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Journal of Petrology Volume 42 Number 9 Pages 1705-1728 2001
© Oxford University Press 2001
Formation of Wollastonite by Chemically Reactive Fluid Flow During Contact Metamorphism, Mt. Morrison Pendant, Sierra Nevada, California, USA
1DEPARTMENT OF EARTH AND PLANETARY SCIENCES, JOHNS HOPKINS UNIVERSITY, BALTIMORE, MD 21218, USA
2GEOPHYSICAL LABORATORY, 5251 BROAD BRANCH ROAD NW, WASHINGTON, DC 20015-1305, USA
Received June 18, 2000; Revised typescript accepted February 27, 2001
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ABSTRACT |
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TOP
ABSTRACT
INTRODUCTIONQuartz–calcite sandstones experienced the reaction calcite + quartz = wollastonite + CO2 during prograde contact metamorphism at P = 1500 bars and T = 560°C. Rocks were in equilibrium during reaction with a CO2–H2O fluid with XCO2 = 0·14. The transition from calcite-bearing, wollastonite-free to wollastonite-bearing, calcite-free rocks across the wollastonite isograd is only several millimeters wide. The wollastonite-forming reaction was driven by infiltration of quartz–calcite sandstone by chemically reactive H2O-rich fluids, and the distribution of wollastonite directly images the flow paths of reactive fluids during metamorphism. The mapped distribution of wollastonite and modeling of an O-isotope profile across a lithologic contact indicate that the principal direction of flow was layer-parallel, directed upward, with any cross-layer component of flow <0·1% of the layer-parallel component. Fluid flow was channeled at a scale of 1–100 m by pre-metamorphic dikes, thrust and strike-slip faults, fold hinges, bedding, and stratigraphic contacts. Limits on the amount of fluid, based on minimum and maximum estimates for the displacement of the wollastonite reaction front from the fluid source, are (0·7–1·9) x 105 cm3 fluid/cm2 rock. The sharpness of the wollastonite isograd, the consistency of mineral thermobarometry, the uniform measured 18O–16O fractionations between quartz and calcite, and model calculations all argue for a close approach to local mineral–fluid equilibrium during the wollastonite-forming reaction.
KEY WORDS: contact metamorphism, fluid flow, wollastonite, oxygen isotopes, reaction front
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INTRODUCTION |
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Flow of chemically reactive fluid may control the mineralogy, stable isotope composition, and trace element chemistry of rocks during contact metamorphism (e.g. Labotka et al., 1988
; Jamtveit et al., 1992a, 1992b; Nabelek & Labotka, 1993; Bowman et al., 1994; Roselle et al., 1999). A number of aspects of metamorphic fluid flow, however, remain the subject of debate, including the direction of flow (Labotka et al., 1988; Ferry & Dipple, 1992; Ferry, 1995a; Hanson, 1995a), whether numerical models that assume uniform rock properties adequately predict the geometry of fluid flow in aureoles that contain a variety of anisotropic structural features (Hanson, 1992, 1995b; Cook et al., 1997), and the degree to which local mineral–fluid equilibrium is attained (Lasaga & Rye, 1993; Bolton et al., 1999; Lasaga et al., 2000). These and related questions were addressed in an investigation of contact metamorphism in the Mt. Morrison pendant, eastern Sierra Nevada, California. The location is an exceptional site for the study of metamorphic fluid flow for several reasons. First, the Mt. Morrison Sandstone, the focus of this study, is ideal for application of continuum models of coupled fluid flow and reaction because it is unusually homogeneous in mineralogy and texture. Second, mineralogical differences among samples from the Mt. Morrison Sandstone within the study area are primarily the result of a single prograde metamorphic reaction:
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GEOLOGIC SETTING |
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The Mt. Morrison pendant is exposed over
200 km2 on the eastern edge of the Sierra Nevada batholith, and is composed of five fault-bounded structural blocks (Rinehart & Ross, 1964; Wise, 1996; Greene et al., 1997; Stevens & Greene, 1999). This study focused on the Mt. Morrison Sandstone in portions of the Convict Lake and McGee Mountain blocks (Fig. 1). The Devonian Mt. Morrison Sandstone is separated into lower and upper members that are both texturally uniform quartz–calcite sandstones by the Prow member, a marl of 50–70 m thickness. Individual sedimentary beds typically are 10–200 cm thick in the sandstone and 2–20 cm thick in the marl. Both sandstone and marl are converted to calc-silicate hornfels where reaction (1) occurred. Other formations of metasedimentary rock in the area are Cambrian to Devonian in stratigraphic age and are composed of combinations of argillite, limestone, quartzite and chert, and their metamorphic equivalents. For clarity, they are not differentiated in Fig. 1. Contact metamorphism in the area of Fig. 1 resulted from emplacement of the Round Valley Peak granodiorite. The age of metamorphism is constrained by both an 89 Ma U–Pb zircon age for the granodiorite (Stern et al., 1981) and an 89·6 ± 1·5 Ma Th–Pb age for metamorphic monazite in a sample of pelitic hornfels from location 6, Fig. 1 (Wing et al., 1999; Ferry, 2000).
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Although the area in Fig. 1 is structurally complex, the major structures fall into only three groups (Greene et al., 1997; Stevens & Greene, 1999). The oldest are the thrust faults. Next youngest are steeply plunging folds. Because of folding, bedding in the area is overturned and approximately homoclinal with dips 55–80°E except near fold hinges (Stevens & Greene, 1999, figs 3 and 7). Strike-slip faults are the youngest major structures. Thrusts, folds and strike-slip faults are all younger than Devonian, the stratigraphic age of the youngest deformed formation. They are older than an undeformed 225 Ma felsic dike that intruded the Laurel–Convict fault (Greene et al., 1997). Because the structures were in place at least 135 m.y. before Cretaceous contact metamorphism, they provide the opportunity to investigate the control of pre-metamorphic structures on the geometry of metamorphic fluid flow.
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METHODS OF INVESTIGATION |
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A total of 143 samples were collected, including 101 samples of sandstone and 32 samples of marl (or their metamorphic equivalents, hornfels) from the Mt. Morrison Formation, eight samples of pelitic hornfels from the Convict Lake Formation, and two samples of Round Valley Peak granodiorite. Sample locations and contacts within areas
10 000 m2 were mapped in three dimensions to an accuracy of several centimeters with a laser rangefinder equipped with a digital fluxgate compass. At larger scales, mapping was conducted either on 1:5400 color air photographs or 1:24 000 topographic maps.
Mineral assemblages were identified by optical and scanning electron microscopy of thin sections. Qualitative analyses of all minerals were obtained with the JEOL JXA-8600 electron microprobe at Johns Hopkins University. Minerals that showed detectable deviation from ideal compositions were quantitatively analyzed with wavelength-dispersive spectrometry using natural silicate and carbonate standards and a ZAF correction scheme (Armstrong, 1988). Modes of 32 samples of Mt. Morrison Sandstone were measured by counting 
2000 points in thin section with back-scattered electron imaging. Any uncertainty in the identification of a particular point was resolved by obtaining an energy dispersive spectrometry (EDS) X-ray spectrum. Modes of 16 representative samples of sandstone, marl, and hornfels are listed in Table 1. Compositions of analyzed minerals in a subset of these samples appear in Table 2.
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Calcite from 77 samples of Mt. Morrison Sandstone was analyzed for oxygen- and carbon-isotope composition following procedures described by Rumble et al. (1991)
. Approximately 2–3 mg finely powdered calcite was obtained with a 2 mm diamond-tipped drill from a polished rock slab. Calcite was dissolved in phosphoric acid (McCrea, 1950) in evacuated reaction vessels at 100°C. Evolved CO2 was analyzed with the Finnigan MAT 252 mass spectrometer at the Geophysical Laboratory. The acid fractionation factor was computed from the expression given by Swart et al. (1991) using their calibration for sealed reaction vessels. Results were normalized to the composition of calcite standard NBS-19 (
18O = 28·65
, VSMOW; 13C = 1·95, VPDB, Coplen, 1988, 1996). Analyses of NBS-19 and other standards indicate that analytical precision for both oxygen and carbon isotopes is approximately ±0·1. The 18O values of quartz from 29 samples of sandstone and hornfels, of quartz from two samples of Round Valley Peak granodiorite, and of wollastonite from three samples of hornfels were measured following procedures of Yui et al. (1995). Oxygen was extracted from 2 mg of mineral separate in an atmosphere of BrF5 using a CO2 laser fluorination system similar to that of Sharp (1990). The O2 gas was collected, purified, and directly analyzed with the Finnigan MAT 252 mass spectrometer at the Geophysical Laboratory. Duplicates of each sample typically were analyzed. Results were normalized to garnet standard UWG2 (18O = 5·8, Valley et al., 1995) whose composition was measured at the beginning and end of each analytical session. With five exceptions, duplicate analyses of quartz agreed within ±0·1; the other five pairs agreed within ±0·2. Because reported values for quartz were referenced to daily analyses of UWG2 standard, they have a precision of ±0·2; precision for wollastonite is slightly poorer, ±0·3.
All calculations of mineral–fluid equilibria used Berman’s (1988) thermodynamic database updated August, 1990. Fluids coexisting with sandstones and marls were considered CO2–H2O solutions that obey the Kerrick & Jacobs (1981) equation of state. Fluids coexisting with graphite-rich pelitic hornfelses were considered CO2–H2O–CH4 solutions with fH2O defined by the appropriate equation from table 1 of Connolly & Cesare (1993). Activities of components in mica, garnet, and cordierite were computed using ideal ionic mixing models; activities used for micas specifically are those listed by Holland & Powell (1990). The activity of the anorthite component in plagioclase was estimated from activity coefficients in table 3 of Carpenter & Ferry (1984). Molar volumes of minerals were taken from data of Berman (1988).
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TEXTURE, MINERALOGY, AND MINERAL REACTIONS |
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Rocks from the Mt. Morrison Sandstone can be divided into two groups on the basis of texture and mineralogy (Fig. 2; Table 1). Rocks that are texturally hornfelses typically contain >90% wollastonite (Wo) and quartz (Qtz) but <1% calcite (Cal). [Mineral abbreviations follow Kretz (1983)
.] Rocks that texturally are sandstones or marls typically contain >90% Cal + Qtz and are devoid of Wo. Diopside (Di), K-feldspar (Kfs), titanite (Ttn), apatite (Ap), zircon, (Zrn), and combinations of accessory allanite, pyrrhotite, chalcopyrite, and sphalerite also occur in all rocks. All but one sample of hornfels contain trace grossular (Grs); the exception (sample 9A) contains plagioclase (Pl) instead. Sandstones and marls also contain graphite (Gr) and combinations of Pl, tremolite (Tr), and phlogopite (Phl). Grossular occurs in one sample of sandstone (30A5-S) within 1 cm of its contact with Wo hornfels in the same thin section.
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Analyzed Wo and Cal in the Mt. Morrison Sandstone are nearly pure CaSiO3 and CaCO3, respectively, with <0·01 Mg + Fe + Mn per formula unit. All but one other mineral display limited chemical variability. K-feldspar is a (K,Na)AlSi3O8 solution with K/(K + Na) = 0·92–0·97. Diopside is close to a Ca(Mg,Fe)Si2O6 solid solution with Fe/(Fe + Mg) = 0·01–0·10 (Table 2). Analyzed Grs has Ca/(Ca + Fe2+ + Mg + Mn) = 0·96–0·99 and Al/(Al + Fe3+) = 0·95–1·00 (Table 2). Tremolite and Phl are close to Ca2(Fe,Mg)5Si8O22(OH)2 and K(Mg,Fe)3AlSi3O10(OH)2 solutions with Fe/(Fe + Mg) = 0·01–0·06 and 0·01–0·04, respectively (Table 2). The one analyzed mineral that has a significant range in composition is Pl, which varies between An02 and An93 in Wo-free rocks. In the single sample with Pl + Wo (9A), Pl is An36 (Table 2).
Selected minerals in three samples of graphitic pelitic hornfels from locations 5–7 (Fig. 2) were analyzed for geothermobarometry. All three samples contain Kfs, muscovite (Mu), biotite (Bt), Gr, and Qtz with combinations of Pl, Zrn, and rutile. In addition, sample 5A contains cordierite (Crd), and samples 6A and 7G contain andalusite (And). Compositions of Mu, Bt, Kfs, and Crd in the three samples are listed in Table 3.
Reactions that produced Tr, Phl, Di, Kfs, and the anorthite component of plagioclase in both sandstones and marls of the Mt. Morrison Sandstone could not be defined because unmetamorphosed equivalents are not exposed in the study area. The difference in mineralogy between the Wo-bearing hornfelses and their sandstone and marl equivalents can primarily be explained by prograde reaction (1) (Table 1). Trace Grs in the hornfelses probably developed by
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; Ferry et al., 1998). Because the focus of this study is formation of Wo by prograde reaction (1), the retrograde reaction is not considered further. |
THE WOLLASTONITE ISOGRAD |
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A wollastonite isograd based on reaction (1) was mapped in three dimensions at a variety of scales in the Mt. Morrison Sandstone as the contact between Wo-bearing hornfels and Wo-free sandstone or marl. The isograd is easily identified in the field and on color air photographs because sandstone and marl appear gray (as a result of finely disseminated Gr) whereas the Gr-free hornfels appears white (Fig. 3). The correlation of the color difference with the presence or absence of Wo was verified by examination of hundreds of rock samples in hand specimen and thin section. The Wo isograd is surprisingly sharp, with the transition from either sandstone or marl to Wo-bearing hornfels developed over a distance of just several millimeters (Figs 2–4). The sharpness of the Wo isograd at locations 29 and 30 was confirmed by measurement of modes along a traverse across the isograd where it could be captured within a single thin section (29F, 30A5, Table 1). Results for location 30 (Fig. 4) are representative of both. The decrease in modal Wo towards the Wo isograd along the traverses at locations 29 and 30 indicates that the isograd stalled during metamorphism at a position in the sandstone where reaction capacity was at a local minimum.
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The 17 exposures of Mt. Morrison Sandstone in Fig. 1 can be divided into three groups. One group (C) contains only Qtz–Cal sandstone and marl with no Wo. A second group (W) contains only Wo-bearing hornfels. The third group (M) contains a mixture of Qtz–Cal sandstone and Wo hornfels. On the basis of the distribution of M- and W-type exposures, Wo appears preferentially developed at the scale of Fig. 1 close to either outcrops of the Round Valley Peak granodiorite or a region where the northern and southern outcrops of the granodiorite are probably connected at depth. A close spatial relationship between occurrences of Wo hornfels and exposures of the Round Valley Peak granodiorite is found in parts of the Mt. Morrison pendant outside the area of Fig. 1 as well (Lackey & Valley, 1999, 2000).
Because of the difficulty in accurately representing the position of the Wo isograd at the scale of Fig. 1, the location of the isograd is illustrated at expanded scales in Figs 5 and 6. The configuration of the isograd in map view indicates that the isograd surface has a number of different orientations in three dimensions. The typical configuration is observed far from faults, dikes, and stratigraphic contacts, such as along sample traverse 28 and in the vicinity of locations 17 and 30 (Fig. 5), where the Wo isograd approximately follows a topographic contour. The isograd therefore is a gently undulating, nearly horizontal surface in three dimensions with Wo hornfels always developed at elevations below the surface and unreacted Qtz–Cal sandstone above. The Wo isograd, however, cuts across topographic contours where it parallels pre-metamorphic faults and dikes (Fig. 5) and stratigraphic contacts (Fig. 6). At these locations the isograd surface parallels the fault, dike, or contact in three dimensions with Wo developed between the isograd and either the fault, dike, or contact. At a scale of 1–10 m, the Wo isograd can define fingers whose long dimensions parallel bedding (Fig. 6). Along the sides of the fingers, the isograd surface parallels bedding planes in three dimensions with Wo developed preferentially within certain beds or groups of beds. Finally, the Wo isograd can form closed loops (Figs 5 and 6) that contain Wo hornfels within them. Each closed loop represents the cross-section through a steeply dipping tube-shaped isograd surface in three dimensions.
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STABLE-ISOTOPE GEOCHEMISTRY |
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Oxygen- and carbon-isotope compositions of minerals
Measured O- and C-isotope compositions of Cal and Qtz in 73 samples of Wo-free sandstone or marl from the Mt. Morrison Sandstone are listed in Table 4. For sandstone samples collected more than several meters from sandstone–marl contacts, values of
18OQtz, 18OCal, and 13CCal are in the range 13·7–16·5, 12·3–15·7, and -5·0 to -0·2, respectively. There is no apparent correlation between isotopic composition and stratigraphic position (Fig. 7). Marl from the Prow member has significantly greater 18OCal (19·5–24·8) but similar 13CCal (-3·1 to -0·7) (Fig. 7). Exceptions to these ranges in 18O are observed in samples collected within several meters of sandstone–marl contacts. Sample 20C from a thin marl layer within the upper sandstone member has significantly lower 18OCal (16·3) than marls from the Prow member (Fig. 7). Sample 37A from a thin sandstone layer within the Prow member has significantly greater 18OCal (19·9) than sandstone from the lower and upper members (Table 4). Similar exceptions are observed in 26 samples of sandstone and marl collected along a traverse perpendicular to the contact between the Prow and upper members at location 22 (Fig. 8). Measured values of 18OCal change smoothly across the contact over a distance of 12 m from 13·4 in sandstone 7 m from the contact to 24·6 in marl 5 m from the contact. Values of 18OQtz for sandstone change sympathetically from 14·9 in sandstone 7 m from the contact to 19·6 <1 m from the contact. The 18OCal profile along traverse 22, the low value of 18OCal in sample 20C, and the high value of 18OCal in sample 37A are considered the result of 18O–16O exchange between sandstone and marl across their contact.
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Measured 18OQtz values for 14 samples of hornfels with sandstone protoliths are in the range 14·0–15·8 (Table 4). The 18O of Cal, Qtz, and Wo were measured in four samples of hornfels with marl protoliths. Values of 18OCal for two samples are 15·2 and 16·7, and 18OQtz for one sample is 14·9 (Table 4). Analyzed 18OWo is 11·5, 9·3, and 9·6 for samples 14C, 15A, and 23C, respectively.
Measured values of 18OQtz for granodiorite from locations 9 and 38 (Fig. 1) are nearly the same, 9·4 and 9·8. There is no overlap in 18OQtz between samples of granodiorite and samples of either sandstone or hornfels.
Oxygen-isotope fractionations
The approach to 18O–16O exchange equilibrium between minerals was evaluated from measured O-isotope fractionations. The largest dataset is for Qtz–Cal pairs (Fig. 9). Within error of measurement, the array of data has a slope of unity, consistent with attainment of equilibrium. The Qtz–Cal O-isotope fractionation, 
18OQtz-Cal, during development of the Wo isograd was estimated by averaging 18OQtz-Cal for the 10 analyzed Qtz–Cal pairs from outcrops that contain Wo hornfels. The result is 18OQtz-Cal = 1·24 (
= 0·2; range = 0·95–1·52). The average 18OQtz-Cal for all 20 analyzed Qtz–Cal pairs is not significantly different (1·26, = 0·2). A value of 18OQtz-Cal = 1·24 ± 0·2 is consistent with equilibrium at the elevated T of metamorphism rather than conditions of sedimentation or diagenesis (Sharp & Kirschner, 1994). Lackey & Valley (2000) reported 18OQtz-Cal = 1·85 ± 0·33 for Qtz–Cal pairs from another portion of the pendant, and also concluded that 18O–16O exchange equilibrium between Qtz and Cal was attained or nearly so. The lower value of 18OQtz-Cal measured in this study is probably explained by a bias towards sampling near the Wo isograd (the focus of the investigation) rather than obtaining samples evenly distributed over the lower-grade portions of the pendant as well.
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Oxygen-isotope exchange equilibrium between Qtz–Wo and Cal–Wo pairs is difficult to evaluate because of limited data: a single measured value of 18OQtz-Wo = 5·6 ± 0·5 (sample 15A), and one value of 18OCal-Wo = 5·5 ± 0·4 (sample 23C). The difference, 0·1 ± 0·9, is barely within error of the range of measured 18OQtz-Cal, 0·95–1·59, and is consistent therefore with both Qtz–Wo and Cal–Wo 18O–16O exchange equilibrium. The relatively large uncertainty in the difference, however, allows for departures from Qtz–Wo and Cal–Wo equilibrium by up to 2. On the basis of a much larger dataset for Qtz–Wo pairs, Lackey & Valley (1999, 2000) concluded that Qtz and Wo approached but did not everywhere attain O-isotope equilibrium in the pendant.
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PRESSURE, TEMPERATURE, AND FLUID COMPOSITION |
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Pressure and temperature
The peak P–T conditions of contact metamorphism were estimated from mineral equilibria. The development of And in pelitic hornfelses at and near locations 5–7 (Fig. 1) limits peak conditions to the And stability field (Fig. 10). The occurrence of Grs, Qtz, and Wo without Pl in all but one sample of Wo hornfels constrains peak T to values less than the Grs–Qtz–An–Wo equilibrium. The greatest upper bound is provided by sample 1A, which contains Grs with the lowest measured activity of Ca3Al2Si3O12 (0·86). The occurrence of Qtz, Wo, and Pl without Grs in sample 9A constrains peak T to values greater than the Grs–Qtz–An–Wo equilibrium computed with the reduced activity of CaAl2Si2O8 (0·72) appropriate for the sample. Conditions of contact metamorphism therefore lie within the quadrilateral defined by the And–sillimanite (Sil) equilibrium curve and the two curves for the Grs–Qtz–An–Wo equilibrium (Fig. 10). The maximum possible P is 2500 bars. An approximate lower bound on P is based on the 1200 m elevation difference between the highest exposure of the Round Valley Peak granodiorite and the study area around Convict Lake below. The granodiorite is coarse grained, equigranular, and contains abundant biotite and hornblende. It is unlikely therefore that the granodiorite magma crystallized at a P below 500 bars. Assuming a lithostatic P gradient of 270 bars/km, peak P in the study area was probably not less than
1000 bars. The range 1000–2500 bars for the Mt. Morrison pendant encompasses the tightly constrained peak P of 1500 bars for contact metamorphism in the Ritter Range pendant (Ferry et al., 1998). Because the Ritter Range pendant is located only 30 km to the NW and because metamorphism occurred there only 2 m.y. earlier, the preferred estimate for peak P in the Mt. Morrison pendant is also taken as 1500 bars with an uncertainty of +1000/-500 bars.
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At 1500 bars, the Grs–Qtz–An–Wo equilibrium in samples 1A and 9A constrains peak T to the range 525–575°C (Fig. 10). The range is further limited by mineral equilibria in graphite-rich pelitic hornfels samples 5A, 6A, and 7G. Fluids in graphitic rocks during metamorphism typically are CO2–H2O–CH4 solutions with bulk H/O = 2 (Connolly & Cesare, 1993). Peak P–T conditions for samples 6A and 7G therefore lie along the curve for the equilibrium among Gr, Mu, Qtz, Kfs, And, and CO2–H2O–CH4 fluid. At 1500 bars the equilibrium defines a narrow range of T = 555–560°C for the two samples (Fig. 10). Peak P–T conditions for sample 5A lie along the curve for the equilibrium among Gr, Mu, Phl, Qtz, Crd, Kfs, and CO2–H2O–CH4 fluid with bulk H/O = 2. At 1500 bars the equilibrium defines a T of 540°C (Fig. 10). Constraints on peak T provided by mineral equilibria in both the Wo and pelitic hornfelses are in good agreement. Because the focus of this study is the development of the Wo isograd, the preferred peak T is 560°C, T recorded by the pelite sample (7G) collected nearest to an occurrence of Wo hornfels. On the basis of calculated results in Fig. 10, the uncertainty is considered to be ±25°C. The preferred value of peak T is consistent with measured 18OQtz-Cal of 1·24 ± 0·2 at the Wo isograd, which records T of 565 + 80/-60°C (Sharp & Kirschner, 1994). Other calibrations of the T dependence of 18OQtz-Cal give lower T values of 280–530°C (Friedman & O’Neil, 1977; Chiba et al., 1989) that nevertheless confirm that the measured 18OQtz-Cal developed at elevated T.
Fluid composition
Metamorphic fluid at the peak of metamorphism at the Wo isograd probably did not contain significant dissolved salts both because of the absence of scapolite from the hornfelses and because apatites analyzed by EDS contain no detectable Cl. Fluid therefore was assumed to be a CO2–H2O solution. The composition of fluid at the peak of contact metamorphism at the Wo isograd is defined by equilibrium (1), XCO2 = 0·14 + 0·09/-0·07 (Fig. 11). The uncertainty in XCO2 is from the uncertainty in peak P and T.
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FLUID FLOW DURING PROGRADE CONTACT METAMORPHISM |
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Evidence for chemically reactive fluid flow
It is impossible to form more than
0·2 modal % Wo in low-porosity rocks (<1%) by reaction (1) in a closed system at P–T conditions where Cal + Qtz + Wo are in equilibrium with CO2–H2O fluid with XCO2 0·1 (Rice & Ferry, 1982). The observed amounts of Wo in hornfelses in the Mt. Morrison pendant could develop only in a closed system (in the petrologic sense) provided that porosity was impossibly large during metamorphism. Using equation (2) of Ferry (1994), for example, porosity would had to have been at least 67% during progress of reaction (1) (assuming modal Wo content the average of all analyzed hornfels samples, initial pore fluid with XCO2 = 0, equilibrium pore fluid with XCO2= 0·14, and reaction at 1500 bars and 560°C). The result is not substantively changed by consideration of plausible uncertainties in P, T, or modal Wo content. The only mechanism that can produce the measured 15–79 modal % Wo in hornfels at the inferred P–T conditions of metamorphism is for reaction (1) to have been driven by infiltration of the Qtz–Cal sandstone by a chemically reactive fluid with XCO2 < 0·14 (see Ferry, 1991). Using the same argument, infiltration has been invoked as the driving force for formation of Wo in numerous instances of metamorphism worldwide (e.g. Rumble et al., 1982; Labotka et al., 1988; Heinrich & Gottschalk, 1994; Cartwright & Buick, 1995; Ferry, 1996; Ferry et al., 1998).
In spite of the firm petrologic evidence for a role of reactive fluid flow in the formation of Wo by reaction (1), there is no complementary O-isotopic evidence for infiltration of the hornfelses. Oxygen-isotope evidence for infiltration of rocks by reactive fluids during metamorphism is a difference in 18O between metamorphic rocks (or any of their constituent minerals) and their low-grade or unmetamorphosed equivalents that is larger than can be explained by mineral reaction and the effects of Rayleigh distillation (Rumble, 1982; Rumble et al., 1982; Nabelek et al., 1984; Roselle et al., 1999). The effect of Rayleigh distillation on 18OQtz during the development of Wo by reaction (1) in the hornfelses was computed using equation (4) of Valley (1986) and his representative value of 
CO2-rock = 1·006. Unreacted sandstone was assumed 75 modal % Qtz and 25% Cal with 18OCal = 14·0 before reaction, and hornfels was assumed to contain no Cal after reaction. Values for 18OQtz-Cal of 1·24 and 18OQtz-Wo of 5·6 were taken as those directly measured. The effect of reaction (1) is to increase 18OQtz by only 0·2–0·3. The 18OQtz in Wo hornfels therefore may be directly compared with that in unreacted Qtz–Cal sandstone to address the question of reactive fluid flow during metamorphism. There is complete overlap in measured 18OQtz values between hornfelses and their sandstone equivalents both at a scale of 0·1–10 m across the Wo isograd (Fig. 12) and for the area of Fig. 1 as a whole (Fig. 13). Thus, there is no O-isotopic evidence for infiltration of the Qtz–Cal sandstones in the study area by isotopically reactive fluids during formation of Wo. Any infiltrating fluids must have been in O-isotope exchange equilibrium with sandstone at the site of reaction (1). It is unlikely that the absence of any difference in 18OQtz between sandstone and equivalent hornfels is explained by a kinetic limitation to the exchange of O isotopes between Qtz and fluid. Calcite is a mineral that exchanges O isotopes rapidly with H2O-rich fluid at the conditions of contact metamorphism [see the discussion by Bowman et al. (1994)]. The close approach to isotope exchange equilibrium between Qtz and Cal (Fig. 9) therefore suggests that Qtz closely approached O-isotope exchange equilibrium with fluid as well. The apparent paradox posed by the presence of petrologic but absence of isotopic evidence for infiltration is explained below.
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Oxygen-isotopic evidence for infiltration of hornfelses with marl protoliths is equivocal. Measured 18OCal for hornfelses with marl protoliths and for unreacted marls collected more than several meters from sandstone–marl contacts do not overlap (Table 4). Values of 18OCal and 18OQtz for hornfelses with marl protoliths, however, overlap the range of corresponding values for hornfelses with sandstone protoliths and for the sandstone protoliths themselves. Although the difference in 18OCal between hornfelses and their marl protoliths could be explained by infiltration by an isotopically reactive fluid during metamorphism, the difference is more probably explained by 18O–16O exchange between sandstone and marl by diffusion (see Fig. 8).
Mechanism of fluid–rock reaction
There are two mechanisms for infiltration-driven metamorphic decarbonation reactions. First, reaction may be driven by fluid flowing along a P and/or T gradient with fluid and rock at local equilibrium everywhere in the flow system (Baumgartner & Ferry, 1991; Ferry, 1991). In this case, reactants and products occur together along a significant length of the flow path provided that time-integrated fluid flux is not excessively large. Second, reaction may be driven by input of fluid at the inlet of the flow system that is chemically out of equilibrium with rock (Ferry, 1991). In the second case, reactants and products coexist along the flow path only at a sharp interface (a reaction front), provided that reaction kinetics are not excessively sluggish. Upstream from the reaction front, the decarbonation reaction will have gone to completion; downstream from the reaction front, there is no reaction at all. There is a sharp interface, corresponding to the Wo isograd, between unreacted Qtz–Cal sandstones and Wo hornfelses that contain <2% Cal (Figs 2–4). The small amounts of Cal are considered to have developed by the reverse of reaction (1) during retrograde metamorphism. At the peak of metamorphism the Wo isograd is inferred to have separated completely unreacted from completely reacted rocks. The distribution of reactants and products of reaction (1) in the Mt. Morrison pendant therefore conforms to that predicted for reaction driven by infiltration of a disequilibrium fluid, but not to that predicted for reaction driven by flow at mineral–fluid equilibrium along P and/or T gradients.
Geometry of fluid flow
There are three constraints on the geometry of reactive fluid flow during prograde contact metamorphism. The first derives from a quantitative analysis of the profile in 18OCal across the sandstone–marl contact at location 22 (Fig. 8). The composition profile in Fig. 8 has the characteristics of an initial step function that has been broadened by diffusion and possibly displaced by advection but without any significant complication by kinetically limited mineral–fluid isotope exchange [see discussions by Bickle & Baker (1990), Bickle et al. (1997) and Baxter & DePaolo (2000)]. The mass continuity equation (Bickle & McKenzie, 1987) describes the evolution of 18O along the profile from the initial step function in space (z) and time (t). For one-dimensional diffusion and advection in a low-porosity medium, local mineral–fluid isotope exchange equilibrium, and considering 18O as the tracer component in fluid,
 |
is porosity; 
is tortuosity; is 18O of the fluid; 
is Darcy velocity of the fluid; Vr and Vf are moles O per unit volume rock and fluid, respectively; and is the isotope fractionation factor, [(18O/16O)rock/(18O/16O)fluid]. A solution to equation (3) was derived that combines the consideration of a possible displacement of the initial isotopic discontinuity by cross-layer advection [solution to equation (3) by Bickle & Baker (1990)] with the provision for different values of in sandstone and marl [solution of equation (3) by Ganor et al. (1989)]. Measured values of 18OCal in Fig. 8 were fitted to the solution by minimizing the 
2 statistic using a non-linear Levinson–Marquhardt method (Press et al., 1986) to obtain estimates of z*, D*Mt*, and D*St*, where subscripts M and S refer to marl and sandstone, respectively; t* is the duration of diffusion and advection; z* = [(t*Vf)/(Vr)];
 |
18O was displaced from the contact by advection; D* is a measure of the distance over which 18O was homogenized by diffusion in marl and sandstone. The analysis assumes that 18O and fluid flux were continuous across the contact and that Vr, Vf, Df, and were the same everywhere in sandstone and marl. The assumption of equal Vr in sandstone and marl is confirmed by modal data (Table 1). Average Vr for all analyzed sandstones and marls is 0·086 mol/cm3 and 0·085 mol/cm3, respectively. The far-field value for 18OCal for marl and sandstone was taken as that for sample P14 and the average for samples U10 and U11, respectively. Constant 18OQtz-Cal along the profile (Fig. 9) confirms the assumption of local mineral–fluid equilibrium.
Results for two cases are illustrated in Fig. 8. All reported errors correspond to 2. Case 1 refers to the common solution that assumes D* is the same for rock on both sides of the contact (Bickle & Baker, 1990; Bickle et al., 1997). The curve for case 1 (dashed) noticeably overpredicts 18OCal for marl 100–400 cm from the contact. The fit is appreciably improved, both visually and as indicated by a significant reduction in the 2 statistic, for the solution to equation (3) that permits different values of D* in sandstone and marl as well as advection of the initial isotopic discontinuity by cross-layer fluid flow (case 2). In case 2 the estimated advection distance, z*, is not significantly different from zero, a result that rules out any cross-layer component to metamorphic fluid flow at location 22 within error of measurement. Fluid flow must have been confined to a direction parallel to lithologic layering. Diffusion exclusively in the solid state over D* 1 m is ruled out because it would take an impossible time of 1011-1015 a [calculated using the compilation of diffusion coefficients given by Eiler et al. (1992)]. Diffusion must have occurred within some region of rapid transport such as grain boundaries.
Fluid flow direction during formation of Wo is further constrained by hornfels sample 9A collected 2 m from the contact between the Mt. Morrison Sandstone and the Round Valley Peak granodiorite (Fig. 1, Table 4). Measured 18OQtz (14·8) for the sample is within the range of 18OQtz values for unreacted sandstone but significantly different from 18OQtz for the granodiorite (Fig. 13). If there was an approach to O-isotopic exchange equilibrium between Qtz and fluid at the peak of metamorphism, the measured 18OQtz for sample 9A rules out any significant horizontal component to fluid flow out of the granodiorite into sandstone at location 9. Otherwise 18OQtz for sample 9A would approach that for the granodiorite, 9·4–9·8 (Ferry & Gerdes, 1998). The limit to the amount of horizontal fluid flow at location 9 is presented below.
The third constraint on the direction of fluid flow during formation of Wo is the orientation of the surface of the Wo isograd in three dimensions. Where the isograd is a reaction front, flow was perpendicular to the surface directed from hornfels to sandstone (Ferry, 1991; Ferry & Gerdes, 1998). Where the isograd is a reaction side (Yardley & Lloyd, 1995) flow was parallel to the plane of the surface. The direction of fluid flow at any location of the isograd therefore was either perpendicular to the surface or parallel to it. At locations where the Wo isograd surface is horizontal or nearly so, flow was either perpendicular, directed upward, or horizontal. Where the Wo isograd defines steeply dipping tube-like features, flow could only have been parallel to the long dimension of the tubes, directed either upward or downward. Where the Wo isograd defines layer-parallel fingers of alternating Wo hornfels and Qtz–Cal sandstone (Fig. 6), flow could only have been within the plane of lithologic layering, directed either upward or downward. At locations where the Wo isograd parallels a pre-metamorphic dike or stratigraphic contact, the direction of fluid flow was either perpendicular to the dike or contact directed from dike or contact into sandstone or parallel to the dike or contact. Only one flow direction is consistent with each of the different constraints imposed by the geometry of the Wo isograd, and it is parallel to the steeply dipping lithologic layering and directed upward. Layer-parallel, upward-directed flow is consistent with the constraints implied both by the quantitative analysis of the 18O profile in Fig. 8 and by the consideration of 18OQtz for sample 9A. The preferred direction of peak metamorphic fluid flow implies that the Wo isograd represents a reaction front only where its surface is approximately horizontal. At all other locations the isograd surface is a reaction side.
Source of reactive fluid
If reactive metamorphic fluid flow was upward, the source of infiltrating H2O-rich fluid that drove reaction (1) must lie at depth. The two plausible sources of H2O fluid at depth are the crystallizing Round Valley Peak granodiorite and dehydrating pelitic rocks from surrounding stratigraphic units. A fluid source in the granodiorite is favored for two reasons. First, at the scale of Fig. 1 there is a spatial association between occurrences of Wo and exposures of the granodiorite, implying a genetic relationship between the two, as concluded independently by Lackey & Valley (1999, 2000). Second, a source of reactive fluid in other stratigraphic units appears unlikely because there is no evidence for a measurable component of cross-bed metamorphic fluid flow either at location 22 (Fig. 8) or anywhere along the Wo isograd. If the granodiorite was the source of fluid, the Mt. Morrison Sandstone is probably cross-cut by the pluton at depth, as is illustrated schematically in Fig. 14.
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Structural control of the geometry of fluid flow
The distribution of Wo hornfels directly images the flow paths of chemically reactive fluids during peak metamorphism in the study area (Figs 5 and 6) and therefore serves to evaluate the control of pre-metamorphic structures on the geometry of fluid flow at a scale of 1–100 m. There is enhanced development of Wo along pre-metamorphic dikes, along thrust and strike-slip faults, along stratigraphic contacts, along certain sandstone beds or groups of beds, and near a small flexure of the contact between the Mt. Morrison Sandstone and the overlying stratigraphic formation. Wollastonite hornfels can be particularly widespread in the vicinity of intersections of thrust faults with either a strike-slip fault or a dike (Fig. 5). The close spatial relationship between occurrences of Wo hornfels and locations of faults, dikes, contacts, the hinge area of folds, and certain sedimentary beds implies that there was enhanced flow of reactive fluid in the vicinity of all of these features. The various structures evidently were more permeable at the peak of metamorphism than surrounding rocks and focused the fluid flow. The preferential development of Wo on the up-dip (NE) side of each of the dikes in Fig. 5, in particular, suggests that the dikes were permeable channelways for reactive fluid flow. The strong structural control to peak metamorphic fluid flow, illustrated schematically in Fig. 14, appears to be almost universal in metamorphic terrains worldwide (reviews by Oliver, 1996
; Ferry & Gerdes, 1998).
With a few notable exceptions (e.g. Cartwright & Weaver, 1997) numerical models of contact metamorphic fluid flow typically lack a sufficiently detailed permeability structure to reproduce the kind of complicated flow pattern documented for the Mt. Morrison pendant (e.g. Hanson, 1992, 1995b; Hanson et al., 1993; Gerdes et al., 1995a). Although these models may successfully predict the general pattern of flow, such as upward flow near the margins of a pluton, they fail to predict real complexities in the flow pattern at a scale <1 km. Consequently, field studies that integrate petrology and isotope geochemistry serve as essential guides to the development of detailed hydrodynamic models of contact metamorphic fluid flow systems (see Gerdes et al., 1995b; Cook et al., 1997).
Amount of fluid flow
Constraints from the distribution of wollastonite
The absence of a measurable displacement of the 18OCal profile in Fig. 8 from the lithologic contact (z* = 0) rules out any significant cross-layer component to fluid flow at any time during contact metamorphism. The absence of evidence anywhere in the study area for complete back-reaction of Wo hornfels to Cal + Qtz by the reverse of reaction (1) rules out both a significant component of cross-layer fluid flow and a significant component of horizontal layer-parallel flow at the peak of metamorphism. Petrologic and isotopic evidence therefore are consistent with primarily a single direction of fluid flow during formation of Wo, and the amount of fluid involved may be computed with a one-dimensional model. The distribution of Wo indicates that the amount of fluid flow was variable during formation of Wo; lower and upper bounds on the amount therefore were computed.
For one-dimensional fluid flow, the distance (zWo) that the Wo reaction front traveled from the inlet to the flow system is related to molar time-integrated fluid flux at the reaction front, q, by
 |
max is the value of reaction progress when reaction (1) goes to completion, XCO2 is the composition of fluid at the reaction front, and the composition of fluid introduced at the inlet to the flow system is considered pure H2O (Ferry, 1996). The term in parentheses that is integrated in equation (5) accounts for the change in fluid composition at the reaction front as it passes through changing P–T conditions. Equation (5) assumes local mineral–fluid equilibrium in the flow system and negligible mass transport by diffusion or dispersion along the fluid flow path. A lower bound on q was computed from equation (5) using the minimum estimate of zWo of 360 m that corresponds to largest vertical exposure of Wo hornfels in the study area (between locations 9 and 11, Fig. 1). The value adopted for max was 0·0086 mol/cm3, based on the average modal Wo content of 15 hornfels samples with sandstone protoliths. Fluid composition at the reaction front was assumed to be defined by the Cal–Qtz–Wo–fluid equilibrium at inferred P–T conditions along the fluid flow path with pCO2 + pH2O = Ptotal. Conditions at the surface were considered to be those of peak metamorphism (1500 bars, 560°C); conditions at depth were taken as those along a P gradient of 270 bars/km and a T gradient of 100°C/km. The P gradient is appropriate for upward near-vertical flow of fluid at lithostatic P through rocks of normal crustal density. The T gradient corresponds to the peak T at the present level of exposure (560°C) divided by the inferred depth of metamorphism in the area (5·6 km). The relationship between q and z specified by equation (5) was then computed by fitting values of (1/XCO2) along the flow path to a second-order polynomial in z, integrating the result with respect to z, and evaluating the integral between zero and zWo = 360 m. The result is q = 1660 mol fluid/cm2 rock (Table 5). Considering the molar volumes of CO2 and H2O at the present level of exposure, the equivalent volumetric time-integrated flux is 0·71 x 105 cm3 fluid/cm2 rock.
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The greatest distance that the Wo reaction front could have traveled upward during prograde contact metamorphism is the difference in elevation between the surface and the point at depth where Cal, Qtz, and Wo would have been in equilibrium with pure CO2. Using preferred values for peak P and T during metamorphism at the present surface and the inferred P and T gradients along the flow path, that point lies 1575 m below the surface. An upper bound on q was computed using equation (5) with zWo = 1575 m and values for the other input parameters as specified above. The calculated maximum q is 3720 mol/cm2 on a mole basis (Table 5) and 1·9 x 105 cm3/cm2 on a volume basis.
If input values of peak P and T, max, and T gradient are varied within geologically plausible limits (see footnotes to Table 5), calculated lower and upper bounds on q are different from the preferred estimates by a factor of no more than 2–3 (Table 5). Calculated values of q in Table 5 formally assume vertical, upward-directed fluid flow. The inferred direction of upward layer-parallel fluid flow, however, deviated from vertical by 20°. To correct for the effect of fluid flow parallel to layering or other structural features with a 70° dip, q would be computed by dividing values in Table 5 by cos(20°), i.e. 0·94. Results would differ from those reported in Table 5 by only 6%.
Constraints from oxygen-isotope data
The upper bound on z* from the quantitative analysis of the 18OCal profile in Fig. 8 sets an upper bound on the cross-layer component to any metamorphic fluid flow. Assuming O-isotope exchange equilibrium between rock and fluid, the time-integrated fluid flux q needed to displace any O-isotope discontinuity a distance zOx is
 |
). The maximum possible value for z* is 17 cm (Fig. 8). Equation (6) was used to compute q for zOx = 17 cm assuming the fluid was H2O (Nf = 1) and taking Vr = 0·085 mol/cm3, based on measured modes and molar volumes of minerals. The value for Vr is relatively insensitive to modal mineralogy. The result is q = 1·4 mol/cm2, <0·1% of the component of near-vertical prograde flow (Table 5), and quantitatively confirms that any cross-layer component to metamorphic fluid flow was small compared with the layer-parallel component.
At the time of emplacement of the Round Valley Peak granodiorite, there was a discontinuity in 18OQtz between 9–10 in the pluton and 14–17 in sandstone of the Mt. Morrison Sandstone (Table 4, Fig. 13). If there was a significant horizontal component to fluid flow across the contact and into sandstone, 18OQtz in hornfels close to the contact should have been altered from 14–17 to 9–10 (provided Qtz–fluid O-isotope exchange equilibrium was approached at the peak of metamorphism). The preservation of 18OQtz of 14·8 in hornfels sample 9A, collected 2 m from the pluton–hornfels contact, thus places an upper bound on any horizontal component of fluid flow across the contact. Setting zOx <200 cm in equation (6), q <17 mol/cm2. Horizontal flow across the pluton–hornfels contact at location 9 was not more than 1% of the component of near-vertical prograde flow (Table 5).
The absence of any O-isotopic alteration of Qtz in Wo-bearing hornfels in the area of Fig. 5 is consistent with the estimated limits on q reported in Table 5. If O-isotope exchange equilibrium was approached during peak metamorphism, infiltration of sandstone by magmatic fluid would have produced both a Wo reaction front and O-isotope alteration front that moved upward from the pluton–sandstone contact into the sandstone (see Ferry & Gerdes, 1998). Equations (5) and (6) predict that the mineral reaction front traveled faster than the isotope alteration front. The bounds on the separation between the two fronts can be computed using equation (6) and the bounds in q in Table 5 [after making a correction for internal volatile production as a result of reaction (1), (zWo)(max), along the flow path]. The separation, zOx, predicted from the best estimate of the bounds on q is 195–1290 m (Table 5). The predicted separation of the two fronts is greater than the largest observed vertical separation between the mineral reaction front and hornfels with unaltered 18OQtz below (135 m between locations 1 and 30, Fig. 5). The analysis therefore predicts that O-isotopic alteration of Wo hornfels should exist, but that isotopically altered hornfels lies at least 60 m below the present level of exposure, as illustrated schematically in Fig. 14. The explanation holds if input parameters to calculations with equations (5) and (6) are varied within geologically plausible limits (Table 5). Alternatively, the absence of O-isotope alteration of Qtz in hornfels could result from lack of an approach to Qtz–fluid O-isotope exchange equilibrium during metamorphism, although this seems unlikely considering the evidence for Qtz–Cal O-isotope exchange equilibrium (Fig. 9).
Comparison with other studies
Hydrodynamic models of contact metamorphic fluid flow provide an independent check on time-integrated fluxes computed from chemical effects (e.g. Norton & Taylor, 1979; Hanson, 1992, 1995b; Hanson et al., 1993). The preferred range in time-integrated fluid flux for contact metamorphism of the Mt. Morrison Sandstone is 1660–3720 mol/cm2 (Table 5), which, considering the compositions of the prograde metamorphic fluids, corresponds to a range in time-integrated mass flux of (4–10) x 105 kg/m2. The published hydrodynamic model of contact metamorphism most relevant to the Mt. Morrison pendant is that of Hanson et al. (1993) for the Ritter Range pendant, 30 km to the NE. Their preferred flow model is one of vertical, upward-directed flow at the margin of the pluton with a time-integrated flux of 4·5 x 105 kg/m2. Both the amount and direction of flow agree well with the results of this study. The values of time-integrated flux reported in Table 5 are consistent with the hydrodynamics of fluid flow around cooling plutons. Values of q of the order of 1000 mol/cm2, like those in Table 5, are emerging from numerous studies of contact aureoles worldwide (e.g. Hanson et al., 1993; Ferry, 1995b, 1996; Cook et al., 1997; Ferry & Rumble, 1997; Ferry et al., 1998; Cook & Bowman, 2000).
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MINERAL–FLUID EQUILIBRIUM |
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The assumption of local mineral–fluid equilibrium used in this study to quantitatively estimate time-integrated flux during contact metamorphism has recently drawn criticism from Lasaga & Rye (1993)
and Lasaga et al. (2000). They have proposed that fluids have compositions specified not by mineral–fluid equilibrium but by a rapidly attained, kinetically controlled steady state during infiltration-driven metamorphic reactions. To quantitatively assess the validity of the assumption of local equilibrium, the difference between the equilibrium and steady-state fluid compositions during metamorphism was estimated by computing the steady-state XCO2 (XCO2ss) from equation (10) of Lasaga et al. (2000):
 |
xsys is the length of the representative elemental volume parallel to the flow direction, XCO20 is the composition of the input fluid at the reaction site (considered zero), and a and b are coefficients that define a linear approximation to the isobaric T–XCO2 curve for reaction (1) at the P–T conditions of the reaction site (1500 bars, 560°C). The reaction rate constant was taken as that computed by Lasaga et al. (2000) from experimental data for the Mu–Qtz–Kfs–And–H2O reaction at 600°C corrected to 560°C using an activation energy of 25 kcal/mol. Quartz was assumed the rate-limiting mineral in the reaction, and its surface area per unit volume was computed by modeling its occurrence as spheres of 0·4 mm diameter with a modal concentration of 65% (Table 1, Fig. 2). The T–XCO2 equilibrium curve for reaction (1) was linearized as the tangent to the curve in Fig. 11 at 560°C and 1500 bars. Fluid velocity of 2 m/a was estimated from a time-integrated flux of 105 cm3 fluid/cm2 rock (Table 5), a flow time of 105 a appropriate for the duration of contact metamorphism (Norton & Taylor, 1979; Dipple & Ferry, 1996), and the representative value for porosity listed by Lasaga et al. (2000, table 1). A value of xsys = 1 m was also taken from Lasaga et al. (2000, table 1).
Adopting these values for the input parameters to the calculation, the computed difference between the equilibrium and steady-state XCO2 is vanishingly small, 3 x 10-5. The result is robust, and is not significantly different if any of the input parameters is changed within geologically plausible bounds. The analysis of Lasaga et al. (2000) validates the assumption of local mineral–fluid equilibrium used in this study and confirms conclusions drawn from more qualitative evidence such as the sharpness of the reaction front (Huang & Bowman, 1993), direct laboratory simulations of reaction (1) driven by H2O infiltration (Zhang et al., 2000), the consistency of mineral thermobarometry in the area, and the uniform measured values of 18OQtz-Cal.
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GRAIN-SIZE CONTROL OF POROSITY DURING METAMORPHISM |
|---|
As measured by the
2 statistic, the best fit to the 18OCal data in Fig. 8 is one in which the value for D* differs in marl and sandstone with DM*/DS* 1·79. If pore fluid was the same in sandstone and marl, then the effective interconnected porosity () was 80% larger in marl than in sandstone [equation (4)]. The principal difference between sandstone and marl is average grain diameter (60 µm in sandstone; 20 µm in marl) rather than mineralogy (Table 1), implying that grain size exerts a significant control on porosity during metamorphism. The empirical correlation between smaller grain size and larger porosity, observed in this study, has been demonstrated experimentally (Wark & Watson, 2000) and results from an equalization of the curvature of pore walls across the lithologic contact as rocks approach textural equilibrium.
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ACKNOWLEDGEMENTS |
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Sorena Sorensen first brought the study area to our attention. Cal Stevens and Rich Schweikert educated us in the field about the geology of the region, and Victoria Avery, Elizabeth Catlos, Martha Gerdes, Sarah Penniston-Dorland, and Sorena Sorensen assisted with the fieldwork. John Valley generously supplied a sample of UWG2 garnet standard. Mike Bickle kindly performed some preliminary fits of the data in Fig. 8 to equation (3). Ian Buick, Simon Harley, and Peter Nabelek provided helpful reviews. Research supported by grant EAR-9805346 from the Petrology and Geochemistry Program, Division of Earth Sciences, National Science Foundation.
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FOOTNOTES |
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*Corresponding author. Telephone: 410-516-8121. Fax: 410-516-7933. e-mail: jferry{at}jhu.edu
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