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Journal of Petrology Volume 42 Number 7 Pages 1301-1320 2001
© Oxford University Press 2001
High fO2 During Sillimanite Zone Metamorphism of Part of the Barrovian Type Locality, Glen Clova, Scotland
DEPARTMENT OF GEOLOGY AND GEOPHYSICS, YALE UNIVERSITY, PO BOX 208109, NEW HAVEN, CT 06520-8109, USA
Received November 8, 1999; Revised typescript accepted December 5, 2000
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONThe redox state of sillimanite zone (650–700°C, 5–6 kbar) metasediments of the Barrovian type area, Scotland, was investigated using estimates of metamorphic oxygen fugacity (fO2), sulfur fugacity (fS2), and fluid chemistry based on new determinations of mineral and rock compositions from 33 samples. A total of 94% of the samples lack graphite, contain both ilmenite–hematite solid solutions (RHOMOX) and magnetite, and had metamorphic fO2 about 2 log10 units above the quartz–fayalite–magnetite (QFM) buffer. The regional variation in metamorphic fO2 for these rocks was minimal, about ±0·3 log10 units, reflecting either a protolith that was homogeneous with respect to redox state, or an initially variable protolith whose redox state was homogenized by metamorphic fluid–rock interaction. RHOMOX inclusions in garnet porphyroblasts that become richer in ilmenite from the interior to the edge of the host porphyroblast suggest that at least some syn-metamorphic reduction of rock occurred. Significant variations in bulk-rock oxidation ratio (OR) that are probably inherited from sedimentary protoliths are found from one layer to the next; OR ranges mostly between
20 and 50 [OR = molecular 2Fe2O3 x 100/(2Fe2O3 + FeO)]. These OR variations are uncorrelated with fO2 and do not indicate that large, order-of-magnitude gradients in fO2 and redox state existed or were preserved between layers during metamorphism. The other 6% of the samples contain ilmenite, lack magnetite, and had low fO2 0–1 order of magnitude below QFM in the stability field of graphite. They are characterized by combinations of the following: large fluid HF/H2O; metasomatic, tourmaline-bearing veins; absence or rarity of primary organic matter; and crosscutting late metamorphic shear zones rich in carbonaceous material. Such observations suggest that locally low fO2 conditions may have been related to the influx of reducing fluids from elsewhere in the area. KEY WORDS: metamorphism; redox; Barrovian; Scotland; oxygen fugacity
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INTRODUCTION |
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The chemical and mineralogical evolution of the crust during orogenesis is critically dependent upon the redox state of metamorphic fluids (James, 1955
; Chinner, 1960; Buddington & Lindsley, 1964; Miyashiro, 1964; French, 1966; Eugster & Skippen, 1967; Thompson, 1970; Harte, 1975; Ohmoto & Kerrick, 1977; Rumble, 1978; Lamb & Valley, 1985; Frost, 1991a, 1991b; Brenan et al., 1995; Harlov et al., 1997; Ague, 1998). The classical Barrovian type locality in Glen Clova, Angus, Scotland, is a superb example of redox controls on metamorphism (Chinner, 1960). Based on field relations, petrography, and the chemical systematics of minerals and bulk rocks, Chinner (1960) showed that large differences in bulk-rock oxidation ratio (OR) existed at local and regional scales between metasediments during amphibolite facies metamorphism in Glen Clova [OR = molecular 2Fe2O3 x 100/(2Fe2O3 + FeO)]. These differences were attributed to chemical contrasts inherited, at least in part, from the sedimentary protoliths, and were taken as evidence that significant differences in fO2 and fluid composition existed between layers during metamorphism (Chinner, 1960). Rock oxidation ratio variations in Glen Clova have been unequivocally established, but important questions remain regarding the metamorphic fluids. For example, what were the quantitative variations in the fugacities of O2 and sulfur species such as S2 in the region? Did large variations in fO2 and fluid composition exist between intercalated metasedimentary layers? Could metamorphic fluid–rock interaction between layers have overcome characteristic buffer capacities inherited from variable protoliths and act to homogenize redox states of fluids over the field area? In an effort to address these questions, we present a petrologic study of the Glen Clova rocks focused on quantification and interpretation of regional variations in the values of intensive variables during amphibolite facies Barrovian metamorphism. Our results suggest that regional mapping of metamorphic fO2 and fluid compositions is useful for assessing possible redox state histories for the Barrovian type area. Throughout the text, we use the term ‘redox state’ in a general way to refer to the fO2 and composition of metamorphic pore fluids.
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GEOLOGIC RELATIONS |
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The metapelitic and metapsammitic rocks of Glen Clova form part of the Dalradian Supergroup and were deformed and metamorphosed during the Grampian orogeny (Harte et al., 1984
; Fettes et al., 1986; Rogers et al., 1989; Dempster & Bluck, 1995); peak metamorphism occurred at 470 Ma (Oliver et al., 2000; Baxter et al., 2000). The timing of prograde porphyroblast growth relative to the four main deformational events (D1–D4) in the area can be summarized as follows (McLellan, 1989): (1) garnet: mostly syn-D2, some syn- to post-D3; (2) staurolite: syn- to post-D2; (3) kyanite: pre- to syn-D3; (4) sillimanite: mostly post-D3. Some rocks also contain staurolite and kyanite that formed after sillimanite zone metamorphism and, presumably, after D3 deformation (see Chinner, 1961; McLellan, 1985). The inferred bedding and the primary planar deformational fabric are subhorizontal (dips are <30°) and, thus, the rocks are part of the ‘Flat Belt’ of the Scottish Highlands (Harte et al., 1984).
The Glen Clova area contains much loose float; we made every effort to sample only from fresh, in situ outcrops. We investigated rocks in and around Chinner’s (1960) zone of ‘Hematite-bearing gneisses and associated semipelitic gneisses’ (Fig. 1), which our mapping indicates is at least 200–250 m thick. Chinner (1960) reported that at the margins of this zone the hematite-bearing rocks are intimately interbedded at the inch scale with hematite-free, graphite-bearing rocks, but we note that most of the graphitic rocks described in his paper are located 2–5·5 km away from the hematite-bearing zone. We were unable to confirm or refute the presence of such inch-scale interbedding during the course of our field and laboratory work.
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McLellan (1985) inferred the main metamorphic reactions for the area. ‘Peak’ sillimanite zone conditions were 650–700°C at 6 kbar (McLellan, 1985), were probably accompanied by some degree of anatexis (McLellan, 1989), and may have required heat input by fluid advection (McLellan, 1989). No change in grade is evident across the field area, and most rocks retain some pre-sillimanite zone kyanite (Chinner, 1960, 1961; McLellan, 1985).
The oxide mineral assemblages provide a firm basis for subdivision of rock types. Chinner (1960) described assemblages of ilmenite + magnetite, ilmenite + magnetite + hematite, and hematite + magnetite. In addition, we recognize a fourth variety containing ilmenite alone.
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PETROGRAPHIC RELATIONS |
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We present new compositional data for silicate and oxide minerals in 33 rocks from the Glen Clova area. Because Chinner (1960)
provided data for garnet, muscovite, biotite, and rhombohedral oxide for his sample 11, we also consider this sample. Symbols are defined in Table 1, mineral assemblages are listed in Table 2, mineral compositions are given in Tables 3–, and analytical and data reduction procedures are described in the Appendix.
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Excellent petrographic descriptions of silicate assemblages have been given by many workers (e.g. Harry, 1958; Chinner, 1960, 1961; Harte & Johnson, 1969; McLellan, 1985, 1989). Consequently, we focus on the oxides and sulfides.
Prograde phases
The rhombohedral oxide (RHOMOX) minerals range primarily between ilmenite (Ilm) and hematite (Hem) in composition. RHOMOX are found throughout the rock matrix and as inclusions in garnet, aluminosilicate mineral, staurolite, and tourmaline porphyroblasts. Most crystals are subhedral to euhedral and 10–900 µm long. Unaltered RHOMOX in magnetite-bearing rocks contain exsolution lamellae; in many samples, these lamellae underwent additional fine-scale exsolution themselves (Fig. 2). The multiple sets of exsolution features almost certainly developed during progressive cooling from ‘peak’ T, as expected from phase relations in the Fe2O3–FeTiO3 system (e.g. Burton, 1991), and suggest, although do not prove, that little chemical alteration of the grains occurred during cooling (Harlov et al., 1997). Rocks without magnetite contain Ilm (XIlm > 0·95) that lacks exsolution features.
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Cubic solid solutions vary primarily between the magnetite (Mag) and ulvöspinel (Usp) endmembers, but are dominated by Mag (XMag > 0·95). Mag crystals are found in the rock matrix and, more rarely, as inclusions within porphyroblasts. Most grains are subhedral to euhedral and 10 and 1100 µm in diameter. Rare martite alteration of probable supergene origin is present in a few samples.
Rutile (Rt) is uncommon. Crystals are generally subhedral, 100 µm long, and found in rock matrices and as inclusions within porphyroblasts. Rt crystals are typically rimmed or partially replaced by RHOMOX. It is unclear if these rim or replacement textures developed during prograde metamorphism or if they are retrograde. Additional Rt parageneses that are almost certainly retrograde are described below.
Some samples contain small amounts (<1 vol. %) of sulfides. Pyrite (Py), pyrrhotite (Po), and chalcopyrite (Ccp) crystals are found mainly in the rock matrix and much less commonly as inclusions within porphyroblasts. Crystals range from anhedral to euhedral and most are between 5 and 1000 µm in diameter. Ccp, Po, and Ni-sulfide are also found as tiny rounded ‘droplets’ within some Py. Evidence for some low-temperature supergene alteration of sulfides to Hem and limonite is present in most samples. Because supergene reactions may have modified the prograde sulfide assemblages (e.g. alteration of Po to Py), interpretation of metamorphic sulfide assemblages is somewhat hazardous.
Retrogression of Fe–Ti oxide phases
Small, Ilm-rich (XIlm > 0·85) RHOMOX ± Rt grains are commonly found (1) along cleavages and margins of biotite and, more rarely, muscovite crystals, and (2) in contact with and as partial rims around Mag grains. This Ilm ± Rt probably formed from retrograde Ti release from micas and Mag (e.g. Ague & Brimhall, 1988). The small Usp contents of the Mag grains probably reflect, at least in part, such retrograde Ti loss. It is unclear if retrograde Ilm nucleated and grew during cooling, or if pre-existing prograde metamorphic grains were compositionally modified during cooling.
Nearby matrix RHOMOX are largely pseudomorphed by fine-grained assemblages containing variable amounts of Rt, Ti-rich Ilm, Ti-poor Hem, and, rarely, titanite. These pseudomorphs are inferred to have originally been prograde crystals because the pseudomorphic replacement has generally preserved relict exsolution textures characteristic of cooling from high T. Fe2O3–FeTiO3 phase relations predict that at T below 500°C, Ti-poor Hem and Ti-rich Ilm should coexist (e.g. Burton, 1991). Thus, the pseudomorphed RHOMOX probably comprise retrograde minerals that may have equilibrated to some degree with nearby retrograde Ilm ± Rt associated with micas and Mag. RHOMOX grains that are far from micas and Mag or that are preserved as inclusions within garnet, aluminosilicate, or tourmaline porphyroblasts are typically not pseudomorphed.
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PRESSURE AND TEMPERATURE |
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P–T estimates for selected samples were computed with the TWEEQU program of Berman (1991)
using the program’s default thermodynamic data and activity models. Mole fractions were computed using version 0.95 of the CMP program available at the TWEEQU Web site (www.gis.nrcan.gc.ca/twq.html). The mole fraction of hydroxyl in micas was adjusted to include F and Cl substitution for OH (e.g. Ague & Brimhall, 1988; McMullin et al., 1991). We calculated simultaneous solutions of: (I) the garnet–biotite Fe–Mg exchange (GARB; Alm + Phl = Py + Ann) and the SGAM reactions of McMullin et al. (1991; Alm + Ms = Ann + 2 Sil + Qtz, Prp + Ms = Phl + 2 Sil + Qtz); or (II) GARB and GRAIL (Alm + 3 Rt = Sil + 2 Qtz + 3 Ilm). Method (II) was applied to sample 62A, which contains rutile that appears not to be retrogressed (see above). Reactions involving grossular in garnet were not used because grossular contents are generally small and subject to significant uncertainties stemming from estimated andradite contents.
Results cluster in the sillimanite field (Fig. 3), averaging 657°C and 5·7 kbar; we used 660°C and 5·7 kbar to calculate fluid composition herein. T was also estimated using GARB at 5·7 kbar for five samples that lack aluminosilicate minerals (36A, 45A, 49A, 50B, 52C). The average T of 668°C agrees with the simultaneous P–T estimates of Fig. 3. Because larger garnet grains preserve bell-shaped Mn growth zoning profiles, the temperature–time evolution of these rocks was insufficient to cause complete diffusional equilibration of cations in garnet. Our P–T results are fully consistent with those of McLellan (1985).
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, method (I); •, Method (II) (see text). Samples 46B, 54A, 55A, 57B, 61B, 62A, 66A, 70A, and 71A. Al2SiO5 phase relations from Berman (1988) |
REACTIONS |
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A fundamental conclusion of Chinner (1960)
was that, for any given sample, silicates and oxides reacted together as part of a thermodynamic system. The primary reaction that we used to estimate fO2 involves both silicates and magnetite
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Given an estimate of the fO2, fS2 was estimated using
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The fHF/fH2O estimates used measured biotite compositions (Table 5) following Zhu & Sverjensky (1992).
The log10(aMg2+/aFe2+) and log10(aMn2+/aFe2+) for the fluid were estimated with
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), but we emphasize that we are concerned with regional gradients in fluid composition that are independent of log K, not absolute values. Because thermodynamic data for Sps are lacking, we plot log10(aMn2+/aFe2+) -log10K for reaction (10). |
ACTIVITY–COMPOSITION RELATIONS AND THERMODYNAMIC DATA |
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We used the activity model of Ghiorso (1990)
for RHOMOX. However, Evans & Scaillet (1997) showed that current RHOMOX models overestimate fO2 for highly oxidized Mt Pinatubo dacites (see below).
Retrograde Ti loss commonly alters the bulk composition of Mag. However, because ‘peak’ metamorphic T was < 700°C, the expected XMag for spinel coexisting with the measured RHOMOX compositions would have been large and between about 0·80 and 1·0 (Andersen & Lindsley, 1988; Ghiorso & Sack, 1991). The Ti-poor Mag inclusions in garnet (Table 4) support this idea because they should have been shielded from retrograde equilibration with matrix phases and thus preserve prograde compositions. For T of 650–700°C in the Fe3O4–Fe2TiO4 system, the appropriate Mag activities for our samples are between about 0·8 and 1·0 (Andersen & Lindsley, 1988; Ghiorso & Sack, 1991); we used a very conservative range between 0·75 and 1·0 herein. Other components such as MgAl2O4 could also have affected Mag activity, but they are present in only trace amounts (Table 4); their concentrations are generally significant only at much higher T (Ghiorso & Sack, 1991).
Po that coexists with Py in the Fe–S system deviates from ideal FeS by as much as Fe0·875S at high T. However, natural high-T Po compositions commonly equilibrate down to very low T, at which many complex phase transitions occur. In view of these retrograde complications, we make the simplification of setting the activity of Po to 1·0. This simplification introduces some error, but the activity corrections for even highly nonstoichiometric Po have minor impact on the resolution of the large variations in fugacities of interest here. Py is relatively pure and its activity was set to 1·0.
Activity coefficients for garnet components, Ann, and Ms were computed following Berman (1990), McMullin et al. (1991) and Chatterjee & Froese (1975), respectively.
Because the rocks are in the sillimanite zone and nearly all samples contain Sil, thermodynamic data for sillimanite were used in the calculations; use of Ky has little impact on the results. Reconnaissance analyses indicate that the aluminosilicate minerals are relatively pure Al2SiO5 (XAl2SiO5 > 0·99) and thus can be modeled using unit activity. Qtz was assumed to be pure SiO2.
Thermodynamic data for Mag and most silicate components were taken from Berman (1988). Data for Hem, Ann, Alm, and F-bearing micas were taken from Ghiorso (1990), McMullin et al. (1991), Berman (1990), and Zhu & Sverjensky (1992), respectively, and are internally consistent with Berman (1988). Results computed with these standard state properties are nearly identical to those computed using recently revised properties for Alm and Mag (Berman & Aranovich, 1996). For example, the two datasets yield fO2 values for reaction (1), the main reaction we used to estimate fO2, that differ by only 0·04 log units. We used data from Johnson et al. (1992), as updated by Shock et al. (1997) and Sverjensky et al. (1997), for sulfide minerals, graphite, and most fluid species. Data for the gases CO, HF, and COS are not included in the Johnson et al. (1992) database and were thus taken from Robie et al. (1979), Stull & Prophet (1971), and Wagman et al. (1982), respectively.
Speciation calculations were carried out for C–O–H–S fluids composed of H2O, H2, CO2, CO, CH4, COS, S2, SO2, and H2S. Fugacity and activity coefficients for H2O and CO2 were from Kerrick & Jacobs (1981), whereas the other fugacities were computed using the equations of Shi & Saxena (1992) assuming ideal mixing of non-ideal gas species. Solution of the non-linear system of speciation equations was carried out using standard procedures (e.g. Ferry & Baumgartner, 1987; Holloway, 1987).
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FUGACITIES AND MINERAL CHEMISTRY |
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Oxygen fugacity
Results are given in terms of
QFM defined here as the difference, at 660°C and 5·7 kbar, between the log10 fO2 for a sample and that for the quartz–fayalite–magnetite (QFM) buffer (QFM log10fO2= -17·46). Because T–fO2 trajectories for rocks are generally subparallel to that for QFM, use of QFM notation largely removes uncertainties caused by errors in T estimates and, therefore, highlights real fO2 variations between samples. Furthermore, the reactions used to estimate fO2 are fairly insensitive to P variations, thus minimizing errors caused by errors in the P estimate.
QFM values computed using Hem-bearing reactions (QFM Hem) are plotted against the corresponding Mag-bearing reactions (QFM Mag) in Fig. 4. Most rocks have QFM Hem between +1·75 and +3·0 and QFM Mag of +1·8 to +2·3. However, QFM values for the Mag-free samples 46B and 62A are 2–3 orders of magnitude lower than those for the other rocks (QFM -0·25 to -0·75; Fig. 4).
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Closer inspection reveals that the QFM Hem estimates are as much as 10 times greater than the QFM Mag estimates. The six samples that do cluster near the 1:1 correlation line (37C-2, 39B, 52C, 61B, 66A, 71A) all have RHOMOX compositions that lie on the Ilm-rich side of the Hem–Ilm solvus. Agreement between QFM Hem and QFM Mag for these samples may be due to the fact that most experiments in the Fe2O3–FeTiO3 system have used Ilm-rich compositions (see Ghiorso, 1990). Consequently, the activity–composition relations for Ilm-rich RHOMOX are probably reasonably well understood. We note that QFM Hem for sample 39B, which plots somewhat farther from the 1:1 line than the other five samples, is based on the composition of matrix RHOMOX. The RHOMOX in this sample may have undergone some retrograde Ti gain by exchange with matrix phases during cooling (see above), and its QFM Hem may be somewhat underestimated.
The samples that plot well above the 1:1 line have compositions that are either on the solvus or to the Hem-rich side (Fig. 4). Problems with current RHOMOX activity models near the solvus have been documented (Evans & Scaillet, 1997), and the Hem-rich side of the solvus remains to be fully explored experimentally. Consequently, the QFM Hem values for these samples are inferred to be less reliable than those for Ilm-rich compositions. Furthermore, the RHOMOX in samples that plot well above the 1:1 line contain complex exsolution textures. It is possible that these grains originally straddled the solvus and were thus composed of two coexisting RHOMOX compositions at peak T. In this case, our ‘integrated’ compositions, which are averaged over all exsolution domains in RHOMOX grains, would be in error and could not be used in thermodynamic calculations. Resolution of the above issues requires a better understanding of phase relations among rhombohedral oxides.
To avoid problems with the thermodynamic treatment of RHOMOX, we focus on QFM values computed using Mag-bearing reactions (Table 7). The two exceptions are samples 46B and 62A, which lack Mag and require QFM estimates based on Hem-bearing reactions. Because these two samples contain Ilm-rich RHOMOX that lacks exsolution features, they should be treated reasonably well with the activity model and standard state thermodynamic data used herein.
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The agreement between QFM estimates made using reactions (1), (2), and (3) for Mag-bearing samples, or reactions (4), (5), and (6) for Mag-free samples (Table 7), indicates that the thermodynamic data and activity models employed are internally consistent.
Mineral chemistry
Prograde RHOMOX vary from XHem 0·02 to 0·75 (Table 3). XHem values for inferred retrograde RHOMOX intergrown with micas are smaller than 0·15 and are significantly smaller than the XHem of crystals included in nearby porphyroblasts (Table 3).
‘Felted mats’ of Sil intergrown with micas and, in some cases, retrograde St or Ky commonly contain significant quantities of RHOMOX. These yield XHem values that are equal to or smaller than those for inclusions in porphyroblasts (Table 3). It is unclear if the smaller values resulted from input of reducing fluid, retrogression, or both. Retrograde effects are difficult to rule out because the RHOMOX are closely intergrown with micas.
Samples 57B and 68A appear to have two coexisting prograde RHOMOX compositions that occur as discrete grains. Using the solvus diagram of Ghiorso (1990), compositions in 68A suggest a T near 680°C, in good agreement with the T estimates based on thermobarometry, whereas compositions in 57B are inconsistent with equilibration at a single T. The significance of these results is unclear given continuing uncertainties regarding T–activity–composition relations for RHOMOX.
Garnet Fe3+ is generally largest in cores and smallest in rims (Fig. 5). Fe3+ and Ti4+ decrease with Ca (Fig. 5), such that the ratio of grossular to andradite component remains fairly constant. Garnet Fe3+ contents are small and no correlation between QFM and garnet Fe3+ is evident, consistent with Chinner (1960). The Fe3+ systematics are apparently sensitive mostly to crystal chemical factors and the availability of Ca, rather than fO2.
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VI and
Garnet Mn content is in general positively correlated with rock Fe3+/Fe2+ (Chinner, 1960), although not necessarily with fO2. The XSps/XAlm ratio is small for QFM < 0, whereas a wide range of XSps/XAlm values seem compatible with QFM of +2 (Fig. 6). Sample 11 of Chinner (1960) is somewhat anomalous relative to our samples and has the largest XSps/XAlm and fO2 (Fig. 6). We note in this regard the possibility that the bulk garnet composition given by Chinner (1960) may be more Mn rich than the peak-T composition at garnet rims.
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The XMg/XFeT ratio for biotite changes little over the QFM range between -1·0 and +2·3 (Fig. 7; FeT is all Fe as FeO). However, XMg/XFeT increases at the largest QFM, consistent with the expectation that higher fO2 should favor more Mg-rich compositions (e.g. Wones & Eugster, 1965). The XMg/XFeT of muscovite (Ms) decreases between QFM -1·0 and +2·3, and then increases somewhat for higher QFM (Fig. 7). Fe3+ is the dominant form of Fe in Glen Clova Ms (Chinner, 1960). Under reducing conditions, the total Fe content of Ms is relatively small because the amount of Fe3+ available is limited relative to more oxidized conditions. Consequently, XMg/XFeT ratios for Ms in the low-fO2 rocks are large. The slight increase in XMg/XFeT above QFM +2 reflects the stability of magnesian compositions at higher fO2.
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Changes in RHOMOX inclusion compositions from core-to-rim in garnet
Some samples were examined in detail to determine if RHOMOX inclusion compositions changed from core to rim in garnets. This effort was hampered by the relatively small size and paucity of inclusions in many of the garnet crystals. None the less, RHOMOX included in the interiors of garnet porphyroblasts in samples 61B and 66A were measured to have significantly smaller XIlm than RHOMOX inclusions near garnet rims, or those included within Sil + Ms aggregates and Ky rims (Figs 2 and 8; Table 3). One interpretation of this shift towards larger XIlm is that reducing fluids, either locally or externally derived, caused reduction of the mineral assemblage during garnet growth. Unfortunately, uncertainties in activity–composition relations for RHOMOX preclude full quantification of possible fO2 changes recorded by the increases in XIlm, although calculations using the Ghiorso (1990) model suggest that drops in fO2 as large as an order of magnitude occurred during garnet growth. Because RHOMOX compositions between the small XIlm values in Grt interiors and the large XIlm values in rims were not observed, any such reduction was probably rapid compared with the rate of Grt growth.
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, inclusions in kyanite rims; 
, inclusions in sillimanite + muscovite aggregates. Generalized changes in RHOMOX compositions during garnet growth denoted by dashed arrows.
Relationships between fO2 and fS2
Most sulfide-bearing samples record fS2 at or slightly below the Py–Po buffer (Fig. 9). The average fS2 of Py-bearing rocks is slightly higher than that for Po-bearing ones, consistent with thermodynamic predictions. The two samples with the smallest QFM (46B, 62A) are the only ones that could have been in equilibrium with graphite. These record fS2 values 2–3 orders of magnitude less than the other rocks (Fig. 9; see below).
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REGIONAL VARIATIONS |
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Regional variations in fO2
Regional fO2 profiles were constructed for the northwest, central, and southeast parts of the field area (Fig. 1). Because the fabric of the rocks is nearly horizontal, sample data were projected onto vertical sections such that the profiles are essentially perpendicular to the fabric. A small number of samples (46–49) are located between the northwest and central profile areas and are thus not shown in Figs 10 and 11.
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QFM for the northwestern and southeastern profiles shows no large variations as a function of elevation (Fig. 10a and c). The main feature of the central profile is the strong localized minimum in QFM associated with sample 62A (Fig. 10b).
Regional variations in activities of Mg, Fe, and Mn species in solution
Surprisingly, the calculated log10(aMg2+/aFe2+) of the fluid varies little and shows no significant correlations with elevation or QFM (Fig. 11). Nearly all values are within ±10% of the mean. Other related ratios, such as log10(aMgCl+/aFeCl+), would show exactly the same patterns of regional variation [of course, the absolute values would be different from log10(aMg2+/aFe2+)]. These results strongly suggest that no large gradients in fluid Mg/Fe existed between layers. QFM for all samples must have been low enough so that Fe2+ species were dominant, otherwise the log10(aMg2+/aFe2+) would vary more strongly as a function of fluid Fe3+/Fe2+.
In contrast, log10(aMn2+/aFe2+) varies significantly and shows a weak positive correlation with fO2 (Fig. 11). In most cases the log10(aMn2+/aFe2+) variations are not abrupt and restricted to individual beds, but rather are gradational at the 10–100 m scale. Examples include the gradual decrease in log10(aMn2+/aFe2+) from 460 to 360 m in the northwest profile (Fig. 11a), and the ‘C’-shaped variation between 300 and 425 m in the southeast profile (Fig. 11c). The sharper discontinuity in the central profile corresponds to the extremely low fO2 sample 62A, and may be of metasomatic origin (see below). Metamorphic fluid–rock interaction was insufficient to homogenize fluid Mn2+/Fe2+ in the sequence, possibly owing to low Mn concentrations in fluids (i.e. large solid–fluid partition coefficient for Mn).
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DISCUSSION OF fO2 SYSTEMATICS |
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Chinner (1960)
used the rock oxidation ratio (OR) as an index of rock oxidation state, where OR = 2Fe2O3 x 100/(2Fe2O3 + FeO) (molecular basis). Although clearly powerful, OR is not a direct indicator of fO2 because rocks of differing OR can equilibrate at the same fO2. For example, for equilibrium at a given P and T, pore fluid in a rock that contains 99% Hem and 1% Mag would be at the same fO2 as fluid in a rock containing 99% Mag and 1% Hem, even though the OR values for the two rocks differ considerably.
In fact, a wide range of OR values, between about 20 and 50, are compatible with essentially the same QFM of about +2 (Fig. 12). Consequently, the observation of intimately interbedded rocks of differing OR (Chinner, 1960) does not necessarily indicate that the rocks had very different fO2 and fluid compositions. Extremes of the OR range are, however, associated with extremes of fO2. Sample 62A, equilibrated at QFM <0, has OR <20, whereas sample 11 of Chinner (1960) has an OR of 75 and QFM > +2·5. These extremes of the fO2 range are discussed in a later section.
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The homogeneity of the bulk of the metamorphic fO2 values can be interpreted in two ways. One possibility is that the lack of large variations reflects a protolith that was fairly homogeneous with respect to mineral and fluid compositions to start with (although OR values must have differed from bed to bed). Metamorphism of such a sequence would not be expected to produce large gradients in fO2 and fluid composition between layers, unless ‘exotic’, nonequilibrium fluids from outside the sequence were introduced.
The second possibility is that the sequence was originally more heterogeneous, such that sharp, layer-to-layer variations in redox state controlled by protolith composition existed. Metamorphic fluid–rock interaction between layers must then have acted to smooth these short wavelength variations so that the entire sequence attained a relatively uniform redox state. The observed increases in the XIlm and decreases in the XHem of RHOMOX inclusions from core to rim in garnet (Figs 2 and 8) suggest that some syn-metamorphic changes in redox state did occur.
If redox states were homogenized during metamorphism, the reaction histories would have been critically dependent on fluid compositions and the amount of fluid transport. The highly oxidized, QFM +2 fluids of Mag-bearing rocks would have been fairly inefficient oxidizers or reducers because such fluids consist mostly of H2O and CO2, and are poor in many species needed for redox reactions including H2, CH4, and CO (Wood & Walther, 1986; Frost, 1991a; Ague, 1998). For example, the reactions
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QFM +2 Glen Clova fluids, however, probably had XH2 of only 10-4 and vanishingly small XCH4 < 10-7 (Table 8). Consequently, large fluid–rock ratios would have been needed to homogenize the rock package. Quantification of these ratios is difficult owing to the uncertain activity–composition relations for RHOMOX, but calculations using simple model reactions such as (11) and (12), and representative fluid compositions (Table 8) and modes (Chinner, 1960) imply that if homogenization occurred, minimum fluid–rock ratios were 102 (Wood & Walther, 1986; Ague, 1998). Fluid exchange between beds would probably have been dominated by diffusion and mechanical dispersion across layers (e.g. Ague & Rye, 1999; Baxter & DePaolo, 2000), although advection may also have played a role. Modeling results suggest that disruption of protolith bed-scale fugacity heterogeneities may occur in as little as 102-103 years (Ague, 1998).
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PETROGENESIS OF SAMPLES AT THE EXTREMES OF THE fO2 RANGE |
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Rocks equilibrated at extremely low fO2
Two samples (46B, 62A), making up
6% of the sample set, are from layers that equilibrated at fO22–3 orders of magnitude lower than the rest of the rocks. The layers are at least meter scale in thickness, but whether contacts with surrounding rocks are sharp, gradational, tectonic, or otherwise could not be determined in the field. Either the sedimentary or diagenetic precursors to these two samples were highly reduced, or reduction resulted from metamorphic fluid–rock interaction. Some evidence from graphite and fluorine systematics suggests the latter possibility.
Graphite, inferred to have been derived from sedimentary organic matter, is common in low-OR rocks far removed (kilometer scale) from the study area (Chinner, 1960). In contrast, graphite is absent from the lowest fO2 sample (62A). This sample probably originally lacked or contained only traces of organic matter and, thus, its protolith may not have been highly ‘reduced’. Consequently, the low fO2 could have been produced during metamorphism.
Traces of graphite may be included within a few mica grains in 46B, but the identification is uncertain. On the other hand, the sample is cut by millimeter- to centimeter-wide, post-peak metamorphic shear zones containing abundant carbonaceous material, plagioclase, muscovite, and chlorite, and minor calcite (Fig. 13). Reducing fluids were therefore able to infiltrate the rock, albeit not necessarily on the prograde path.
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The low fO2 samples record larger HF/H2O than the others (Fig. 14). In addition, sample 46B is cut by veins that contain abundant F-rich tourmaline crystals as much as 1 cm long of probable metasomatic origin (see Dutrow et al., 1999). We suggest that the HF/H2O systematics reflect the local infiltration of chemically active fluids into the low fO2 rocks, and draw the connection that these fluids also drove syn-metamorphic reduction. The HF/H2O relations might also have been inherited from sedimentary protoliths, but it is difficult to envision what sedimentary process could produce a negative correlation between HF/H2O and QFM (Fig. 14).
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The above evidence suggests that infiltration of reduced metamorphic fluids drove reaction and lowered fO2 in samples 46B and 62A. If these fluids were at or near graphite saturation, they would have contained abundant H2 and CH4 (sample 62A, Table 8). Because these species are powerful reducing agents, alteration would have occurred even if fluid–rock ratios were rather small—of the order of one or less (Wood & Walther, 1986; Ague, 1998). The fluid could have been derived from graphitic rocks, such as those that surround the sequence studied here, or from reduced magmas that intruded during peak metamorphism. If the low fO2 was produced during metamorphism, the localized reduction suggests channelized fluid infiltration.
High fO2 sample
Sample 11 of Chinner (1960) has both large OR and QFM relative to our samples (Fig. 12). These characteristics were probably inherited from the protolith because (1) the large Mn and Fe contents of the rock suggest a protolith with high bulk-rock Fe2O3/FeO [see discussion by Chinner (1960)], and (2) no local or regional sources for highly oxidized metamorphic fluids needed to oxidize the rock are evident. One possibility is that sample 11 is an unusual relic that escaped to some degree redox equilibration with the rest of the sequence, perhaps because of extremely low porosity and/or permeability. Another possibility is that the bulk garnet composition of Chinner (1960) used in our calculations has lower Fe/Mn and almandine activity than the peak-T rim composition. If so, then the fO2 estimate made using the bulk garnet composition would be somewhat too high.
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CONCLUSION |
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Estimated fO2 for the bulk of the study area was about +2 log10 units above the quartz–fayalite–magnetite buffer (
QFM +2) at sillimanite zone conditions, although rare low fO2 rocks (QFM < 0) were also present. The high fO2 of QFM +2 is comparable with the highest known from arc settings (e.g. dacite from Mt Pinatubo; Evans & Scaillet, 1997). With the exception of the rare low fO2 rocks, variations in metamorphic fO2 across the study area were generally small, about ±0·3 log10 units. This minimal variability may reflect either (1) a protolith that was initially homogeneous with respect to redox state or (2) a protolith characterized by significant differences in redox state from bed to bed that was homogenized by metamorphic fluid–rock interaction. Rhombohedral oxide (RHOMOX) inclusions in garnet grains that become richer in ilmenite and poorer in hematite towards garnet rims suggest that some syn-metamorphic reduction of rock occurred. In detail, the QFM +2 rocks comprise intimately intercalated, graphite-free metasedimentary layers with whole-rock oxidation ratios (OR) that vary mostly between 20 and 50 from one layer to the next [OR = molecular 2Fe2O3 x 100/(2Fe2O3 + FeO)]. These OR variations are probably inherited from sedimentary protoliths, are uncorrelated with fO2, and do not indicate that large, order-of-magnitude gradients in fO2 and redox state existed between these layers during metamorphism. As pointed out by Chinner (1960), metamorphic fluid–rock interaction did not homogenize fluid compositions at scales of several kilometers, because the high fO2, QFM +2 rocks are in regional contact with Fe2O3-poor, graphite-bearing lithologies. The rare low fO2 rocks (QFM < 0) in our field area that occur among the dominant high fO2 (QFM +2) rocks may have been reduced by regionally migrating fluids derived from these graphite-bearing surroundings. |
APPENDIX: ANALYTICAL AND DATA REDUCTION PROCEDURES |
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Mineral composition determinations and electron microscopy used the JEOL JXA-8600 electron microprobe at Yale University. Quantitative analyses employed wavelength-dispersive spectrometers, 20 nA Faraday cup current, 15 kV accelerating voltage, natural and synthetic standards, off-peak and fluorescence-corrected, non-linear mean atomic number background corrections, and
(
z) matrix corrections. Electron microscope image processing and qualitative energy-dispersive chemical analyses were carried out with systems mounted on the JXA-8600; the early stages of our study used a Kevex Delta system and the later stages used an EDAX Phoenix system. All Fe was assigned to FeO for the matrix corrections. This procedure can lead to inaccurate results for Fe3+-rich phases if the ‘excess’ oxygen associated with Fe2O3 is not accounted for. Consequently, oxygen was incorporated by difference into the matrix correction iterations for RHOMOX and Mag such that the sum of all oxides and the ‘excess’ oxygen was 100 wt %. Similarly, H2O was specified by difference in the matrix corrections for micas such that the sum of the oxides and the H2O was 100 wt %.
Bulk compositions for the exsolved RHOMOX were estimated by making as many as 40 spot determinations on several grains per thin section. A 5 or 10 µm diameter beam was used in nearly all cases, but a focused beam was required for several fine-grained samples. Data were typically collected along linear traverses perpendicular to the long axes of the exsolution lamellae. At least five spot analyses on 1–2 grains per thin section were carried out for unexsolved RHOMOX (10 µm beam), Rt (5 µm beam), Mag (focused beam), garnet (focused beam), and micas (15 or 20 µm beam). Linear traverses across garnet involved as many as 50 determinations. Unless noted otherwise, reported garnet analyses are for inferred ‘peak’ compositions having the highest Mg and lowest Mn contents (usually several tens of µm from the rim).
Structural formulae for RHOMOX and magnetite were computed assuming 2 moles of cations per three oxygens for RHOMOX, and 3 moles of cations per four oxygens for magnetite. These procedures were judged successful because total oxide sums incorporating estimated FeO and Fe2O3 are always near 100 wt % (Tables 3 and 4). FeO and Fe2O3 in garnet can be estimated in several ways; we assumed 2 moles of octahedrally coordinated cations per 12 oxygens. This procedure gives the best results when tested against the wet chemical determinations by Deer et al. (1992), and gives total cation site occupancies very close to the ideal number of three for the four-fold and eight-fold coordinated sites (Table 6).
Bulk-rock X-ray fluorescence analyses were performed by X-ray Assay Laboratories, Don Mills, Ontario, Canada. FeO was determined using standard wet chemical titrations.
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ACKNOWLEDGEMENTS |
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We thank D. M. Rye, E. L. McLellan, R. L. Masters, and C. J. Carson for thoughtful discussions, M. S. Ghiorso for sharing both software and insightful advice for calculating activity–composition relations among Fe–Ti oxides, B. W. Evans, B. R. Frost, and an anonymous referee for critical reviews, and S. S. Sorensen for editorial handling. Financial support from the National Science Foundation, Division of Earth Sciences, grant EAR-9405889, is gratefully acknowledged.
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FOOTNOTES |
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*Corresponding author. Telephone: +1-203-432-3171. Fax: +1-203-432-3134. E-mail: jay.ague{at}yale.edu
Present address: Division of Earth and Planetary Sciences, California Institute of Technology, MC 170-25, 1200 E. California Blvd., Pasadena, CA 91125, USA.
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