-mail: [email protected]

1 Present address: Mineralogisk-Geologisk Museum, Sars-gate 1, N-0562, Oslo, Norway.

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olivines from island arc basalts (1^3 wt% H2O)[1] ; (b) by the high H2O/CO2 ratios (10^100) involatile £uxes from arc volcanoes [2] ; (c) by thecomposition of £uid inclusions entrapped in high-pressure minerals [3] ; and (d) by the large solubil-ity of silicates in H2O-rich £uids, which accountfor modal metasomatism in the mantle wedge [4].As a consequence, H2O transport and releasethroughout the subduction process have been ex-tensively reproduced in high pressure experiments[5^8] and numerically modeled [8,9].

Nevertheless, once composite analyses of thealtered oceanic crust are considered (Table 1,Fig. 1), such enrichment in H2O is at ¢rst sightunexpected. Aragonite in vugs and veins as wellas calcium carbonate in the groundmass a¡ect theoceanic crust down to at least 1 km depth leadingto bulk CO2 contents which can be as high as theH2O contents [10]. Even though of minor impor-tance compared to hydrous phases, carbonatesare present as secondary minerals also in gabbroicrocks in the lower oceanic crust, and ophicarbon-ates are found along fracture zones [11]. Carbo-nate-bearing breccias are commonly found in nu-merous Mediterranean ophiolites [12] suggestingmajor compositional heterogeneities of the upperoceanic crust on scale lengths between 10 and 100m [13].

Therefore the altered basaltic crust as well ascarbonate/basalt breccias should be regarded notonly as H2O reservoirs in subducted slabs, butalso as CO2 reservoirs, as relevant as the over-lying sedimentary layer. This is further demon-

strated by the occurrence of dolomite- and/ormagnesite-bearing ma¢c eclogites from high-pres-sure and ultra-high-pressure terrains [14^16].

The observed H2O fractionation and the imbal-ance of CO2 £uxes via arc-related phenomenahave led numerous authors to argue about shal-low vs. deep recycling of CO2 into the mantle [17^19]. Experimental determination of the reactionsdolomite+2 enstatite = forsterite+diopside+2 CO2,enstatite+2 magnesite = 2 forsterite+2 CO2, dolo-mite+2 coesite = diopside+2 CO2 ([20] and refer-ences therein) would suggest that carbonates arestable in a large spectrum of mineral assemblages(both saturated and undersaturated in SiO2) atpressure^temperature conditions of subductingslabs [21]. This idea has been further exploredexperimentally on carbonate-bearing eclogites[22] and modeled numerically on ophicarbonates[19] leading to the conclusion that decarbonationof the ma¢c and ultrama¢c portions of the sub-ducted oceanic crust is hardly achieved.

Nevertheless, as long as most solid phases insubducted basaltic and gabbroic rocks are com-plex solid solutions, continuous exchange and nettransfer reactions are expected to involve not onlyhydrates (amphibole, zoisite, lawsonite, etc.[6,8,23]) and H2O, but also carbonates and CO2.It is therefore the goal of this experimental work

Table 1Bulk composition of gels (wt%)a

HFBb OTBc HMBd AOBe

SiO2 49.32 52.38 51.49 49.79Al2O3 11.91 16.93 15.39 16.91FeOT 17.94 10.29 8.24 9.74MgO 7.08 7.13 10.38 7.23CaO 9.68 10.05 12.39 13.92Na2O 4.06 3.21 2.12 2.41mg 41.31 55.26 69.19 56.94aComposition recalculated in the model system NCFMAS.bHigh-Fe basalt [24,25].cOlivine tholeiite [26].dHigh-Mg basalt [27].eAverage of altered oceanic basalt compositions [10].

Fig. 1. H2O and CO2 contents (wt%) in altered oceanic ba-salts. Shaded ¢eld is the composition displayed by compositeanalyses after Staudigel et al. [10]. Star is the estimate ofglobal subducting sediments (GLOSS) by Plank and Lang-muir [51]. Arrow indicates direction of the ophicarbonatecomposition [11].

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to document variations in mineral assemblagesand to evaluate the coexisting £uid compositionsas a function of pressure, temperature and whole-rock composition. We show that the progressivedecomposition of hydrates coupled to the increas-ing stability of carbonates with increasing pres-sure is responsible for a continuous displacementof £uid composition towards H2O with pressure.Therefore, if signi¢cant decarbonation in the fore-arc region is possible, dehydration is expected tobe the dominant devolatilization process at pres-sures typical for arc magmatism.

2. Experimental methods

2.1. Bulk chemistry

All experiments were performed on three start-ing compositions (Table 1) ranging from high-Feto high-Mg tholeiites in the model system Na2O^CaO^FeO^MgOÂl2O3^SiO2 (NCFMAS) in thepresence of a CÔ^H £uid. Composition 1(HFB), which closely approaches Fe^Ti gabbrosfound in eclogite facies terrains [24] and in ODPdrills [25], has a mg ( = 100 Mg/(Mg+Fe))W41and low Al2O3. Composition 2 is a representativeMOR basalt (olivine tholeiite basalt, OTB) havingmgW55, close to the LT and TB starting materi-als used in previous H2O-saturated high-pressureexperiments [8,26]. Composition 3 is a high-Mgbasalt (HMB) with mgW69, similar to magnesianolivine tholeiites which are thought to representprimitive MORB melts [27]. Bulk compositions 1and 3 represent end members of MORB magmas,whereas composition 2 is also very similar to boththe average of the altered oceanic basalts reportedby Staudigel et al. [10] (see Table 1) and the GA1basalt employed by Yaxley and Green [22] tostudy near solidus phase relationships in carbo-nate eclogites, although their addition of 10 wt%of calcite in most experiments led to higher CaOcontents.

2.2. Starting materials

The starting materials used consist of mixturesof gels and crystal seeds. Because amphibole and

garnet are expected to show the slowest nuclea-tion rate, V10 wt% of a mixture of natural trem-olite, natural pyrope, and synthetic almandineseeds (the latter synthesized following [28]) wasadded to gels. As a true bracketing of mineralcompositions (e.g. re-running experimentalcharges at di¡erent pressure^temperature condi-tions) is hampered by the extremely low reactivityof fully crystalline run products, growth rims onalmandine and pyrope seeds were also used toevaluate approach to equilibrium. Gels were pre-pared following the methods of Hamilton andHenderson [29], using diluted nitrate solutions assources for aluminum, magnesium, calcium, andsodium, tetraethyl orthosilicate (TEOS) for sili-con, and ferric benzoate dissolved in N,N-dime-thylformamide for iron.

2.3. Oxygen bu¡ering and £uid speciation

Oxygen fugacity (fO2) was bu¡ered by Ni^NiO(NNO) bu¡er using inner Ag75Pd25 capsules andouter Au capsules. Up to three AgPd capsules(outer diameter 2 mm) were packed into onelarge, 8 mm long, welded Au capsule with ca.200 mg NNO bu¡er and 30^50 mg H2O. Weldedinner capsules contain the starting materials plusV10 wt% oxalic acid dihydrate (OAD) (see Table2). In the present experimental con¢guration,OAD breaks down to yield approximately a 1:1molar mixture of H2O^CO2 according to the re-action:

C2O4H2W2 H2O � 2 H2O� 2 CO2 �H2 �1�

at fO2 values ¢xed by the NNO bu¡er [30]. How-ever, because iron in ¢red gels is completely oxi-dized, an extra amount of H2O is generated insidethe AgPd capsule due to the reduction of the Fe3�

in the starting materials to get the bulk Fe3�/FeT

ratio at the speci¢ed pressure, temperature, andfO2 conditions of the experiment:

Fe2O3 �H2 � 2 FeO�H2O �2�

where H2 is provided by Eq. 1 and by the oxygenbu¡er equilibrium:

Ni�H2O � NiO�H2 �3�

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Furthermore, because the graphite boundaryin the CÔ^H system closely approaches thejoin H2O^CO2 at high pressure [31], fugacitiesof CO and CH4 species are negligible at theO:H ratios investigated here, and £uids producedby decomposition of OAD can be safely approxi-mated by binary mixtures of H2O and CO2. Wewill therefore estimate the total amount of H2Oin each AgPd capsule following iterative non-linear mass-balance and thermodynamic calcula-tions (see below) assuming a binary H2O^CO2

£uid.

2.4. Experimental procedures and analyticaltechniques

All experiments were carried out in a single-stage piston cylinder apparatus with a 22 mmbore using both full-salt and NaCl^MgO assem-blies. Temperatures were measured with K-typethermocouples accurate to þ 5³C. Pressure uncer-tainties are within þ 0.03 GPa. A temperaturerange from 665 to 730³C, at pressures from 1.2to 2.0 GPa, was investigated in this study which iscomplementary to experiments by Yaxley andGreen [22]. Run durations range from 200 to

Table 2Run table

Run number Bulka T (³C) P (GPa) Time (h) OAD (wt%)b Run productsc

R1 HFB 665 2.0 430 12.1 gar, omp, pg, dol, mag, qR3 HFB 680 1.2 309 11.5 amp, plg, cc, qR10 HFB 680 1.4 353 13.2 amp, plg, dol, qR2 HFB 680 1.6 473 11.7 gar, amp, plg, dol, mag, qR9 HFB 680 1.8 312 12.8 gar, amp, pg, dol, mag, qR6 HFB 730 1.2 282 8.8 amp, di, plg, cc, qR7 HFB 730 1.4 345 11.7 gar, amp, plg, cc, qR4 HFB 730 1.6 363 7.8 gar, amp, cc, qR8 HFB 730 1.8 183 8.5 gar, amp, omp, dol, mag, qR5 HFB 730 2.0 257 7.2 gar, omp, dol, mag, qR1 OTB 665 2.0 430 11.7 gar, amp, pg, dol, mag, qR3 OTB 680 1.2 309 8.8 amp, ep, plg, cc, qR10 OTB 680 1.4 353 5.3 amp, ep, plg, dol, qR2 OTB 680 1.6 473 11.6 gar, amp, pg, dol, qR9 OTB 680 1.8 312 9.4 gar, amp, pg, dol, qR6 OTB 730 1.2 282 13.8 amp, ep, plg, cc, qR7 OTB 730 1.4 345 8.4 amp, ep, plg, cc, qR4 OTB 730 1.6 363 10.9 gar, amp, ky, dol, qR8 OTB 730 1.8 183 9.9 gar, amp, ky, dol, qR5B OTB 730 2.0 291 10.8 gar, amp, omp, ky, dol, mag, qR1 HMB 665 2.0 430 12.6 gar, amp, ep, ky, dol, mag, qR3 HMB 680 1.2 309 11.0 amp, ep, dol, qR10 HMB 680 1.4 353 9.8 amp, ep, dol, qR2 HMB 680 1.6 473 9.3 gar, amp, ep, ky, dol, qR9 HMB 680 1.8 312 10.2 gar, amp, ep, ky, dol, mag, qR6 HMB 730 1.2 282 7.6 amp, ep, plg, cc, qR7 HMB 730 1.4 345 9.9 amp, di, ep, cc, qR4 HMB 730 1.6 363 9.2 gar, amp, ep, qR8 HMB 730 1.8 183 11.5 gar, amp, ep, ky, dol, qR5 HMB 730 2.0 257 8.6 gar, amp, ep, ky, dol, mag, qaStarting materials with 10 wt% of crystal seed mix.bAmount of OAD added to the starting material.cMineral abbreviations: alm = almandine; amp = amphibole; cc = calcite; di = diopsidic clinopyroxene; dol = dolomiteânkerite sol-id solution; ep = epidote; gar = garnet; ky = kyanite; mag = magnesite^siderite solid solution; omp = omphacitic clinopyroxene;pg = paragonite; plg = plagioclase; q = quartz; £ = H2O^CO2 £uid.

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more than 470 h (Table 2). After unloading,charges were pierced, and the presence of excess£uid both in the outer Au capsule and in the innerAgPd capsules was veri¢ed by observation of liq-uid bubbling out of the capsule or by weight lossafter heating at 110³C. Charges were sectioned,embedded in epoxy resin and polished for electronmicroprobe analysis and scanning electron mi-croscopy (SEM). Analyses were obtained bywavelength dispersive spectrometry (WDS) withan ARL-SEMQ electron microprobe (CSGAQ-CNR, Milan) operating at 15 kV acceleratingvoltage and 20 nA sample current.

3. Experimental results

3.1. Description of the synthetic mineralassemblages

Amphibole (Fig. 2a) is always present, exceptfor the HFB composition, where amphibolebreaks down at PW1.9 GPa. Amphiboles aremagnesio-hornblende in low-pressure runs mov-ing to barroisite at higher-pressure conditionsdue to an increase of the Na occupancy in theM4 site (0.37^0.87 atoms per formula unit, apfu,NaM4 in HMB, 0.45^1.07 apfu NaM4 in OTB, and

Fig. 2. SEM pictures of run products: (a) euhedral garnet and, on the left, prismatic amphibole in OTB composition at 1.8 GPaand 730³C; (b) amphibole aggregates in HFB composition at 1.2 GPa and 680³C; (c) rhombohedral crystal of dolomite in HFBcomposition at 2.0 GPa and 730³C; (d) euhedral plagioclase in OTB at 1.2 GPa and 730³C.

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0.59^1.28 apfu NaM4 in HFB, mineral formulaebased on 24 atoms of oxygen).

Garnet appears at pressures higher than V1.4GPa, which is more than 0.5 GPa higher than inH2O-saturated experiments on basalt [26]. Gar-nets have almandine fractions ranging from 0.45to 0.65 (Table 3).

Unexpectedly, omphacitic clinopyroxenes onlyoccur at the highest pressures investigated andtheir appearance is clearly related to the mg ofthe bulk composition. Omphacite (Na from 0.53to 0.68 apfu) in HFB (mg = 41) forms at 1.8 GPaand 730³C, it appears at 2.0 GPa in the OTBcomposition (mg = 55) at 730³C, but it is still ab-sent in HMB (mg = 69) at 2.0 GPa. This result issigni¢cantly di¡erent from the appearance of om-phacite in H2O-saturated ma¢c systems, whichtakes place at 1.5^1.6 GPa at 650³C [7,26]. Diop-sidic clinopyroxene is present in low-pressure ex-periments (1.2^1.4 GPa) at 730³C in the HFB andHMB compositions (Fig. 3).

Epidote stability is reduced by reduced H2Oactivities and by a decrease in mg. Epidote(Fe3�/(Al-2+Fe3�)W0.2^0.4) is always absent in

the HFB composition, it occurs at P6 1.4 GPaat 680³C, and P6 1.6 GPa at 730³C in the OTBcomposition, whereas it is always present in theHMB composition.

The stability ¢eld of plagioclase again shows astrong dependence on bulk composition. In theHFB composition, plagioclase occurs up to 1.6GPa at 730³C, with a composition XAbW0.88^0.92. Plagioclase disappears at Ps 1.4 GPa(XAbW0.75^0.83) in the OTB composition,whereas it occurs only at 1.2 GPa and 730³C(XAbW0.70) in the HMB composition. Plagio-clase is replaced with increasing pressure by para-gonite in the HFB and OTB compositions, at665³C and 680³C, and by kyanite in the OTBcomposition at 730³C. Kyanite is absent in theHFB composition, probably as a consequence ofthe low alumina content (V12 wt%) (Fig. 3).

These data reveal a large stability ¢eld for am-phibole+carbonate at the expense of the pair gar-net+clinopyroxene. Carbonate-bearing amphibo-lites in eclogite facies terrains should thereforebe regarded with caution as the absence of thediagnostic garnet+omphacite assemblage is not

Table 3Representative microprobe analyses of the mineral phases

Mineral phase Amp Amp Gar Gar Gar Cpx Cpx Cc Dol Dol MagP (GPa) 1.4 1.8 2.0 1.6 2.0 1.4 2.0 1.2 1.6 2.0 2.0T (³C) 680 730 665 730 730 730 730 680 680 730 730Bulk composition OTB HMB HFB OTB OTB HMB HFB OTB HMB HFB HFB

(wt%)SiO2 46.0 48.9 39.1 38.4 39.8 52.3 55.8Al2O3 15.2 13.5 22.6 22.0 22.2 8.7 16.3FeOT 14.2 10.7 26.8 26.9 23.7 7.0 7.6 4.7 9.52 17.3 32.8MgO 10.3 12.9 2.5 5.5 5.5 10.8 3.5 4.0 14.1 10.8 20.0CaO 8.6 7.7 8.4 8.7 9.9 20.2 7.1 47.1 32.9 27.7 3.5Na2O 3.1 2.7 2.4 9.9Total 97.3 96.4 99.4 101.5 101.2 101.5 100.2 57.0 56.5 55.8 56.3Atomic proportionsa

Si 6.561 6.836 3.060 2.957 3.031 1.887 1.966Al 2.555 2.225 2.085 1.997 1.993 0.372 0.677Fe 1.694 1.251 1.754 1.732 1.509 0.212 0.224 0.065 0.124 0.240 0.450Mg 2.190 2.688 0.296 0.635 0.630 0.581 0.186 0.099 0.327 0.267 0.489Ca 1.314 1.147 0.702 0.724 0.809 0.781 0.270 0.836 0.549 0.493 0.061Na 0.846 0.743 0.167 0.678Fe2� 0.846 0.185 0.202 0.189Fe3� 0.848 1.066 0.010 0.035NaM4 0.686 0.853aAtoms of oxygen in mineral formulae: 24 in amphibole; 12 in garnet; 6 in clinopyroxene; 1 in carbonates.

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indicative of a lower-pressure metamorphic event,but simply as a displacement of phase equilibriain the presence of a H2O^CO2 mixed £uid.

3.2. Adjustment to chemical equilibrium

In spite of the presence of seed relics, the com-positions of newly formed clinopyroxene grainsand of growth rims on garnet (both on almandineand on pyrope) and on amphibole seeds arehomogeneous. The application of widely usedgeothermometers suggests a close approach tochemical equilibrium. A garnet^clinopyroxenethermometer [32] yields V680³C for runs at665³C, whereas temperatures ranging from 730to 790³C are calculated for experiments at730³C. A garnetâmphibole thermometer [33] (as-suming all iron as Fe2� to be consistent with thecalibration procedure) leads to temperatures from695 to 780³C for runs at 730³C, and from 610 to735³C for experiments at 665³C and 680³C.

4. Fluid composition and phase proportions:mass-balance approach

Once the chemical composition of the coexist-ing phases, including hydrates and carbonates, isknown at each pressure and temperature condi-tion, and hydrogen is considered as the only per-fectly mobile component, the abundances of the

phases and the composition of the coexisting £u-ids can be calculated by a mass-balance approach.The following overdetermined non-linear systemhaving i equations (i = number of chemical com-ponents) and j+1 unknowns (NP = abundance ofsolid-phase P, j = total number of phases,X fluid

CO2= fraction of CO2 in the £uid) was solved:

A1 �Xj31

P�1

NP WY P1

T

Ai32 �Xj31

P�1

NP WY Pi32

AH2O �Xj31

P�1

NP WY PH2O �N fluidW�13X fluid

CO2�

ACO2 �Xj31

P�1

NP WY PCO2�NfluidWX fluid

CO2

where Ai is the fraction of component i in thesystem and YP

i is the fraction of component i inphase P.

The in£uence of Fe3�/(Fe2�+Fe3�) was esti-mated solving the system at variable Fe3�/(Fe2�+Fe3�) and physically meaningful, i.e. pos-itive, values of NP . All Fe as Fe2� and Fe3�/(Fe2�+Fe3�) = 0.5 were used as limiting condi-tions (Table 4). Because Fe3�/(Fe2�+Fe3�)W1 in

Fig. 3. Experimentally determined phase relationships for basalts in the presence of a H2O^CO2 mixed £uid. Reaction albite(Ab) = jadeite (Jd)+quartz (Q) is computed after Holland and Powell [34]. Albite solidus at X fl

CO2= 0.5 from Eggler and Kadik

[52].

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the starting material, these limits will correspondto the maximum and the minimum hydrogen ad-ditions to the charge, respectively, and thereforeto the maximum and minimum H2O contents inthe £uid. The abundance of solid phases in theruns estimated by the mass-balance calculationsare displayed in Fig. 4. Oxygen equivalent wasused as unit because of the direct relation to mo-dal (volume) proportions.

In all bulk compositions, amphibole shows thehighest abundance (60^40 oxygen%), moderatelydecreasing with increasing pressure. The appear-ance of omphacite is responsible for a suddendecrease of amphibole, and an increase of garnetcontent, suggesting a reaction of the type:

amphibole! garnet� omphacite

which is continuous with increasing pressure.The most remarkable di¡erence of computed

phase proportions compared to data obtainedon a model basalt at H2O-saturated conditions[26] is the displacement of omphacite appearanceto pressures higher than 1.8^2.0 GPa. As a resultof this displacement, plagioclase breakdown isno more directly responsible for the appearanceof sodic clinopyroxene. Therefore, even thoughcounter-intuitive, the addition of CO2 to the sys-tem expands the ¢eld of àmphibolitic' assemblag-es at the expense of typical eclogitic assemblages,the latter de¢ned by the garnetômphacite pair.This can be easily explained once phase relation-ships are analyzed even in simpli¢ed model che-mographies (see below).

The abundance of carbonates (0^20 oxygen%)increases with increasing pressure and decreasingtemperature. This is consistent with the positivedP/dT slope for all decarbonation reactions sug-

Table 4Fluid estimates by mass-balance calculations

Runnumber

Bulk T(³C)

P(GPa)

Fe3�/FeT X CO2 Nfluid

(mol%)

R1 HFB 665 2.0 0 0.02 15.3R1 HFB 665 2.0 0.5 0.02 12.8R3 HFB 680 1.2 0 0.35 15.8R3 HFB 680 1.2 0.5 0.42 13.4R10 HFB 680 1.4 0 0.36 18.3R10 HFB 680 1.4 0.5 0.43 16.1R2 HFB 680 1.6 0 0.03 11.0R2 HFB 680 1.6 0.5 0.03 8.6R9 HFB 680 1.8 0 0.04 11.5R9 HFB 680 1.8 0.5 0.05 9.1R6 HFB 730 1.2 0 0.39 11.7R6 HFB 730 1.2 0.5 0.52 9.1R7 HFB 730 1.4 0 0.30 15.4R7 HFB 730 1.4 0.5 0.37 13.1R4 HFB 730 1.6 0 0.22 9.8R4 HFB 730 1.6 0.5 0.31 7.1R8 HFB 730 1.8 0 0.13 11.2R8 HFB 730 1.8 0.5 0.20 8.7R5 HFB 730 2.0 0 0.16 13.7R5 HFB 730 2.0 0.5 0.21 11.2R1 OTB 665 2.0 0 0.05 9.1R1 OTB 665 2.0 0.5 0.05 7.4R3 OTB 680 1.2 0 0.49 12.4R3 OTB 680 1.2 0.5 0.56 10.9R10 OTB 680 1.4 0 0.44 6.1R10 OTB 680 1.4 0.5 0.62 4.4R2 OTB 680 1.6 0 0.20 9.3R2 OTB 680 1.6 0.5 0.24 7.7R9 OTB 680 1.8 0 0.05 6.2R9 OTB 680 1.8 0.5 0.07 4.6R6 OTB 730 1.2 0 0.49 12R6 OTB 730 1.2 0.5 0.57 10.5R7 OTB 730 1.4 0 0.50 11.7R7 OTB 730 1.4 0.5 0.58 10.3R4 OTB 730 1.6 0 0.25 11.1R4 OTB 730 1.6 0.5 0.29 9.7R8 OTB 730 1.8 0 0.05 8.6R8 OTB 730 1.8 0.5 0.06 7.1R5B OTB 730 2.0 0 0.10 12.5R5B OTB 730 2.0 0.1 0.19 12.2R1 HMB 665 2.0 0.1 0.00 9.6R1 HMB 665 2.0 0.4 0.42 12.8R3 HMB 680 1.2 0 0.56 14.1R3 HMB 680 1.2 0.5 0.61 13R10 HMB 680 1.4 0 0.40 9.4R10 HMB 680 1.4 0.5 0.47 8.3R2 HMB 680 1.6 0 0.11 6.5R2 HMB 680 1.6 0.5 0.14 5.3R9 HMB 680 1.8 0.5 0.02 6.8R6 HMB 730 1.2 0 0.62 9.9R6 HMB 730 1.2 0.5 0.72 8.6R7 HMB 730 1.4 0 0.59 13.2

Table 4 (continued)

Runnumber

Bulk T(³C)

P(GPa)

Fe3�/FeT X CO2 Nfluid

(mol%)

R7 HMB 730 1.4 0.5 0.65 12.1R4 HMB 730 1.6 0 0.60 12.2R4 HMB 730 1.6 0.5 0.67 11R8 HMB 730 1.8 0 0.28 10.3R8 HMB 730 1.8 0.5 0.32 9.2R5 HMB 730 2.0 0.6 0.32 5.8

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gested by experiments both in simple model sys-tems ([20] and references therein) and on basalts[22]. The abundance of quartz also increases withincreasing pressure and decreasing temperature,whereas paragonite is very abundant in the OTBcomposition at low temperature conditions (10^20oxygen%), contrasting with the lower abundance(V5 oxygen%) reported in the H2O-saturated ex-periments on the LT composition [26]. This canbe explained by the increase of the abundance ofCa^Mg carbonates, which favor an enrichment inSi, Al, Na and H2O in the silicate portion of theassemblage.

Table 4 and Fig. 5a,b illustrate the variation ofcomputed £uid compositions with pressure, tem-perature and bulk composition via a mass-balanceapproach. A systematic decrease of CO2 contentin the £uid with increasing pressure and decreas-ing temperature is coupled to the increasingamount of carbonates and to the decrease of hy-drates in the solid assemblage. Furthermore thedata obtained at 730³C may suggest a decreaseof CO2 in the £uid as a result of a decrease inthe Mg/(Mg+Fe) ratio of the bulk composition.

It should be emphasized that the compositionof the coexisting phases and their proportions area function of the amount of OAD added to thestarting material. Fig. 6 is a schematic chemogra-phy where anhydrous components plot at the topof the triangle and volatile components at thebase. Solid assemblages including hydrates andcarbonates will plot in the upper portion of thechemography. Analyzed solid-phase compositionsand calculated £uid compositions reveal a rota-tion of £uid^solid tie lines leading to carbonate-rich solid assemblages and H2O-rich £uids at highpressure. Fig. 6 shows that the intersection ofequilibrium tie lines with mixtures of starting ma-terials at variable contents of OAD in the bulkcomposition lead to di¡erent solid-phase compo-sitions and proportions, as well as di¡erent £uidcompositions. As a consequence, even though theamount of OAD added to the experimentalcharges is always close to 10 wt%, variations ofinitial weight proportions in the order of 2^3 wt%(Table 2) are responsible for some of the scatter-ing in the data of Figs. 4 and 5. In conclusion,even though the general trend depicted in Fig.

Fig. 4. Abundances of mineral phases (oxygen%) in the experimental runs calculated by non-linear least squared methods. Inter-polated curves are schematic.

EPSL 5381 29-2-00


4a,b is well constrained and further corroboratedby thermodynamic analysis presented in the fol-lowing section, application to natural rocksshould be undertaken with caution as the compo-sitions of the phases and of the coexisting £uidsare strongly dependent of the bulk composition ofthe system.

5. Fluid composition: thermodynamic approach

Knowledge of the composition of the coexistingsolid phases at known pressure and temperatureconditions allow us to calculate the compositionof the coexisting £uid, solving the equation:

0 �Xj

P�1

X P WW P �R T ln �XCO2 WQCO2�

where XP and WP are the reaction coe¤cient andchemical potential of phase, respectively.

Because of the poor knowledge of thermody-namic mixing properties of amphibole, only reac-tions involving garnet, carbonates, plagioclaseand epidote were used to compute the X CO2 ofthe £uid (Table 5). The calculations were per-formed using the thermodynamic data by Hollandand Powell ([34] and references therein), the activ-ity model for garnet by Berman [35], for clinopyr-oxene by Holland [36], and for plagioclase byFurhman and Lindsley [37], and assuming anideal model for epidote. Error propagation (1cstandard deviations) based on uncertainties ofthe enthalpy of formation [34] and of mineralcompositions was performed through a MonteCarlo method [38].

Despite the simpli¢ed approach used here,which ignores the presence of the most abundantphase, i.e. amphibole, the resulting values of X CO2

closely coincide with the values de¢ned by mass-balance calculations (Fig. 5c,d). Displacement ofcomputed £uid compositions at 1.2 GPa and730³C towards lower X CO2 may be ascribed tothe assumption of ideality of epidote solid solu-tion. Both mass-balance and thermodynamic cal-culations suggest that the maximum X CO2 ob-tained in this study occurs at V1.2^1.4 GPaand 730³C and X CO2W0.5. Conversely, a nearlypure hydrous £uid is expected at 2.0 GPa and665^680³C, close to amphibole breakdown.

The coexistence of carbonates with H2O-rich£uids at high pressure and low temperature con-ditions has also been predicted in several simpli-

Fig. 6. Schematic compositional diagram A^H2O^CO2 whereA represents non-volatile components. Fluids plot at thebase of this chemography, volatile-free phases at the top,and hydrate- and/or carbonate-bearing assemblages plot closeto the top of the triangle. Starting materials (bulk) composedof mixtures of ¢red gels and OAD lie close to the verticalline. Equilibrium assemblage and £uid composition are afunction of the amount of OAD used in the experiments.

Fig. 5. Estimates of X CO2 content in coexisting £uids ob-tained by (a,b) mass-balance and (c,d) mineral-£uid equilibri-um calculation. Squares: experiments on the compositionHFB; triangles: experiments on the composition OTB; dia-monds: experiments on the composition HMB. Open sym-bols are estimates at 665^680³C; ¢lled symbols at 730³C.

EPSL 5381 29-2-00


¢ed systems, e.g. in the systems CaOÂl2O3^SiO2^H2O^CO2 [39] and in the system CaO^MgO^SiO2^H2O^CO2 [19]. Such experimentaland theoretical predictions are further con¢rmedby the composition of £uid inclusions in high-pressure ma¢cs [40], by the estimation of £uidcomposition via phase equilibrium calculationson natural carbonate-bearing assemblages [41],and by stable isotope geochemistry on eclogitesand amphibolites [42].

6. Carbonate and hydrate stability in eclogitefacies metabasalts

As previously stated, the dominant, but unex-pected, e¡ect of addition of CO2 to the system is¢rst the displacement of the appearance of clino-pyroxene to much higher pressures than observedin H2O-saturated systems, and second the persis-tence of hydrous phases, namely amphibole, but

also paragonite and epidote. A complete and rig-orous treatment of phase relationships would re-quire a compositional space including nine com-ponents. However, if aluminous phases (garnet,paragonite, kyanite, clinozoisite) are ignored forthe sake of simplicity, most features can be easilyexplained by the chemography CaO^MgO^H2O^CO2 projected from SiO2 (Fig. 7). In this simpli-¢ed tetrahedron, amphibole solid solution is rep-resented by tremolite, and clinopyroxene by diop-side. Tremolite and dolomite plot on the facesCaO^MgO^H2O and CaO^MgO^CO2, respec-tively, whereas calcite and magnesite lie on theedges CaO^CO2 and CaO^MgO, respectively.The composition for a tholeiitic basalt+10 wt%OAD lies close to the line which connects trem-olite to the point H2O:CO2 = 1.

Because the system has four components andsix phases and the assemblage calcite+magnesiteis forbidden by currently available experimentaland thermodynamic data, only two non-degener-ated reactions are possible (indicated by the ab-sent phases) :

tremolite� calcite �

diopside� dolomite� quartz� fluid �Mag�

tremolite� dolomite� quartz �

diopside�magnesite� fluid �Cc�

Due to the colinearity between diopside, dolomiteand CO2, the degenerate reaction:

dolomite� quartz �

diopside� 2 CO2 �Tr; Cc; Mag�

is also possible.At relatively low pressure conditions (Fig. 7,

tetrahedron I) the assemblage tremolite+calciteþ dolomite coexists with a CO2-rich £uid. Thecompositional barrier de¢ned by tremolite+dolo-mite+calcite is responsible for the absence of di-opside in the basalt+10% OAD composition.With increasing pressure the displacement of theequilibrium £uid composition toward the H2O

Table 5Fluid estimates by mineral-£uid equilibrium calculations

Runnumber

Bulk T(³C)

P(GPa)

Reactiona X CO2b

R1 HFB 665 2.0 2 0.04 (0.01)R2 HFB 680 1.6 1 0.12 (0.03)R6 HFB 730 1.2 5 0.26 (0.05)R7 HFB 730 1.4 6 0.35 (0.10)R8 HFB 730 1.8 2 0.19 (0.04)R5 HFB 730 2.0 2 0.14 (0.02)R6 OTB 730 1.2 4 0.24 (0.18)R7 OTB 730 1.4 4 0.44 (0.22)R4 OTB 730 1.6 3 0.43 (0.05)R8 OTB 730 1.8 3 0.22 (0.05)R5B OTB 730 2.0 3 0.12 (0.02)R1 HMB 665 2.0 3 0.04 (0.01)R2 HMB 680 1.6 3 0.13 (0.02)R9 HMB 680 1.8 3 0.08 (0.01)R6 HMB 730 1.2 4 0.22 (0.17)R7 HMB 730 1.4 5 0.14 (0.02)R8 HMB 730 1.8 3 0.19 (0.04)R5 HMB 730 2.0 3 0.10 (0.01)aReactions employed in mineral-£uid equilibria calculations:1, 2 Gro+1 Pyr+6 CO2 = 3 Dol+3 An+3 Q; 2, Di+2CO2 = Dol+2 Q; 3, Pyr+Gro+6 CO2 = 2 Ky+3 Dol+4 Q;4, 2 Cz+CO2 = 3 An+Cc+H2O; 5, Di+2 CO2 = Cc+Mag+2Q; 6, Gro+2 CO2 = An+Q+Cc.bNumbers in parentheses are 1c standard deviations calcu-lated after 50 iterations.

EPSL 5381 29-2-00


component leads to a progressive increase in theamount of dolomite and ¢nally to the disappear-ance of calcite and the appearance of magnesite inthe experimental charges. Contemporaneously, re-action [Mag] causes the break of the composition-al barrier and assemblages bearing diopside+calcite+dolomite are expected in H2O- and/orCaO-enriched bulk compositions (Fig. 7, tetra-hedron II). At bulk compositions where CO2/(H2O+CO2) = 0, the stable assemblage is tremo-lite+diopside+water. In contrast, the clinopyrox-ene-free assemblage tremolite+magnesite+dolomi-te+£uid develops in the starting materials used.Again, a compositional barrier de¢ned by tremo-lite+dolomite+£uid is responsible for the absenceof diopside. It is therefore only at the highestpressures investigated (Fig. 7, tetrahedron III)that reaction [Cc] stabilizes the plane diopside+magnesite+£uid and clinopyroxene becomes ubiq-uitous in most rock compositions.

The occurrence of calcite/aragonite at relativelylow pressure, followed by dolomite and then bymagnesite at increasing pressure, is consistentboth with the experiments of Yaxley and Green[22] and with natural assemblages from high-pres-sure and ultra-high-pressure terranes [14^16].Magnesite in coesite eclogites from the Dabiemetamorphic terrane [16] is followed by dolomiteduring retrogression and similar relations havebeen described even in carbonate-bearing ultra-ma¢c systems mainly as a result of the reaction:

diopside� 2 magnesite � dolomite� enstatite

The arrangement of this reaction sequence in apressure^temperature diagram has been calculatedby vertex [43] using the thermodynamic data byHolland and Powell [34] (Fig. 8). Tremolite+cal-cite and tremolite+dolomite+quartz lie on the lowtemperature side of the univariant reactions [Mag]and [Cc], respectively. These univariant reactionsshow negative dP/dT slopes at high pressure butthe continuous shift of £uid composition towardsCO2-rich compositions with decreasing pressure isresponsible for backbending and positive dP/dTat low pressure. As tremolite is the only hydrousphase considered, its reaction coe¤cient vanishesat singular points S1 and S2, where the £uid iscomposed of pure CO2.

Even though the calculated pressure^tempera-ture diagram cannot be quantitatively used to un-ravel the evolution of carbonate-bearing eclogites,it con¢rms the sequence of assemblages observedexperimentally with increasing pressure and thecontinuous H2O enrichment in the £uid at highpressure when low geothermal gradients are at-tained.

The extrapolation of phase relations in the sys-tem CMS^H2O^CO2 to the Fe-bearing system isstraightforward, because addition of iron simplycauses an increase in the variance of the mineralassemblages. In contrast, addition of Al and Nacauses the presence of plagioclase, epidote, garnet,paragonite and kyanite. Nevertheless, reactionsde¢ned in the subsystem CMS^H2O^CO2 systemare still of relevance, unless the addition of Aland Na in amphibole and clinopyroxene is re-

Fig. 7. Phase relations among tremolite, diopside, calcite, dolomite, magnesite and H2O^CO2 mixed £uids in the tetrahedronCaO^MgO^H2O^CO2 projected from quartz. Numbers (I, II and III) refer to relative P^T conditions in the diagram of Fig. 8.

EPSL 5381 29-2-00


sponsible for inversions of reaction coe¤cients.In such a case, singular-point nets would occur[44]. The consistency between phase relations inthe CMS^H2O^CO2 system and the experi-mental results suggests that phase relations mod-eled in simple chemical systems are appropriateto understand the evolution of extremely com-plex systems such as carbonate-bearing metaba-salts.

7. The contribution of altered ma¢c volcanics tothe recycling of CO2 at convergent margins

The decarbonation of altered metabasalts hasbeen considered as a possible mechanism in orderto explain CO2 £uxes at convergent margins[10,19]. Because devolatilization reactions in Fig.8 have a £at positive dP/dT slope at low pressure,and because decarbonation is feasible only on thehigh-temperature and/or low-pressure portion ofthese reactions, the release of CO2-rich £uids fromsubducted altered metabasalts is unlikely for mostthermal regimes predicted by numerical models[21] of nearly steady-state subduction. Both theexperimental results presented here and the con-clusions from the supersolidus experiments ofYaxley and Green [22] suggest that carbonatestability at high pressure strongly promotes a £uidfractionation towards H2O-rich compositions.Even for early subduction of a young oceaniccrust at slow convergent rates, the hottest possiblescenario for a convergent margin (see Fig. 9 in[21]), the maximum X CO2 at V70 km depthdoes not exceed 0.2 and it is expected to approacha null value at pressures of the amphibole break-down reaction. It is therefore probable that themain source for CO2 of arc magmas is not thema¢c portion of the oceanic crust but the carbo-nates of marine sediments [2].

Nevertheless, when hot geotherms are assumed,relatively CO2-rich £uids can be transferred fromthe altered oceanic crust to shallower reservoirs inthe forearc region. Here, either volatiles migratetoward the surface, e.g. via upwelling of serpen-tine bodies [45], or CO2 is entrapped in ultra-ma¢cs, and downdragged by the mantle £ow inthe wedge region. It is then at higher pressures

and temperatures, in the hot portion of the man-tle wedge, that such CO2-enriched peridotite maycontribute to arc magmatism. However, it shouldbe noted that evidence of carbonated £uids/meltsin sub-arc environments is rare [46,47] and relatedto unusual tectonic settings. Nevertheless, thismechanism can account for the CO2 enrichmentof lithospheric mantle on a long-term scale and itmay explain the occurrence of carbonates in peri-dotite xenoliths [48] as well as in some campto-nitic lamprophyres [49], for cases where a carbo-natite metasomatism is disregarded on the basis ofgeochemical arguments.

The most likely fate of the CO2 componentstored in altered oceanic metabasalts is thereforethe almost complete fractionation into the solid,i.e. in dolomite and/or magnesite. The huge stabil-ity ¢eld of dolomite to more than 20 GPa and ofmagnesite down to lower mantle pressures [50], attemperatures far higher than 1000³C (even atpressures as low as 2.0 GPa), suggests a process

Fig. 8. Computed P^T diagram in the SiO2-saturated portionof the system CMS^H2O^CO2 to 2.5 GPa using the thermo-dynamic data base and the equation of state for H2O andCO2 of Holland and Powell [34]. Numbers (I, II and III)stand for chemographies in Fig. 7. Numbers (0.0, 0.5, 1.0)represent the CO2 fraction in the £uids along univariant(thick solid lines) and divariant (thin dashed lines) equilibria.

EPSL 5381 29-2-00


of deep mantle recycling of carbonates in mostsubduction zones. Such a process is responsiblefor the global-scale excess of subducted CO2

compared to CO2 released via arc magmatism[18].

8. Conclusions

The experimental determination of subsolidusphase relationships in three H2O^CO2-bearing ba-saltic systems at pressures to 2.0 GPa reveals thathydrates such as amphibole, paragonite, and epi-dote coexist with carbonates, i.e. calcite, dolomiteand magnesite, over a wide range of pressure andtemperature conditions typical for the formationof eclogite-facies rocks. Calcite is the stable car-bonate in the lower pressure range investigated(P91.4 GPa in tholeiite), its stability ¢eld is en-larged in Fe-rich bulk compositions. Dolomite-and dolomite+magnesite-bearing assemblages areformed with increasing pressures.

Mass-balance and thermodynamic calculationsof the composition of coexisting £uids demon-strate the progressive displacement toward H2O-rich compositions with increasing pressure, lead-ing to nearly pure aqueous £uids at 2.0 GPa and665³C. The stability of Ca^Mg carbonates and theH2O enrichment of the £uid at high pressure areresponsible for the absence of clinopyroxene tomore than 1.8 GPa in all basaltic compositions.As a consequence a `carbonateâmphibolite' de-velops within the eclogite facies P^T ¢eld.

Devolatilization reactions in this system are de-carbonation reactions at relatively low pressureand high temperature conditions, but they are de-hydration reactions at high pressure and low tem-perature conditions. As a consequence, the releaseof CO2-rich £uids from subducted altered oceanicbasalts is unlikely. Only hot thermal ¢elds (e.g. inthe case of early subduction of young oceaniccrust at slow convergent rates) allow for signi¢-cant CO2 transfer to shallow reservoirs, mainly inthe forearc region.

The global-scale imbalance of CO2 £uxes atconvergent margins is related to the deep recy-cling of CO2 via dolomite+magnesite-bearingmetabasalts and metagabbros.

Acknowledgements

We would like to thank Hubert Staudigel forhis review of the manuscript. J.-F.M. was sup-ported by the Spanish DGICYT, Project PB96-1266, and by a Spanish FPI postdoctoral grant.All experimental and analytical work was fundedby CNR (CSGAQ)-MURST (Co¢n98). The helpof Christine Laurin in improving the English ofan earlier version of this paper is gratefully ac-knowledged.[EB]

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