Contrib Mineral Petrol (1988) 100:34%360 Contributions to Mineralogy and Petrology �9 Springer-Verlag 1988

Granulite facies amphibole and biotite equilibria, and calculated peak-metamorphic water activities William M. Lamb 1 and John W. Valley 2

Department of Geology, Texas A & M University, College Station, TX 77843, USA 2 Department of Geology and Geophysics, University of Wisconsin, Madison, WI 53706, USA

Abstract. Samples located near the Oregon Dome anortho- site massif in the south-central Adirondack Mountains, New York contain the fluid-buffering mineral assemblages: amphibole + clinopyroxene + orthopyroxene + quartz or biotite + quartz + orthopyroxene + K-feldspar. These rocks were metamorphosed under granulite-facies conditions (T=725 ~ 750 ~ C, P=7 .5 kbar) during the Grenville oro- geny. The Mg-rich nature of amphiboles, micas, and pyrox- enes allow accurate calculation of water activities because corrections for the effects of solid solution are relatively small. The activity of water was low during the peak of granulite-facies metamorphism, with e H20 ~ 0.15 _+ 0.14. Wollastonite occurrences indicate that the c~CO2 was low (< 0.3) in nearby rocks, demonstrating that large quantities of CO2 did not infiltrate in a pervasive manner. The combi- nation of low ~HzO with low ~C02 is consistent with the hypothesis that magmatic processes were dominant, gener- ating dry, fluid-absent conditions.

Introduction

Low values of J t t 2 0 are a universal characteristic of granu- lite-facies metamorphism, and understanding the causes of dehydration in granulites is frequently cited as the most important problem in deep crustal metamorphism. Reduced f H 2 0 may result from a variety of processes. For example, partial melting can lower fHzO, if H20 is partitioned into the melt which may then be removed by intrusion to shal- lower crustal levels (Fyfe 1973). If the magma is not extract- ed then H20 will be released after the peak of metamor- phism, possibly causing retrograde hydration. Other pro- cesses that could reduce H20 fugacities are: (1) infiltration of other fluids, particularly CO2 (Newton et al. 1980; Con- die et al. 1982; Hansen et al. 1984; Friend 1985; Touret 1985; Newton 1986), (2) metamorphism of rocks that inher- it low f H 2 0 (examples may include both previously meta- morphosed rocks and orthogneisses; Valley 1985), and (3) passage of magmas through the crust (Frost and Frost 1987). Clearly, CO2-infiltration requires the presence of a peak-metamorphic fluid that is both pervasive, and CO2-

Abbreviations: f i Fugacity of species i in a fluid; Xi mole fraction of component i in a phase; T temperature; P lithostatic pressure; Pv fluid pressure; c~7 activity of component i phase X

Offprint requests to: W.M. Lamb

rich. In contrast, magmatic processes can generate either fluid absent conditions or a CO2-rich peak metamorphic fluid (CO2-enrichment may result because the solubility of CO2 in magmas is low relative to H20). Quantification of f H 2 0 will aid in distinguishing among these various models for the genesis of granulites. Furthermore, the pres- ence or absence of fluids in the deep crust has wide signifi- cance in terms of metamorphism, heat transport, metaso- matism, and the seismic character of presently active meta- morphic belts.

In granulites, low f H 2 0 is inferred from mineral as- semblages, particularly those containing orthopyroxene and/or unmelted lithologies, but quantitative determina- tions are rare. Previous studies have estimated that the mole fractions of H20 in peak metamorphic fluids range from 0.01 to 0.6 (Phillips 1980; Bohlen et al. 1980; Valley et al. 1983; Hansen et al. 1984; Powers and Bohlen 1985; Bhatta- charya and Sen 1986; Erhard 1986; Seal 1986; Edwards and Essene 1988). However, estimation offHzO generally involves large uncertainties that arise, in part, from the lack of adequately calibrated activity composition rela- tions for the amphibole and biotite compositions that are commonly found in granulites.

In this study we report on unusually Mg-rich gneisses from the southern Adirondack Mountains, New York, that contain fluid-buffering mineral assemblages. The Mg-rich nature of amphiboles, micas, clinopyroxenes and orthopy- roxenes in some of these rocks allows the calculation of f H 2 0 with high accuracy. Graphite in two samples permits calculation of the composition of fluids in the C - O - H - F system. All samples also contain coexisting orthopyroxene and clinopyroxene, a potential geothermometer, and results from the application of two-pyroxene thermometry are re- ported in Lamb (1987).

Adirondack geology and sample locations

The Adirondack Mountains form a southeastern outlier of the Grenville Tectonic Province and were metamorphosed during the Grenville orogeny (,,~1.1 Ga). This terrane is composed of meta- sedimentary and metaigneous rocks which surround several meta- morphosed anorthosite bodies.

The samples analyzed in this study were taken from outcrops located near the Oregon Dome anorthosite. Horizontal distances to the anorthosite range from 8 km to <1 m (Fig. I). Table one lists the mineralogy of the samples and abbreviations of mineral names. Accurate sample locations are given by Lamb (1983). Meta- morphic temperatures, determined largely by application of two-

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Fig. 1. Map of the Oregon Dome region of the Adirondack Moun- tains, New York. Sample locations are shown as solid dots and anorthosite bodies are ruled

feldspar and Fe -T i -ox ide geothermometers, increase from 725 ~ C___30 ~ C in the southern part of Fig. 1 to 750 ~ C_+30 ~ C in the north, at a pressure of 7.5 _+0.5 kbar (Bohlen et al. 1985).

The geology of this region has been described by McLelland and Isachsen (1980). Many of the rocks are metasediments (e.g., marbles, calc-silicates, and pelites). However, the protolith for some lithologic units is uncertain, and may be either igneous or sedimen- tary. Most samples (10 of 12) have high whole rock Mg/Fe and thus contain Mg-rich micas, amphiboles, and pyroxenes. Many of the Mg-rich samples (8 of 10) contain a tremolitic hornblende; however, in two of the samples the only hydrous mineral is phlogo- pite. All Mg-rich samples, with the exception of 81-TL-2, were collected from non-carbonate-bearing portions of the Blue Moun- tain Lake Formation, a marble-rich unit that also contains quart- zite, amphibolite, and granitic gneisses. Sample 81-TL-2 is also from a non-carbonate-bearing portion of a marble-rich unit, the Cedar River Formation, which also contains amphibolite, quart- zite, and sillimanitic garnet-biotite gneiss. The two samples that

do not contain Mg-rich minerals are: (1) 81-SP-1, from a gabbroic dike that crosscuts the anorthosite, and (2) 81-SP-40, and orthopy- roxene-bearing (charnockitic) gneiss from the Lake Durant Forma- tion (McLelland and Isachsen 1980).

Analytical procedure

Analyses were made by electron microprobe (Tables 2 through 6) in order to evaluate the effect of solid solution on fluid-buffering mineral equilibria. Major and minor elements were analyzed using an ETEC Autoprobe (Rice University) and an ARL-SEMQ micro- probe (University of Wisconsin-Madison) with wavelength-disper- sive LiF, PET and TAP crystal spectrometers. The accelerating potential was 15 keV except for analysis of F and C1 when 10 keV was employed to avoid volatilization. Sample currents ranged from 0.015 to 0.020 IxA. Analyses are accurate to _+3% of the amount present for major elements and _+ 5-10% for minor elements (< 2 wt%). All minerals were checked for the presence of unanalyzed elements (Z> 10, >0.2 oxide weight percent) by energy dispersive analysis (500 to 1000 s).

The original analyses of these minerals used topaz as the F standard. However, this procedure resulted in systematic errors in F content, presumably due to differences in the shape of the FKe-peak between topaz and phlogopite (Solberg 1982). All F analyses reported here are relative to synthetic fluorphlogopite.

Pyroxene formulae were normalized to four cations with Fe 3 + inferred from charge balance, Fe 3 + = AI TM- A1 w - 2 Ti + Na. Am- phibole formulae were normalized to 13 cations (Ma, M2, M3, and IV sites), with Fe3+=AIIV--AlVl--2Ti--NaA-KA+Na M4. Trioctahedral micas were normalized to seven VI + IV cations with Fe 3 + = B A + (A1 w - 1 ) - A1 vl - 2 Ti -- Ca A. This normalization scheme assumes no oxymica component or octahedral vacancies. Feldspar analyses were normalized to five cations. The hydrogen content of a biotite was measured by fusion at 1500 ~ C, oxidation of evolved gases over CuO at 550 ~ C, cryogenic separation of HaO, reduction of HaO to Hz with Zn metal at 450 ~ C, and manometric measurement of H2.

Stable isotopic analyses were performed on a Finnigan/MAT 251 mass-spectrometer (University of Wisconsin-Madison). Extrac- tion of oxygen from silicates was performed using BrF5 (Clayton and Mayeda 1963). Pyrrhotite and graphite were combusted in the presence of excess 02 at 1150 ~ C and 1000 ~ C, respectively.

Origin of magnesian gneisses

Analyses of coexisting minerals show that 10 samples contain the assemblage Mg-rich amphibole or phlogopite + diopside + enstatite (see Tables 2-6). Thus, regardless of the modal abundances of these

Table 1. Mineralogy of samples examined in this study

SL Am Cpx Opx Qtz Bi Kfs P1 Po Ilm Other RT

SP-79-9 2 X X X X X X X A Sp-79-I 1 1 X X X X X X A SP-79-12 1 X X X X X X A SP-79-19 1 X X X X A SP-79-20 1 X X X X A 81-SP-13 3 X X X X Sph A 81-TL-2 10 X X X X X X A JM-TRD 1 X X X X A 82-SP-5a 4 X X X X X Gr A 82-SP-51 9 X X X X X Gr A 81-SP-I 6 X X X X X X X X B 81-SP-40 7 X X X X X X X X X Mag/Ap C

Abbreviations: Am Amphibole, Ap Apatite, Bi Biotite, Cpx Clinopyroxene, Di Diopside, En Enstatite, Gr Graphite, Ilm Ilmenite, Kfs K-feldspar, Mag Magnetite, Opx Orthopyroxene, Ph Phlogopite, Pl Plagioclase, Po Pyrrhotite, Qtz Quartz, Sa Sanidine, Sph Sphene, Tr Tremolite, SL Sample Location, corresponding to locations on Fig. 1, RT Rock Type: A Sample with high M g / M g + Fe (see text); B Gabbroic Dike, crosscutting anorthosite; C charnockitic gneiss (see text)

Table 2. Amphibole analyses

SP-79-9 SP-79-11 SP-79-19 SP-79-20 81-SP-13 81-TL-2 JM-TRD a 81-SP-1 81-SP-40

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SiO2 51.36 51.75 51.98 50.83 51.62 51.45 53.00 45.84 39.85 TiO2 1.27 1.33 0.92 1.47 0.72 1.30 0.71 0.85 2.08 A1203 8.39 8.22 6.67 8.26 7.39 7.58 7.29 9.45 11.92 Fe203 b 2.73 0.80 0.13 0.51 2.36 1.69 0.00 4.48 6.40 FeO 0.00 0.00 0.00 0.00 0.00 0.00 0.78 12.58 14.12 MnO 0.13 0.14 0.12 0.15 0.08 0.10 <0.06 0.15 0.28 MgO 21.17 21.64 22.98 22.62 22.58 21.26 21.01 11.36 8.41 CaO 12.86 12.66 13.17 13.00 12.31 12.56 12.91 12.01 11.17 Na20 0.98 1.31 1.20 1.37 1.21 1.02 0.82 0.47 1.23 K20 0.41 0.41 0.41 0.49 0.71 0.44 0.75 1.19 2.01 F 0.60 0.68 0.73 0.60 1.23 0.49 0.69 0.12 0.22 C1 < 0.05 < 0.05 < 0.05 < 0.05 < 0.05 < 0.05 < 0.05 < 0.05 0.73 HzO b 1.92 1.87 1.84 1.94 1.64 1.94 1.85 1.90 1.66

Total 101.82 101.81 100.15 101.24 101.85 99.83 99.81 100.40 100.08 - - O = F + C I 0.25 0.29 0.31 0.25 0.52 0.21 0.29 0.05 0.25 Total 101.57 100.52 99.84 100.99 101.33 99.62 99.52 100.35 99.83

Formulae normalized to 13 I V + V I cations (see text)

Si 5.960 7.051 7.115 6.891 6.962 7.084 7.325 8.728 5.104 A1 Wb 1.040 0.949 0.885 1.109 1.038 0.916 0.675 1.272 1.896 A1 vl 0.300 0.371 0.191 0.210 0.137 0.314 0.510 0.456 0.257 Ti 0.129 0.137 0.094 0.150 0.073 0.135 0.073 0.094 0.240 Fe 3 + b 0.278 0.082 0.013 0.052 0.240 0.175 0.000 0.495 0.737 Fe 2 + 0.000 0.000 0.000 0.000 0.000 0.000 0.090 1.544 1.809 Mn 0.015 0.016 0.013 0.018 0.009 0.012 <0.006 0.019 0.037 Mg 4.277 4.394 4.689 4.570 4.541 4.364 4.327 2.485 1.920 Ca 1.867 1.849 1.931 1.888 1.780 1.854 1.912 1.888 1.832 Na vlb 0.133 0.151 0.069 0.112 0.220 0.146 0.088 0.112 0.168 NaA 0.126 0.194 0.249 0.249 0.097 0.128 0.t32 0.022 0.196 K 0.071 0.071 0.071 0.085 0.122 0.078 0.132 0.223 0.393 F 0.258 0.292 0.318 0.256 0.526 0.216 0.300 0.058 0.107 C1 < 0.006 < 0.006 < 0.006 < 0.006 < 0.006 < 0.006 < 0.006 < 0.006 0.185 OH b 1.742 1.708 I. 682 1.744 1.474 1.784 1.700 1.942 1.708

Mg Mg + Fe 0.939 0.982 0.997 0.989 0.950 0.961 0.980 0.549 0.430

a From Valley et al. 1983 b From normalized formula (see text)

minerals, many of these rocks combine very low Fe contents with M g / M g + C a > 0 . 5 (samples SP-79-20, 81-SP-13, JM-TRD, 82-SP- 5a, and 82-SP-51). The other Mg-rich samples contain plagioclase, and so may have higher Ca contents, al though estimates of modal abundances indicate that M g / M g + Ca is also > 0.5. This unusual bulk chemistry cannot result from isochemical metamorphism of a normal marine limestone or marl (calcite +dolomite) . Instead, such compositions are best explained as the result of: (1) metamor- phism of evaporites, or (2) metasomatism.

Evaporites are characterized by high Mg/Ca ratios (Moine et al. 1981), and so these rocks may originally have been part of an evaporite rich sequence. Support for an evaporitic origin is also inferred from the presence of tremolite granulites and talc + tremolite + quartz schists in the northwest Adirondacks that have a major element chemistry similar to the Mg-rich gneisses of this study. The schists are associated with halite and anhydrite-bearing marbles that are clearly metaevaporites (Brown and Engel 1956; Wiener et al. 1984). Furthermore, the Gouverneur Marble, which is the unit that contains the metamorphosed evaporites, is strati- graphically correlative with the Cedar River and Blue Mounta in Lake Formations, which contain the Mg-rich samples examined in this study (Wiener et al. 1984).

Oxygen isotopic ratios measured in two samples give further support to an evaporitic origin. Values of 6180 (quartz) for the two phlogopite-bearing samples (82-SP-5a and 82-SP-51) are high

(~21.5%o), and are similar both to Adirondack marbles (average 180 (Calcite)~20%o), and to the t a l c+ t r emol i t e+qua r t z schist

(24.5%o, Valley and O'Neil 1984). These " h e a v y " values indicate a sedimentary origin, are consistent with an evaporite protolith, and would not be expected if the rock was formed by metasoma- tism. Stable isotopic values from tremolite-bearing samples are less distinctive (~180=12 .4 to 14.6%0; Table 7). While these values are lighter than most calc-silicates, marbles, or evaporites (Hoefs 1987; Valley and O'Neil 1984), they are still likely to be metasedi- mentary in origin. Values of 634S in pyrrhotite (-5.1%o to -3.1%0 CDT) are lower than sulfates in evaporites (8%0 to 28%0 in the Balmat region, Whelan et al. 1984; 10%o to 30%0 in general for evaporites, Hoefs 1987), and lower than primary sulfides in Adiron- dack anorthosites (2.9%o-3.2%0, Morrison 1988), suggesting post- depositional modification. Nevertheless, these values are consistent with an evaporitic origin for these rocks and the high 6180 values rule out an origin by metasomatism during intrusion of the nearby anorthosite pluton.

Calculation of H 2 0 activities

M i n e r a l a s s emb lages t h a t bu f f e r ed p e a k m e t a m o r p h i c f H 2 0 by two equ i l ib r i a h a v e b e e n a n a l y z e d in s amples lo- ca t ed n e a r the O r e g o n D o m e a n o r t h o s i t e : a m p h i b o l e + or-

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Table 3. Mica analyses

82-SP-5a 82-SP-31 81-SP-1 81-SP-40

SiO 2 43.55 41.89 37.59 37.95 TiO2 2.09 2.14 4.50 4.05 A12Oa 12.68 14.30 14.27 13.76 Fe2Oa a 0.00 0.00 0.00 0.00 FeO 1.13 1.40 17.79 20.03 MgO 25.40 24.62 12.44 10.35 CaO 0.04 0.04 0.07 0.18 Na20 0.11 0.07 <0.05 <0.05 K20 10.12 9.84 9.81 8.69 F 2.08 1.40 0.52 0.69 C1 <0.10 <0.10 <0.10 0.18 H20 a 3.25 3.55 3.63 3.45

Total 100.45 99.05 100.62 99.33 -- O = F + C1 0.88 0.60 0.22 0.33 Total 99.57 98.45 100.40 99.00

Formulae normalized to 7 cations

Si 3.083 2.985 2.884 2.971 A1 Ira 0.917 1.015 1.116 1.029 A1 vI 0.141 0.186 0.175 0.241 Ti 0.111 0.114 0.260 0.239 Fe 3 + " 0.000 0.000 0.000 0.000 Fe 2+ 0.067 0.084 1.142 1.312 Mg 2.681 2.616 1.423 1.208 Ca 0.003 0.003 0.006 0.0l 5 Na 0.015 0.009 <0.008 <0.008 K 0.914 0.895 0.960 0.867 F 0.465 0.315 0.127 0.171 C1 <0.012 <0.012 <0.012 0.024 OH a 1.535 1.685 1.873 1.805

Mg Mg+Fe 0.976 0.969 0.555 0.479

a From normalized formula (see text)

thopyroxene + clinopyroxene + quartz; and biotite + quartz + orthopyroxene + K-feldspar. Calculation of f H 2 0 for these rocks is based on the equilibria:

Tremolite = Diopside + Enstatite + Quartz + H 2 0 (1) 2 Ca2MgsSisO22(OH)2 = 4 CaMgSi20 6 + 3 Mg2Si20 6 + 2 SiO 2 + 2 H 2 0

Phlogopite + Quartz = Enstatite + Sanidine + H 2 0 (2) 2 KMgaA1Si301 o(OH)2 + 6 SiOz = 3 Mg2Si20 6 + 2 KA1- Si308 + 2 H 2 0

Equilibrium (1) has been experimentally calibrated by Boyd (1959), Skippen (1970, and Yin and Greenwood (1983). Based on these experiments a temperature of 840~ at P H 2 0 = 1 . 4 kbar (Yin and Greenwood 1983) was chosen as the starting point for calculations. Equilibrium (2) has been experimentally reversed by Bohlen et al. (1983) at 790 ~ C, 5 kbar and X H 2 0 =0.35.

Figure 2 shows the P - T - a H 2 0 conditions for equilib- ria (1) and (2) that have been calculated using the computer program E Q U I L I (Wall and Essene 1972; see also Perkins et al. 1987) which we have modified to account for non- ideal mixing of H20-CO2 (Kerrick and Jacobs 1981). Ther- modynamic data are in Table 8. While some portions o f this diagram are metastable with respect to melting (Wones and Dodge 1977; Bohlen et al. 1983), this does not affect the calculation of water activities.

Accurate calculation of ~zH20 requires an assessment of the effect o f solid solution on equilibria (1) and (2). Ideal ionic activity models were employed for these calcula- tions (Kerrick and Darken 1975):

Arm = (XCaM4)2 (XMgMI ' 2, 3)5 (XOH)2 (3) Bi 3 ep~=(XK) (XMg) (XOH) 2 (4) Cpx M2

~ D i =(XCa ) (XMg M1) (5) Opx M2 c%n =(XMg ) (XMg ~I1) (6) Kfs

asa = K/(K + Ca + Na + Ba) (7)

The distribution of cations between the octahedral M I and M2 sites in pyroxenes is according to Wood and Banno (1973).

The small deviation o f the minerals in the Mg-rich sam- ples from end-member compositions and the large A G of equilibria (1) and (2) minimize the uncertainty arising from the application of activity models. Application of the activi- ty model o f Sharma et al. (1987) to the orthopyroxenes in this study produces negligible changes in calculated water activities for the samples containing Mg-rich minerals. However, when applied to the two samples with higher Fe contents (81-SP-1 and 81-SP-40) the calculated water activi- ties show a slight reduction (for example the e H 2 0 calculat- ed for 81-SP-1 from equilibrium (1) is reduced from 0.14 to 0.06).

The change in free energy due to reduced activities of the solid phases can be calculated from the relations:

AG=RTln K e q ( 8 )

Cpx 4 Opx 3 Am 2 K,q(1)=(~oi ) (~n~) /((XTr ) (9)

Kfs 2 Opx 3 Bi 2 K~q(2)=(~sa ) (aEn) /(~eh) (10)

for equilibria (1) and (2). The activities of the various miner- al phases are given in Table 9, as well as the calculated aHzO in the metamorphic fluid (0.07 to 0.18).

Other equil ibria and the es t imat ion o f water act ivit ies

Accurate calculation of peak metamorphic fluid composi- tions based on equilibria (1) and (2) requires experimental calibration, and such data are seldom perfectly consistent. Our calculations indicate that equilibria (1) and (2) intersect at T=910~ C and P=5.7 kbar for X H 2 0 = I . 0 (Fig. 3a). This intersection forms an invariant point that involves the equilibria:

Phlogopite + 4 Quartz + 2 Diopside = Tremolite + K-Felds- par (11)

Phlogopite + 3 Tremolite = 6 Diopside + 6 Enstatite + K- Feldspar + 4 H 2 0 (12).

This calculated placement and slope o f equilibria (11) differs markedly from that determined experimentally (Wones and Dodge 1977), although the calculations are very sensitive to small differences in A G (see the discussion in Valley and Essene 1980, p. 783). Experiments indicate: (1) equilibrium (11) has a vertical slope, (2) phlogopite + diop- s ide+quar tz is the high temperature assemblage, and (3) invariant point (A) should be located at P < 1 kbar. Inter- section of (1) with the experimentally determined P - - T lo- cation of (11) yields an invariant point at T = 750 ~ C and P = 300 bars (Fig. 3 b). Use of this low T, P invariant point to calculate the stability of equilibrium (2) yields approxi- mately 60 ~ higher temperatures (at 5 kbar) than if the rever-

Table 4. Clinopyroxene analyses

SP-79-9 SP-79-11 SP-79-12 SP-79-19 SP-79-20 81-SP-13 81-TL-2 JM-TRDa82-SP-5a 82-SP-51 81-SP-1 81-SP-40

353

SiO2 53.70 54.80 53.97 54.26 52.92 53.72 53.82 54.55 56.22 55.37 51.90 50.45 TiO2 0.18 0.21 0.25 0.11 0.26 0.21 0.27 0.17 <0.05 <0.05 0.28 0.27 A1203 1.92 2.28 1.60 1.78 2.01 1.97 2.00 1.99 0.50 1.08 1.40 1.71 Fe203 b 1.63 0.00 2.04 0.18 0.31 2.08 1.39 0.76 0.00 0.00 0.00 2.33 FeO 0.70 0.55 0.68 0.00 0.00 0.00 0.00 0.00 1.02 1.09 12.19 12.76 MnO 0.15 0.22 0.06 0.16 0.23 0.10 0.11 0.07 <0.05 0.14 0.22 0.69 MgO 18.01 17.93 17.43 18.50 18.71 18.86 17.98 18.84 17.92 18.40 11.13 11.14 CaO 23.61 25.58 24.35 25.02 24.55 23.87 24.68 24.91 23.99 23.81 21.95 20.]1 Na20 0.25 0.43 0.38 0.36 0.35 0.23 0.30 0.21 0.17 0.14 0.23 0.32

Total 100.15 100.01 100.76 100.37 99.34 101.04 100.55 101.51 99.92 100.03 99.30 99.78

Formulae normalized to 4 cations

Si 1.941 1.973 1.945 1.942 1.91] 1.918 1.934 1.934 2.034 1.998 1.982 1.932 A1 lvb 0.059 0.027 0.055 0.058 0.085 0.082 0.066 0.066 0.000 0.002 0.018 0.068 A1 w 0.023 0.070 0.013 0.017 0.000 0.001 0.019 0.017 0.026 0.044 0.045 0.009 Ti 0.005 0.006 0.007 0.003 0.007 0.006 0.007 0.005 <0.001 <0.001 0.008 0.008 Fe 3 + b 0.044 0.000 0.055 0.005 0.008 0.056 0.037 0.020 0.000 0.000 0.000 0.067 Fe z+ 0.021 0.017 0.020 0.000 0.000 0.000 0.000 0.000 0.031 0.033 0.390 0.409 Mn 0.005 0.005 0.002 0.005 0.007 0.003 0.003 0.002 < 0.002 0.004 0.007 0.022 Mg 0.970 0.962 0.936 0.987 1.007 1.004 0.963 0.997 0.967 0.989 0.634 0.636 Ca 0.914 0.910 0.940 0.959 0.950 0.914 0.950 0.945 0.930 0.920 0.899 0.825 Na 0.018 0.030 0.027 0.025 0.025 0.016 0.021 0.014 0.012 0.010 0.017 0.024

Mg M g § 0.937 0.983 0.926 0.995 0.992 0.947 0.963 0.980 0.969 0.968 0.619 0.572

a From Valley et al. 1983 b From normalized formula (see text)

Table 5. Orthopyroxene analyses

SP-79-9 SP-79-11 SP-79-12 SP-79-19 SP-79-20 81-SP-13 81-TL-2 JM-TRD" 82-SP-5a 82-SP-51 81-SP-1 81-SP-40

SiO2 57.82 59.74 57.59 58.79 57.38 56.17 58.20 58.81 59.95 59.24 52.40 50.03 TiO2 0.09 0.07 < 0.05 0.08 0.08 0.08 0.08 < 0.05 < 0.05 < 0.05 < 0.05 0.08 A1203 1.50 1.51 0.71 1.08 1.03 1.89 1.31 1.54 0.33 1.25 0.64 0.51 Fe203 b 0.62 0.00 1.01 0.46 1.45 2.82 0.00 0.00 0.00 0.00 0.00 0.64 FeO 6.42 1.74 7.65 0.00 0.19 2.94 4.71 2.63 3.29 3.58 30.54 31.28 MnO 0.27 0.46 0.21 0.40 0.48 0.17 0.23 0.18 0.16 0.18 0.62 1.68 MgO 34.72 37.34 34.4] 39.52 39.19 37.07 35.72 36.73 36.24 35.88 16.38 15.55 CaO 0.40 0.38 0.23 0.32 0.33 0.47 0.30 0.23 0.49 0.34 0.59 0.60

Total 101.84 101.24 101.81 100.65 100.13 1 0 1 . 6 1 100.55 100.12 100.46 100.47 101.17 100.37

Formulae normalized to 4 cations

Si 1.960 2.000 1.964 1.959 1.928 1.889 1.982 1.996 2.036 2.013 2.011 1.951 A1 tvb 0.040 0.000 0.029 0.041 0.041 0.075 0.018 0.004 0.000 0.000 0.000 0.023 A1 vt 0.020 0.060 0.000 0.001 0.000 0.000 0.034 0.056 0.013 0.050 0.029 0.000 Ti 0.002 0.002 <0.001 0.002 0.002 0.002 0.002 <0.001 <0.001 <0.001 <0.001 0.002 Fe 3 + b 0.016 0.000 0.026 0.012 0.037 0.071 0.000 0.000 0.000 0.000 0.000 0.019 Fe z+ 0.184 0.048 0.218 0.000 0.005 0.083 0.134 0.074 0.093 0.102 0.980 1.020 Mn 0.008 0.014 0.006 0.011 0.014 0.005 0.006 0.006 0.005 0.005 0.020 0.055 Mg 1.756 1.862 1.749 1.963 1.961 ].858 1.814 1.858 1.835 1.818 0.936 0.905 Ca 0.014 0.014 0.008 0.011 0.012 0.017 0.010 0.008 0.018 0.012 0.024 0.025

Mg M g + F e 0.898 0.975 0.878 0.994 0.979 0.923 0.931 0.962 0.952 0.947 0.489 0.466

From Valley et al. 1983 b From normalized formula (see text)

sal o f Boh len et al. (1983) is used. It is r eassur ing tha t even these seemingly large incons i s tenc ies be tween expe r imen ta l a n d t h e r m o c h e m i c a l d a t a yield very small unce r t a in t i e s o f 0.03 to 0.05 in es t imates o f c~H20 based on equ i l ib r ium (2).

Al-rich equilibria

The h igh A1 c o n t e n t o f the a m p h i b o l e s in this s tudy (6.67 to 11.92 w t % A1203) affects the stabil i ty a n d ca lcu la ted wa te r activit ies for the a m p h i b o l e - b e a r i n g samples . C a o

354

Table 6. K-Feldspar analyses

82-SP-5a 82-SP-51 81-SP-1 81-SP-40

SiOz 65.41 64.07 63.80 65.36 A1203 19.05 19.46 19.21 19.55 BaO 1.47 1.02 1.30 0.40 CaO 0.08 0.73 0.05 0.06 Na20 2.1t 1.19 0.87 0.78 K20 12.55 13.74 15.06 15.43

Total 100.67 100.21 100.29 101.58

Formulae normalized to 5 cations

Si 3.011 2.966 2.956 2.976 A1 1.034 1.062 1.049 1.049 Na 0.188 0.107 0.078 0.069 Ca 0.004 0.036 0.003 0.003 K 0.737 0.811 0.890 0.896 Ba 0.026 0.018 0.024 0.007

K K+Na+Ca+Ba 0.772 0.834 0.894 0.919

Table 7. Stable isotopic analyses

61sO (SMOW) Qtz 0348 (CDT) Po 613C (PDB) Gr

SP-79-9 13.9 SP-79-11 14.6 SP-79-19 13.6

13.3 SP-79-20 14.4 81-TL-2 12.7

12.4 82-SP-5a 21.6

21.4 82-SP-51 21.5

- 5 . 1 -3.1

-4.2

-14.4

-10.73

et al. (1986) have experimentally determined the stability of tremoliteso-tschermakit%o (Trso-Tsso). These experi- ments place broad constraints on the reaction

2 Trs 0 - Ts5 o = 2 Anorthite + 2 Diopside + 3 Enstatite + 2H20 (13),

and provide a qualitative estimate of c~H20. When applied to the Oregon Dome samples that contain plagioclase ( An3o to Anvs) the results indicate that c~H20 < 0.6, in sup- port of the other estimates.

Experiments in Fe-rich systems

Thus far, the estimation of ~H20 has been based on experi- ments performed in Fe-free systems. Use of these experi- ments is particularly appropriate given the Mg-rich nature of many of the samples. However, two samples that contain fluid buffering equilibria are more Fe-rich (81-SP-1 and 81-SP-40), and so experimental data on Fe-rich analogues of equilibria (1) and (2) would be useful in quantifying ~H20. While experimental data of this type are lacking for equilibrium (1), the stability of both biotite and annite have been investigated experimentally (Eugster and Wones 1962; Wones and Eugster 1965; Rutherford 1969; Hoffer and Grant 1980). Application of these experiments to the Fe-rich samples suggest that water activities were higher

Ol

o

ul"

W

9 r

7 ~-

5 -

- / / 3 7

600

o71ol ' /o.;5 / " ' / " / / /

/Ore~176176 Dome P-T

/ / /

/ /

/ , / . 700

I i 0.50 / / ",./

/ /

/ /

/ / /

800 go0 TEMPERATURE, ~

I i 1.00 /

/ "~/0 r

,k--

/ /

i

1000

Fig. 2. Pressure - temperature diagram showing equilibria (1) solid eurves and (2) dashed curves at aH20=0.10, 0.25, 0.5, and 1.0 (see text). The box represents the peak metamorphic conditions, and at these conditions the peak metamorphic c~H20 must have been considerably < 1.0 (see text)

(~H20=0.5) than those calculated from the Mg end- member reactions for the same samples (0.07 to 0.17, Table 9).

If we accept the experimental data on Fe-rich micas, then it might be tempting to conclude that these samples record local differences in H20 activities. However, incon- sistencies between experimental data on the breakdown of annite + quartz and the observed compositions of naturally coexisting biotite and orthopyroxene suggest that the exper- iments do not accurately model biotite and annite stability (Clemens et al. 1987). Thus, conclusions regarding water activities in natural samples that are based on experiments on Fe-bearing micas are premature.

U n c e r t a i n t i e s in c a l c u l a t e d ~ H 2 0

Tetrahedral order-disorder in micas

The calculated values of ~H20 based on equilibrium (2) assume tetrahedral A1-Si disorder in phlogopite. Thus, 18.70 J/mol-deg have been added to Sz9sK to account for this additional configurational entropy (Fyfe et al. 1958). The assumption that metamorphic phlogopites are disor- dered is supported by published structure refinements of phlogopite and biotite (Bailey 1975; Bohlen et al. 1980). While these refinements are consistent with disordered A1- Si, in none of these studies could short-range ordering with- in domains be ruled out. However, spectroscopic work sug- gests that biotites from high temperature (T< 600 ~ C) para- geneses are nearly completely A1-Si disordered (Clemens et al. 1987).

If some ordering in natural micas does occur, then the configurational contribution to entropy has been overesti- mated with the result that ~ H20 has been slightly underesti- mated. For example, if the configurational entropy contri- bution is reduced by 4 J/mol-deg, then the calculated ~HzO would increase by approximately 0.04 (i.e., from 0.10 to 0.14).

355

Table 8. Thermodynamic data for minerals used to calculate equilibria Entropy: S~ = S~98 + A • In T + B • 10- 3 T + C • 102 T- 2 + D (T in ~ S in J/(mol • ~ V~98in cc/mol)

V298 ref. 8298 A B C D ref.

Diopside 66.10 (6) 142.7 231.06 22.849 34.877 --3036.68 Enstatite 31.35 (8) 66.27 115.88 13.607 19.318 -- 685.69 Sanidine 109.05 (7) 232.90 276.57 42.945 42.120 - 1635.85 Phlogopite 149.91 (7) 334.6 472.46 58.559 64.421 -2781.82 Quartz 22.59 (8) 41.51 78.94 - 3.486 18.853 -469.43 Tremolite 272.92 (7) 548.9 800.23 245.13 97.521 -4742.27

(4, 5) (4, 5, 8) (7)

(12) (8) (4, 7)

Expansivity: V~ = V~98K + V~98K (a + bT + cT 2 +dT 3)/100 (T in ~ Compressibility: P o - 6 - 3 V298K=V298~ (1 -m • 10 P + n • 10 p2) (p in kbar)

a• -2 b• -3 c• -7 d• -1~ ref. m n ref.

Diopside - 5.05 3.10 0.I 1 2.44 (2, 3) 0.93 3.10 (3, 6) Enstatite -- 8.55 2.27 9.64 -- 2.82 (8) 0.84 -- 5.42 (8) Sanidine 2.59 -0.24 48.63 -20.41 (9) 2.12 14.50 (1) Phlogopite -10.04 3.49 5.51 -10.20 (11) 2.34 18.20 (1) Quartz - 31.08 7.42 -41.92 39.32 (8) 1.98 7.86 (8) Tremolite -44.18 4.45 -15.70 6.11 (10) 1.30 0.0 (1")

1 Birch (1966); * actinolite 2 Cameron et al. (1973); 3 Hazen and Finger (1980); 4 Krupka et al. (1985a); 5 Krupka et al. (1985b); 6 Levien and Prewitt (1981); 7 Robie et al. (1979); 8 Robinson et al. (1983); 9 Skinner (1966); 10 Sueno et al. (1973); 11 Takeda and Morosin (1975) fluorphlogopite; 12 estimated, phlogopite = fluorphlogopite - sellaite + brucite, Robie et al. (1979)

Table 9. Mineral activities and calculated ~ H20

Trem c~ Phlog ~ En c~ Di ~ Kfs ~H2 0 a C~ H 2 0 b

SP-79-9 0.303 0.771 0.830 SP-79-11 0.327 0.868 0.826 SP-79-12 0.764 0.852 SP-79-19 0.478 0.963 0.935 SP-79-20 0.432 0.961 0.938 81-SP-13 0.266 0.862 0.856 81-TL-2 0.347 0.823 0.890 JM-TRD 0.320 0.863 0.907 82-SP-5a 0.385 0.874 82-SP-51 0.422 0.837 8t-SP-I 0.026 0.090 0.22t 0.527 81-SP-40 0.005 0.046 0.199 0.460

0.18 0.17

<0.50 0.17 0.15 0.14 0.17 0.12

0.772 0.10 0.834 0.12 0.894 0.14 0.17 0.919 0.07 0.07

a Calculated using equilibrium (1). b Calculated using equilibrium (2)

H 2 0 content o f biotite

The majority of mica and amphibole formulae reported here are based solely on electron microprobe analysis which does not measure H 2 or Fe2+/Fe 3+. Direct measurement of H2 in biotites f rom semi-pelitic metasediments in the N W and NE Adirondacks shows substantial OH deficien- cies suggesting a large oxybiotite component (Bohlen et al. 1980; Edwards and Essene 1988). In order to evaluate this possibility two biotite separates from sample 82-SP-51 were analyzed for their H2 content. These direct analyses of H2 yielded 3.60 and 3.96 wt% H 2 0 in reasonable agreement with 3.55 wt% H 2 0 which was calculated f rom microprobe data alone. Consequently, H 2 0 contents of micas and am- phiboles in this study are estimated assuming OH + F + C1= 2, but whenever mineral separation is possible, this assumption should be checked by independent analysis.

Stabil i ty o f synthetic vs. natural tremolite

An additional uncertainty in the application o f equilibrium (1) arises from results showing that some synthetic tremo- lites react with forsterite to produce enstatite + diopside + H 2 0 at a temperature that is more than 85 ~ C below the same reaction for a natural tremolite. Such a shift corre- sponds to a 9 kJ change in enthalpy of this reaction and may be due to a high dislocation density in the synthetic tremolites (Skippen and McKinstry 1985). In the extreme case this correction to the experimental data would shift the ~ H 2 0 ' s calculated in this study from 0.17 to 0.05 for equilibrium (1).

Calculation of water activities using equilibrium (1) are based on experiments that used synthetic tremolite as start- ing materials. However, end-member tremolites are difficult (impossible?) to synthesize, as the amphibole is Mg-en-

356

7

03 Og < m o 5 --I v"

d tY

0 3 3 03 LU i'Y rt

[ i [ I [

_ @

tz - I T r ~ f s ~ ' - - - - - ~

Tr i Ph / / 0i

Kfs i Qtz

�9 i ' + ' ~

i

~ / 0 0

/ I ~.~",' c ~

7 0 0 8 0 0 9 0 0

TEMPERATURE, ~

!

? ~ A

t000

03 t ~ < tw

o 5 . .J

&

0 3 3 03 laJ r

7 -

I

Tr Kf$

A

700 b

' ' ~ I ' I I l a H 2 0 = l . 0 1

XH20 =0.35 / W /-1

o,, ./ 4 o ? 1

8 0 0 9 0 0 t 0 0 0 TEMPERATURE, ~

Fig. 3a, b. Pressure - temperature diagram showing equilibria (1), (2), and (11), and illustrating inconsistencies that arise from using a mixed data set (see text for details), a (top) shows the location of equilibrium (11): the solid curve is from this study (see text), the dashed curve is from Valley and Essene (1980), and the dash-dot curve is from Wones and Dodge (1977). b (bottom) the stabilities of (1), (2) and ( l i ) at eHzO= 1 are calculated by using the experi- mentally determined stabilities of (i) and (11) to locate invariant point A. This location of (2) differs from the experimental reversal at XH20=0.35 by 60 ~ C (dashed curve vs. reversal bracket in b). These seemingly large inconsistencies translate to only small differ- ences of 0.03 to 0.05 in the eHzO estimated from the analysis of natural assemblages (eHzO ~0.15, see text)

riched (with Ca/Mg < 2/5; Jenkins 1987). Consequently, ac- tivity models which assume Ca/Mg=2/5 are not entirely consistent with those synthetic amphiboles used to deter- mine the stability of tremolite. A decrease in temperatures of 20 to 30~ is required to correct equilibrium (1) for the non end-member characteristics of the synthetic amphi- bole (Jenkins 1987). This corresponds to a small increase

in calculated eH20 ' s (Table 9) of approximately 0.05 (i.e. from 0.1 to 0.15).

Estimation o f total uncertainty

Sources of uncertainties in the estimation of a HaO, beyond those discussed above, include probe analyses and activity models. Uncertainties arising from activity models (particu- larly when applied to impure micas and amphiboles) are the most difficult to quantify and may make the largest contribution to possible errors in a H 2 0 determinations. For example, in a study of tremolite-edenite solid solutions, Graham and Navrotsky (1986) found qualitative evidence of large deviations from ideality. I f no correction is made for reduced activities of the near end-member minerals in this study then Fig. 2 shows that the ~H20 would range from 0.25 to 0.30 in these rocks. These values represent

a n approximate upper limit on c~H20 for the Mg-rich sam- ples. This limit, combined with other uncertainties, suggests that a realistic estimation of peak metamorphic aHzO in these rocks is 0.15 _ 0.14.

C - O - H - F f lu id c o m p o s i t i o n s

The presence of graphite in those rocks that record ~H20 (e.g. equilibrium 2) allows calculation of all fluid fugacities in the C - O - H system at a given value of Pc. Four inde- pendent reactions can be written, relating the six most im- portant fluid components (French 1966; Valley et al. 1982; Lamb and Valley 1984, 1985):

C+O2 =CO2 (14)

C O Ac 0 . 5 0 2 = C O 2 ( 1 5 )

H2 + 0.5 02 = HzO (t 6)

C H 4 -/- 2 02 = CO2 + 2 H20 (17)

The equilibrium constants for each reaction can be calculat- ed from a knowledge of the appropriate free energies at a given temperature (Robie et al. 1979). This yields four equations, and six unknowns (when graphite is present at known P and T), making the system divariant. Univariance is obtained if a fluid is present:

P F = P H 2 0 + P C O z + P C H 4 + P C O + P H 2 + P 0 2 (18)

Given a value of e H 2 0 (Table 9) it is then possible to calculate the fugacity of the six fluid species. It is important to emphasize that Eq. (18) assumes that a free fluid phase existed at the peak of metamorphism, an assumption that may not be justified. In our calculations the fugacity coeffi- cients for H20, COz, and CH4, are calculated from the modified Redlich-Kwong equation of state of Kerrick and Jacobs (1981) and Jacobs and Kerrick (1981). For CO and H2, fugacity coefficients are from Ryzhenko and Volkov (1971).

Using values of ~H20 estimated for the graphite-bear- ing samples (82-SP-5a and 82-SP-51) Eqs. (14) through (18) were solved simultaneously. This calculation involves a qua- dratic with two solutions. Thus, two sets of values are de- rived indicating that the peak metamorphic fluid (if present) was either CO2- or CHg-rich (Table 10).

The fugacities of HF and Fz can also be estimated for these two samples by applying results of fluorine-hydroxyl exchange experiments (Munoz and Ludington 1974; Lud- ington and Munoz 1975). These experiments allow the ca1-

357

Table 10. C - O - H -- F fluid compositions. Log fugacities of the C - O - H - F fluid species with mole fractions of selected C - O - H species, calculated for the graphite-bearing assemblages (82-SP-5a and 82-SP-51) at XH20=0.06, Pv=P=7.5 kbar, and T=725 ~ C. Two self-consistent solutions result from this calculation (see text)

Log fugacity of the fluid

H20 CO2 CI-I4 CO H 2 02 HF F2

Solution 1 2.89 4.96 0.75 2.67 0.77 -15.91 -0.81 -31.47 Solution 2 2.89 0.73 4.98 0.56 2.88 -20.14 -0.81 -33.58

X Fluid

H20 CO2 CI-I4 CO H2

Solution 1 0.06 0.93 5.3 x 10-5 6.4 x 10-3 1.9 x 10 .4 Solution 2 0.06 5.3 x 10 -s 0.92 4.9 x 10 5 0.02

culation o f f H 2 0 / f H F at a given P and T. The value of f H F can then be calculated from a knowledge of f H 2 0 (Table 9). This value can be combined with the calculated values o f f H 2 and AG~(HF) (Stull and Prophet 1971) to yield values of f F2 via the reaction

H2 + Fz = 2 HF (19)

The estimated values of f H F and fF2 are given in Table 10 and are very low in spite of the fluorine-rich nature of the phlogopites (1.40-2.08 wt%). Thus, the high fluorine contents of these minerals is stabilized in part by l owfH20 and low Fe rather than unusually fluorine-rich conditions (see Valley et al. 1982).

The causes of low H 2 0 activities

Fluid infiltration

If a free fluid phase were present during the peak of meta- morphism, then low water activities (Table 9) would mean that the metamorphic fluid was enriched in another species (e.g. CO2). The presence of a pervasive CO2-rich fluid can be ruled out, however, because sample 82-SP-5a (c~H20 0.17 is located < 2 m from a sample containing wollaston- ite + calcite + quartz + diopside, which requires that c~CO2 <0.3 at 730 ~ C and 7.5 kbar (Greenwood 1967; Zie- genbein and Johannes 1974; Valley 1985; Lamb et al. 1987). The presence of a mineral assemblage that requires low eCO2 does not preclude CO2 infiltration as the cause of low water activities if either: (1) the infiltrating CO2 was channelized and selectively avoided the wollastonite-bear- ing rocks, or (2) the CO2 infiltrates in a pervasive manner, but the amount of infiltrating CO2 is insufficient to exhaust the buffering capacity of the wollastonite-bearing sample. Selective infiltration (case 1) implies that fluids in the lower crust do not move in a pervasive manner, but instead are channelized and may be restricted to particular lithologies or zones of weakness. If the CO2 infiltrates in a pervasive manner (case 2), then the presence of wollastonite allows limits to be placed on the amounts of CO2 involved. These limits depend upon when the wollastonite formed; either (1) during the regional granulitc-facies metamorphic event, or (2) prior to the regional event, perhaps during the intru- sion of the Oregon Dome anorthosite which outcrops ap- proximately 800 m to the east of sample 82-SP-5a.

If the regional granulite-facies event was responsible for the formation of wollastonite by the reaction of calcite and quartz, then CO2 will be evolved; yet the presence of wollas- tonite shows that the peak metamorphic fluid was H20-rich (assuming a n H 2 0 + C O 2 fluid). Thus, formation of wollas- tonite in the amounts present in this sample (mode: 8.1% wollastonite, 12.6% quartz, 21.2% calcite, 57.7% diopside, and 0.4% sphene), would require infiltration of H20 rather than CO2. Infiltration of H20 is inconsistent with estimates of ~H20 in sample 82-SP-5a (which is < 2 m from the wol- lastonite-bearing sample), and is unlikely during granulite facies metamorphism.

Wollastonite may have formed prior to the regional metamorphic event during a contact metamorphic episode. For example, Valley (1985) shows that the formation of wollastonite, monticellite, and Akermanite in marble near anorthosite indicates that the Mt. Marcy anorthosite in- truded to relatively shallow depths (< 10 km) prior to the peak of granulite-facies metamorphism. If the Oregon Dome anorthosite intruded under similar conditions, then wollastonite may also have formed in a low pressure envi- ronment. Stabilization of wollastonite prior to the regional event means that the granulite facies CO2/rock ratio may be calculated by assuming the all of the quartz in this sam- ple formed by the breakdown of wollastonite to calcite + quartz under the influence of the infiltrating CO2. This calculation yields a maximum CO2/rock of approximately 0.2 (volumetrically). A CO2/rock ratio may also be calculat- ed for sample 82-SP-5a (the sample located <2 m from the wollastonite occurrence) by assuming that the orthopy- roxene in this sample formed as a result of infiltrating CO2 destabilizing phlogopite and forming orthopyroxene (e.g. via equilibrium (2)). Sample 82-SP-5a contains less than one mode percent orthopyroxene, and so the calculated CO2/rock ratio is relatively small (< 0.05 by volume). Thus, pervasive CO2-infiltration could have occurred only in small amounts and only if the wollastonite was formed prior to the regional granulite-facies event. However, an earlier low pressure contact metamorphic event could have also stabi- lized the orthopyroxene-bearing equilibria found in the samples investigated in this study. Thus, the low, peak metamorphic granulite-facies water activities reported here could be the result of anorthosite intrusion prior to the regional metamorphic event (i.e. metamorphism of already dry rock). If so, then it would be unnecessary to call on COz-infiltration to reduce water activities.

358

I l l ' I I i f l i ' I i I

t .0 I t4U', o.5 I o.3 IOA LI' (o.~11 , !

i2L l ! 1 I I Granite H20-CO z 1 l

~- I II t 0 0 .5 ~0 .5 \ 0 . t I ~oF m " �9 �9 ,nn ,

o l-/i

[ / \ I N , s \ \ \ \ \ \ \\ i -

F \ \ ' , \ \ \ .: -

4h . _ \ \ \ \ \ \ ." - L 1 . 0 \ V . \ \ :

21 \ a ~,. \ \ \ .. - r \ \ ,o .s o . s Z . , , \ \ \ , / - I- 1 . 0 ~ �9 " - �9 u.l �9 ~ \ - .. -

600 700 800 900 t000 TEMPERATURE, *C

Fig. 4. Melting relations for the fluid saturated system granite- H20--CO2 at variable XH20. The solid curve shows melting for pure water (XH20 = 1.0, see text), and the dotted line shows dry melting (Huang and Wyllie, 1975). Calibrations at reduced XH20 are from Newton (I 986), dashed curves; Kerrick (1972), dash dotted curve; Clemens (1984), solid circles; and Wyllie (1977), squares. The stippled box shows peak metamorphic P--T conditions. The dashed curves are projected from experiments of Bohlen et al. (1982) on the albite-H20 --CO2 system and indicate that the XH20 in an H20-CO2 fluid in equilibrium with a granitic melt would be 0.5 (at peak metamorphic P -T) . If this calibration is correct, then partial melting could not generate the low values of c~H20 (0.07 to 0.18) estimated in this study. However, the experimental data, and the other calculations all indicate lower XH20 at the P - T of interest, perhaps as low as 0.1, suggesting that partial melting during granulite facies metamorphism could cause the low water activities reported here

Melt ing

Melting reactions are a viable mechanism for producing low water activities, and various determinations of melt equilibria in the g r an i t e -Hz O-CO2 system are plotted on Fig. 4. Melting under water-rich conditions ( X H 2 0 = 1) has been determined by a variety o f workers both on granites (Boettcher and Wyllie 1968; Piwinskii 1968) and on A b + Or + Qtz + H 2 0 (Tuttle and Bowen 1958 ; Luth et al. 1964; Merrill et al. 1970). The results generally agree to within 20 ~ C and are shown by the solid curve on Fig. 4. Also shown are granite melting equilibria for various X H 2 0 in the H z O - C O 2 system. These equilibria have been deter- mined by a variety of methods (Fig. 4): (1) the dashed curves are from Newton (1986) and were derived by using melting experiments in the system Ab + HzO + COz (Bohlen et al. 1982) as a projection base to derive contours of X H z O ; (2) the dash-dot curve was calculated by Kerrick (1972); and (3) the individual points are based on melting experiments (circles and squares are from Clemens 1984, and Wyllie 1977, p. 63, respectively).

Clearly there is disagreement between the various deter- minations of melting in the g r a n i t e - H z O - C O 2 system at ~HgO < 1.0 (Fig. 4). Newton's (1986) contours indicate that an H z O - - C O 2 fluid in equilibrium with a granitic magma would have an X H : O of approximately 0.5 at P ~ 7.5 kbar and T ~ 7 3 0 ~ (c~HgO~0.6). However, both Kerrick's (1972) calculations and the experimental work indicate that

the X H 2 0 in may be as low as 0.1 under these same condi- tions. This low value of X H 2 0 is consistent with the water activities calculated in this study. Furthermore, the presence of F, C1, and B will depress the solidus of granitic melts (Burnham and Nekvasil 1986; Pichavant 1987). Conse- quently, at any given P and T the addition of F, C1, or B will reduce the ~ H 2 0 in the fluid phase that is in equilibri- um with a melt. It is very possible, therefore, that the water activities calculated in this study were generated by partial melting.

The widespread and common occurrences of migrna- tites, feldspathic sweats, and small localized dikes all sup- port the hypothesis that partial melting was important for dehydrating the Adirondacks both at the peak of metamor- phism and before (e.g. MeLelland and Husain 1986). No evidence has yet been found in the Adirondacks to indicate that this terrane was infiltrated by significant amounts o f a pervasive CO2-rich fluid during the peak of granulite- facies metamorphism (Valley and Essene 1980; Valley and O'Neil 1984; Lamb and Valley 1984, 1985; Valley 1985; Powers and Bohlen 1985), although COz infiltration may have been important on a very local scale.

Acknowledgements. We thank Jean Morrison for stable isotopic analyses, James McLelland for assistance in the field and for pro- viding sample JM-TRD, and Kevin Baker for analysis of Hz in biotite. This paper has benefitted from discussions with Eric Es- sene, and reviews by Steve Bohlen and John Ferry. Everett Glover provided valuable help with the microprobe, and Susan Smith and Helen Finney drated the figures. This study was supported by grants from the National Science Foundation (EAR 81-21214 and EAR 85-08102), the Gas Research Institute (5083-260-0852 and 5086-260-1425), and from the Geological Society of America (3017- 82).

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Received October 5, 1987 / Accepted July 11, 1988 Editorial responsibility: J. Ferry