Fault-fluid compositions from fluid-inclusion observations and solubilities of fracture-sealing minerals

ELSEVIER

Tectonophysics 290 (1998) 1–26

Fault-fluid compositions from fluid-inclusion observations and solubilities of fracture-sealing minerals W.T. Parry * Department of Geology and Geophysics, University of Utah, Salt Lake City, UT 84112-1183, USA Received 3 July 1996; accepted 26 November 1997

Abstract Host-rock chemical alteration and syntectonic veins in and near fault zones are evidence for episodic fracturing and fluid transport during faulting. Alteration minerals, vein fillings, and fluid inclusions may be used to estimate fault-fluid chemistry, temperature, and pressure. Fluid inclusions in thrust faults, reverse faults hosting mesothermal gold deposits, and exhumed footwall rocks of normal faults show that fluid components include NaCl, CO2 , CH4 and CaCl2 in addition to H2 O. Fluid composition, temperature, and pressure are spatially and temporally variable on most faults; a typical fault fluid does not exist. NaCl concentrations in fault fluids vary from 0 to 39 wt.%, CaCl2 concentrations range up to 19 wt.% and CO2 concentrations range up to 32 mole% in fluid inclusions, but some inclusions are present that are 100 mole% CO2 . Homogenization temperature measurements and pressure estimates confirm that these fluids were trapped at elevated pressure at depth on the faults. In CO2 -bearing fault fluids, pressures fluctuated, and a range of CO2 contents indicate effervescence. Varying solution densities of NaCl–H2 O fluids have been interpreted to result from entrapment of fluids in inclusions at constant temperature and varying pressures. Diverse fluid compositions are present on some faults with similar homogenization temperatures and estimated pressures suggesting similar depths on the faults. Pressure, temperature and fluid composition determine the solubilities of fracture-filling minerals calcite and quartz and the formation of alteration minerals that are related to the mechanical behavior of the rock. Quartz may precipitate as a result of cooling or pressure reduction, but calcite solubility increases with cooling and decreases with decreased PCO2 : Higher salinities increase solubilities of calcite and quartz and decrease the pH for equilibrium among feldspars, muscovite and solution. Mineral assemblages provide evidence of depressurization of the fluid as fluid moves from higher- to lower-pressured reservoirs. Precipitation of quartz, calcite, and K-feldspar or albite in fractures may result from fluid depressurization. Fault-zone rocks containing stilbite and laumontite reacted with fluid that contained little CO2 at comparatively low temperature and pressure; kaolinite, prehnite, muscovite, epidote, and chlorite formed from fluids at higher temperature and pressure. Variations in mineralogy and fluid-inclusion characteristics on individual faults suggest separate fluids that differ in chemical composition, temperature, and pressure.  1998 Elsevier Science B.V. All rights reserved. Keywords: fluid inclusions; faults; fluid pressure; fluid temperature; mineral solubility

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1. Introduction Faults have clearly acted as fluid conduits (Kerrich et al., 1984; Kerrich, 1986a,b; McCaig, 1988; Losh, 1989; McCaig et al., 1990), and chemical interaction of the fluids with altered fault-zone rocks is evident in fault zones (Sibson, 1981; Sibson et al., 1988; Parry et al., 1988, 1991). Fractures are filled with precipitated minerals, alteration product minerals are present in the host rock, and fluid inclusions are trapped in vein and matrix minerals. Heated aqueous fluids play a significant role in determining the mechanical stability of faults. Fault rupture is affected by healing and sealing of cracks with precipitated minerals, by alteration product minerals with lower frictional or flow strength (Kirby and Kronenberg, 1987; Janecke and Evans, 1988; Wintsch et al., 1995), and by elevated fluid pressures where rupture is governed by effective stress (Hubbert and Rubey, 1959). Fluid flow in a fault zone is controlled by the characteristics and proportions of fault core that consists of fault gouge and breccia, wall-rock damage zone of fractured country rock, and the protolith (Caine et al., 1996). The characteristics and proportions of each of these components varies both laterally and vertically (Byerlee, 1993; Bruhn et al., 1994; Caine et al., 1996). Fluids with a variety of sources and compositions interact with fault rocks at a range of temperatures and pressures. Fault zones initiated at high temperatures and pressures with locally derived fluids at low water=rock ratios evolve to higher fluid fluxes as the structures propagate. Near-surface fluids circulate on the faults later, at shallower depths and lower temperatures and pressures (Kerrich, 1986b). Cycles of fluid pressure fluctuations and vein formation often accompany faulting in the models of Sibson et al. (1988), Parry and Bruhn (1990), Cox et al. (1991), Boullier and Robert (1992), Cox (1995), and Robert et al. (1995). Byerlee (1993) has proposed a model for episodic fluid flow in faults. In a dilatant fault zone, core and damage zone are filled with fluid until the fluid pressure in the fault equals the hydrostatic pressure in the country rock. During compaction of the fault zone under ambient stress conditions, water saturated with silica and other minerals will seal fluid pathways in the fault rock, and fluids circulating up the fault zone will

precipitate additional minerals sealing and isolating a series of fluid compartments. The result, according to Byerlee (1993), will be formation of a series of seal-bounded fluid compartments with no hydraulic communication with each other. Further reduction of pore space in a sealed compartment by compaction results in increased fluid pressure. Rupturing the seal between compartments results in fluid flow from high-pressure to adjacent low-pressure compartments lowering the pressure in one compartment and raising the pressure in adjacent initially lowerpressure compartments. The minerals that seal the fractures precipitate because of a decrease in solubility that is the result of changes in temperature, pressure, and fluid composition. Crack healing is also affected by solution chemistry. Moderate increases in NaCl content and a change (increase or decrease) in pH increase healing rate (Brantley et al., 1990); dissolution or precipitation inhibitors (Fe, Al, Zn) and high mole fraction of nonaqueous phases such as CO2 could decrease healing rate. Macrofractures that are interconnected and held open transport most of the fluid volume and heal slowly, but microcracks of the order of 100 µm long and 10 µm wide allow pervasive penetration of fluid but heal quickly (Brantley and Voigt, 1989; Brantley et al., 1990). The fault models discussed above involve cycles of fluid pressure fluctuation accompanied by temperature-, pressure- and composition-dependent crack healing and sealing. Reviewed here are the best examples of measurements of the chemical composition, temperature and pressure of fault fluids in fluid inclusions associated with thrust faults, reverse faults hosting mesothermal gold–quartz veins, and normal faults. The solubilities of vein calcite and quartz and the stabilities of K-feldspar, muscovite, clay, and other minerals are then calculated for fluid compositions, temperatures, and pressures estimated from fluid-inclusion measurements. Mass and volumetric quantities of mineral precipitated as a result of changes in fluid temperature, pressure, and composition are then estimated. The objective is to provide a framework for utilizing characteristics of fault fluids determined from fluid-inclusion observations to estimate the efficiency of crack sealing due to changes in temperature, pressure and composition of fault fluids.

W.T. Parry / Tectonophysics 290 (1998) 1–26

2. Methods 2.1. Fluid inclusions Fluid present in faults exists in contact with fault rocks. Minerals precipitated from the fluid seal fractures and sometimes trap samples of the fluid as fluid inclusions. Measurements of the temperature of phase changes in fluid inclusions provide constraints on fluid temperature, pressure, and composition. Identity, density, and homogenization temperatures of fluid-inclusion contents were determined by observation of phase changes in doubly polished plates on a heating–freezing microscope stage using procedures outlined by Roedder (1984). Phase changes that were used to determine fluid composition were the temperature of the CO2 solid–liquid–vapor triple point, clathrate melting temperature, melting of ice, hydrohalite, and halite, CO2 liquid–vapor homogenization, and overall homogenization of fluid-inclusion contents. The temperature of overall homogenization of fluid-inclusion contents is a minimum temperature of entrapment because entrapment may take place along an isochore in the one-phase region. Salinities of moderate-salinity fluid inclusions are estimated from ice melting temperatures and the regression equation of Potter et al. (1978). Salinities of carbon dioxide bearing inclusions were estimated from clathrate melting temperatures (Bozzo et al., 1975; Collins, 1979). Salinities of inclusions with low eutectic temperatures (CaCl2 bearing) were estimated from phase melting temperatures and the equations and coefficients of Oakes et al. (1990), and salinities of inclusions that contained a halite crystal on the halite liquidus were estimated from halite dissolution temperatures and the equation and coefficients of Sterner et al. (1988). Densities of fluid-inclusion fluids at their homogenization temperatures were estimated using equations and data of Potter and Brown (1977) for moderate-salinity fluids and the MRK equation of state and data presented by Bowers and Helgeson (1983) for carbon dioxide containing fluids. Fluid pressures on faults have been estimated in a variety of ways (Roedder and Bodnar, 1980): (1) Solution density is estimated from measurement of homogenization temperature and pressure, then the pressure and temperature of the isochore

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are calculated using an appropriate equation of state, and the isochore is projected to its intersection with lithostat or hydrostat (pressure of a column of rock or water and temperature at depth). (2) Intersection of the solution isochore with the temperature determined by mineral equilibria, metamorphic conditions, or isotope fractionation. (3) The isochore intersection method utilizes fluid inclusions that contain CH4 , and=or CO2 in addition to water. The isochores of CH4 , CO2 and water have different slopes and when two or more of these phases are present the corresponding isochores can be calculated and extended to the pressure and temperature of their intersection. (4) The presence of substantial CO2 together with the CO2 –H2 O–NaCl phase relationships permits an estimate of minimum pressure and minimum temperature of entrapment. Entrapment of a homogeneous fluid must be demonstrated and post-entrapment changes must not have altered fluid densities. Pressure is estimated at the homogenization temperature on the two-phase boundary curve of CO2 – H2 O–NaCl phase diagram (Parry, 1986). The greatest source of error in fluid-inclusion measurements is a consequence of post-entrapment changes (Roedder, 1984). Dissolution and reprecipitation of host mineral surrounding the trapped fluid may result in a necked inclusion. If the necking takes place after separation of a vapor phase, then the result is varying liquid to vapor ratios and varying density of fluid-inclusion contents. Necking would not alter the salinity of the liquid phase. Measurements of fluid inclusions with liquid=vapor ratios affected by necking have been avoided by rejecting measurements on fluid inclusions in close proximity to one another with widely varying liquid to vapor ratios and by checking nearby inclusions with apparently similar liquid=vapor ratios to insure similar homogenization temperatures. Care is also taken to avoid measurements of fluid inclusions that have stretched, collapsed or leaked during deformation. For detailed discussion of the methods used in fluid-inclusion measurements and interpretations the reader is referred to Roedder (1984) and references therein, Parry (1986), and Parry et al. (1988, 1991). For detailed fluid-inclusion petrography, and for geology and displacement history of each fault the

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reader is referred to references for each specific fault system described below.

mHC C 2mCa2C C mNaC C mCaClC C mCaHCO3C

2.2. Mineral solubilities

mCa2C total D mCa2C C mCaCO3 ;aq C mCaHCO3

Calcite solubility has been reviewed, measured in experiments, and calculated as a function of temperature, CO2 pressure, and salinity by Ellis (1959, 1963), Malinen (1963), Holland and Borcsik (1965), Garrels and Christ (1965), Holland and Malinin (1979), Plummer and Busenberg (1982), Fournier (1985b), Wood (1986), Stumm and Morgan (1996), and others. Solubility of calcite is calculated here as a function of temperature, CO2 pressure and salinity. The activity of HC and calcite solubility were estimated from calcite saturation and charge balance at fixed PCO2 using the method described by Garrels and Christ (1965). We choose to calculate solubilities for externally fixed PCO2 rather than the closed-system solubilities reported by Wood (1986) because of fluid-inclusion observations that indicate variations in CO2 pressures in many fault fluids. Mass action expressions for the following chemical equilibria together with charge balance and an appropriate solution model form a series of equations which may be solved for activities and concentrations of aqueous species. CaCO3 ; calcite C CO2 ; g C H2 O D Ca2C C 2HCO3 (1) CO2 ; g D CO2 ; aq

(2)

CO2 ; aq C H2 O D HCO3 C HC

(3)

HCO3 D HC C CO23

(4)

C

CaCl D Ca

2C

C Cl

(5)

CaCl2 ; aq D Ca2C C 2Cl

(6)

2C C HCO3 CaHCOC 3 D Ca

(7)

H2 O D HC C OH

(8)

NaCl; aq D NaC C Cl

(9)

CaCO3 ; aq D Ca2C C CO3 2

(10)

D mOH C mCl C mHCO3 C 2mCO23

C mCaClC C mCaCl2 ;aq

(11)

(12)

Combining mass action statements for reactions 1 through 9 with charge balance Eq. 11 results in a fourth-order polynomial that may be solved for HC concentration. Calculated HC concentrations are then used in combination with activity coefficients and equilibrium constants to calculate the concentrations of aqueous species in reactions 1 to 10. Concentrations of aqueous calcium species permits the use of Eq. 12 to calculate the solubility of calcite. Equilibrium constants for the reactions 1 to 10 were calculated using the computer program SUPCRT92 (Johnson et al., 1992). Activity coefficients for aqueous species were calculated using the extended Debye–Huckel equation of Helgeson et al. (1981). Solubility of quartz in water and NaCl brine has been reviewed, measured, and calculated by Kennedy (1950), Morey (1962), Holland and Malinin (1979), Fournier (1983, 1985a), Wood (1986), Rimstidt (1997), and many others. Quartz solubility as a function of pressure and temperature was calculated using SUPCRT92 (Johnson et al., 1992). Solubility of quartz as a function of fluid salinity and temperature was calculated using the equations of Fournier (1983). 3. Composition, temperature, and pressure of fault fluids 3.1. Thrust faults The best examples of fluid characteristics in fluid inclusions related to thrust faults are: the studies of the Gavarnie thrust fault in the central Pyrenees (McCaig et al., 1990, 1995; Banks et al., 1991); the White Oak Mountains thrust in the southern Appalachians (Foreman and Dunne, 1991); a fault bend fold and the Yellow Springs thrust in the Appalachian fold and thrust belt (Srivastava and Engelder, 1990, 1991); the Rector Branch thrust in North Carolina (O’Hara and Haak, 1992); the Hope Valley shear zone in Connecticut and Rhode Island

W.T. Parry / Tectonophysics 290 (1998) 1–26

(O’Hara, 1991); the Window fault in the Apennines (Hodgkins and Stewart, 1994); the Himalaya (Boullier et al., 1991); and the Willard thrust fault in the fold and thrust belt of the western U.S. (Yonkee et al., 1989). Fluid composition, temperature, and pressure for these fault fluids are summarized in Table 1. Lithologies involved in these structures include unmetamorphosed, deformed Paleozoic and Mesozoic sedimentary rocks, greenschists, gneisses, and granites (Table 1). Constituents identified in the fault fluids are NaCl, CaCl2 , CO2 , CH4 , and H2 O. Concentrations of these constituents vary enormously. NaCl varies from 3 wt.% (Rector Branch thrust) to nearly 40% (Window fault), CaCl2 varies from undetectable amounts on most faults to 9.5 wt.% in the White Oak Mountains thrust, CO2 varies from undetectable amounts on most faults to 85 mole% in the structures in the Himalaya, and CH4 was detected on only one structure, the Yellow Springs thrust. No correlation of fluid composition with lithology is apparent. Evaporites are the source of saline brines on the Gavarnie thrust and the Window fault, and salinities vary upward on the Rector Branch thrust due to dilution by infiltration of near-surface water. Measured homogenization temperatures and estimated entrapment temperatures and pressures are also variable (Table 1). Fluid pressures are estimated using the extension of isochores from fluid-inclusion homogenization temperature measurements to the intersection with a metamorphic temperature (Gavarnie, Rector Branch) or the line defined by lithostat or hydrostat (White Oak Mountains, Window, fault bend fold). On the Willard thrust, fluid pressure is estimated from a combination of fluidinclusion measurements and thermal modeling to be 0.7 lithostatic at 325º to 350ºC (Yonkee et al., 1989). Pressure is often interpreted to have fluctuated, with lower homogenization temperatures (higher-density fluids) corresponding to entrapment at near-lithostatic pressures, and higher homogenization temperatures (lower-density fluids) trapped at lower pressures. Both pressure and salinity variations are interpreted by O’Hara and Haak (1992) to result from combined dilatancy and hydraulic fracturing. Depths of entrapment of lithostatically pressured fluids are estimated at 2.4 km (White Oak Mountains) to 10 km (Window fault). Fluid inclusions in highly strained quartz should

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be interpreted with caution because fluid density variations could be the result of inelastic volume change, decrepitation, diffusion, and leakage of the fluid inclusions during deformation (Boullier et al., 1991; O’Hara and Haak, 1992). A more detailed evaluation of fluid pressure is possible on the Hope Valley shear zone. Two types of fluid inclusions are present (O’Hara, 1991): (1) H2 O–NaCl–CO2 inclusions with 5 to 14 mole% CO2 and 7 wt.% NaCl homogenize at 262º to 350ºC and often decrepitate; and (2) aqueous inclusions with 3% NaCl that homogenize at 130º to 220ºC (mode at 160ºC). Isochores for the dense CO2 fluid inclusions indicate pressures of 750 to 900 MPa at peak metamorphic temperatures of 500º to 600ºC. Lowerdensity CO2 and aqueous fluid-inclusion isochores intersect at retrograde temperatures of 325º to 400ºC at pressures of 300 to 400 MPa. We have applied the methods outlined in Parry (1986) to the detailed fluid-inclusion measurements given by O’Hara (1991) to estimate minimum pressures of 95 to 280 MPa and 3 to 20 mole% CO2 . These pressure and temperature points are plotted on Fig. 1. The range in pressures and CO2 content suggests CO2 loss during episodes of lower pressure and that the fluids are trapped in inclusions near the two-phase boundary in the NaCl–H2 O–CO2 system as described by Spooner et al. (1987) and Parry et al. (1988). 3.2. Reverse faults hosting mesothermal gold–quartz veins Mesothermal quartz veins in greenschist facies country rocks occur in Archean greenstone belts of Canada, Brazil, western Australia, southern Africa, and in Phanerozoic vein deposits of New Zealand, eastern Australia, and the North American Cordillera from Alaska to the Sierra Nevada Mother Lode belt (Roberts, 1987; Vearncombe et al., 1989; Craw and Koons, 1989; Groves et al., 1989; Nesbitt and Muehlenbachs, 1989; Cox et al., 1991). Mesothermal gold–quartz veins display the role of fluids in faulting (Cox et al., 1986, 1991; Sibson et al., 1988; Boullier and Robert, 1992). The Superior province in Canada, one of the world’s largest Archean cratons, contains 120 large gold deposits of this type. Most of the largest deposits occur in the Abitibi subprovince and the single largest deposit is McIntyre-

Fault and lithology

Thrust faults Gavarnie: Hercynian phyllites on Cretaceous limestones

Reverse faults hosting mesothermal gold veins Sigma: metavolcanics Pamour: metasediment and metavolcanics Hollinger-McIntyre: Archean metasediments and metavolcanics Kerr-Addison: Archean metasediments and metavolcanics Canadian Cordillera: Proterozoic to Cretaceous metasediments and metavolcanics Contact Lake: Proterozoic granite Klamath Mtns, California: Paleozoic to Triassic metasediments and metavolcanics Mother Lode, California: metasediments and metavolcanics Normal faults Wasatch, Utah: Oligocene quartz monzonite Dixie Valley, Nevada: Oligocene granite to granodiorite Teton, Wyoming: Archean quartz monzonite White Wolf, California: Cretaceous granites Corral Canyon, Utah: Tertiary granites West Stansbury, Utah: Paleozoic carbonates and sandstones

Fluid composition wt.% NaCl, CaCl2 mole% CO2 , CH4

Temperature (ºC)

17 to 22 NaCl C CaCl2

130 to 210

6 to 7 NaCl 17.5 NaCl, 9.5 CaCl2 4.9 to 39.7 NaCl

130 to 240 70 š 22 250 to 320

104 to 182 62 to 94 105 to 240

Banks et al. (1991), Grant et al. (1990) McCaig et al. (1995) Foreman and Dunne (1991) Hodgkins and Stewart (1994)

20.5 to 23.4 NaCl

179 to 267

116 to 180

Srivastava and Engelder (1990)

CH4 <2% CO2

155 to 165

40 to 80

Srivastava and Engelder (1991)

84 CO2 3 to 26 NaCl 7 NaCl, 5 to 14 CO 2

120 to 320 262 to 350

260 at 300ºC 95 to 280

Boullier et al. (1991) O’Hara and Haak (1992) O’Hara (1991)

325 to 350

175

Yonkee et al. (1989)

<10 NaCl, 15 to 30 CO2 25 to 34 NaCl C CaCl2 , ¾100 CO2 2 to 9 NaCl, 2 to 14 CO2 0 to 20 NaCl, 3 to 24 CO2

285 to 395

200 to 300

Robert and Kelly (1987)

325 š 24 160 to 385

100 to 200 100 to 300

3 NaCl, 10 CO2 1 to 5 NaCl, 10 to 20 CO2

270 to 300 275 to 300

100 š 30

Walsh et al. (1988) Smith et al. (1984), Channer and Spooner (1994) Kishida and Kerrich (1987) Nesbitt and Muehlenbachs (1989)

2.5 to 6.3 NaCl 0 to 16 NaCl, 0 to 51 CO2

270 to 360 150 to 375

50 to 150 77 to 138

300 200 to 300

100 to 200 67 to 250

150 to 350

60 to 295

Parry et al. (1988)

340 263 to 347 137 to 272 155 to 335 100 to 230

37 to 157 80 to 160

Parry et al. (1991) Parry (1994)

4.7 NaCl HW, 17.4 FW

2 NaCl, 10 CO2 3.5 NaCl, 2 to 10 CO2 2 to 17.3 NaCl, 0 to 30 CO2 0.1 to 39.2 NaCl, 17 CaCl2 , 3 to 16.7 CO2 1.8 to 9.4 NaCl, 7 to 30 CO2 2.2 to 6 NaCl 0 to 27.8 NaCl 0 to 18 NaCl

Pressure (MPa)

50 to 182 at 300ºC

Reference

Fayek and Kyser (1995) Elder and Cashman (1992) Weir and Kerrick (1987) Coveney (1981)

Cady (1982), Parry (1994) Parry (1994)

W.T. Parry / Tectonophysics 290 (1998) 1–26

White Oak Mountains: thrust in Cambrian shale and limestone Window, Alpi Apuane, Italy: unmetamorphosed, brittlely deformed sediments on ductilely deformed greenschists Fault Bend, Pennsylvania: second-order fold in Ordovician carbonates Yellow Springs, Pennsylvania: brittle–ductile shear zones in Ordovician carbonates in H.W. Himalaya, Nepal: metamorphic rocks Rector Branch, N. Carolina: gneisses on greenschists Hope Valley, Connecticut: ductile shear in Precambrian and Devonian gneisses Willard, Utah: Proterozoic metasediments on Paleozoic carbonates

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Table 1 Summary of fault-fluid compositions, temperature, and pressure from fluid-inclusion observations

W.T. Parry / Tectonophysics 290 (1998) 1–26

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Fig. 1. Pressure and temperature of individual fluid inclusions in mesothermal gold–quartz veins (Hollinger, Smith et al., 1984; Sigma, Robert and Kelly, 1987; Klamath Mtns, Elder and Cashman, 1992; Hope Valley, O’Hara, 1991; Wasatch, Parry et al., 1988; Teton, Parry, 1994; Dixie Valley, Parry et al., 1991). Fluid pressures have been estimated from published fluid-inclusion measurements using the methods in Parry (1986). Lithostatic and hydrostatic pressure gradients are shown as fine lines (35ºC=km, 10, 26 MPa=km).

Hollinger (Card et al., 1989). These gold deposits consist of veins, disseminations, and replacements in high-angle reverse to reverse oblique, ductile to brittle–ductile shear zones (Sibson et al., 1988; Card et al., 1989) spatially associated with major fault zones tens to hundreds of kilometers long. The mineralized veins, emplaced in a dynamic tectonic environment during and after the formation of ductile shear zones (Robert and Brown, 1987), consist dominantly of quartz and carbonate. They are surrounded by zoned alteration envelopes a few meters to a few tens of meters thick characterized by carbonate, sericite, and sulfides (Card et al., 1989). Introduction of CO2 as carbonate minerals constitutes the major feature of visible alteration, and the abundance of carbonate diminishes systematically away from the mineralized veins (Robert and Brown, 1984; Kishida and Kerrich, 1987; Dube’ et al., 1987; Guha et al., 1991). The wall rock is depleted in SiO2 (Kerrich et al., 1977; Kerrich and Allison, 1978; Boyle, 1979; Coveney, 1981; Bohlke, 1989). Most

vein quartz is derived by bulk mass transport during shear zone deformation (Kerrich and Allison, 1978) under hydraulic pressure gradients (Kerrich and Allison, 1978). Silica is also redistributed by intercrystalline diffusive mass transport with transfer of material into extension veins (Kerrich et al., 1977). Hydrothermal minerals were deposited by fluids with access to the wall rocks through permeability generated by shearing (Colvine, 1989). Fluid-inclusion and mineralogical studies show that CO2 –H2 O–NaCl fluids of low to moderate salinity at temperatures from 250º to 450ºC and pressures from 200 to 400 MPa (Table 1; Fig. 1) were dominant in these faults (Smith et al., 1984; Kishida and Kerrich, 1987; Weir and Kerrick, 1987; Robert and Kelly, 1987; Card et al., 1989; Craw and Koons, 1989; Channer and Spooner, 1994). The development of these faults is consistent with the seismic cycle expressed in the fault-valve model of Sibson (1981, 1992) and Sibson et al. (1988). Unmixing of the CO2 –H2 O–NaCl fluid under decreased pressure

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conditions results in two dominant fluids: a CO2 -rich fluid and H2 O–NaCl fluid (Coveney, 1981; Spooner et al., 1985, 1987; Robert and Kelly, 1987; Walsh et al., 1988; Guha et al., 1991; Bowers, 1991; Boullier and Robert, 1992). Estimates of fluid pressure and composition from fluid-inclusion studies further indicate that fault-vein development takes place in two stages. Fluid pressure increases in the first stage result in opening of subhorizontal extension veins by hydraulic fracturing (Wood et al., 1986). In the second stage, slip along shear veins is accompanied by a drop in fluid pressure, collapse of the horizontal extension veins, and unmixing of the homogeneous CO2 –H2 O–NaCl fluid into CO2 -rich and an NaCl–H2 O-rich components (Wood et al., 1986; Boullier and Robert, 1992). Fluid pressure reduction implies increased porosity and=or escape of fluid into adjacent lower-pressure reservoirs. The high-angle reverse faults promoted cyclic fluctuation in fluid pressure from lithostatic to hydrostatic values as a result of cyclic episodes of rupture and vein filling. Dramatic reductions in fluid pressure caused quartz precipitation and phase separation of CO2 , an important mechanism for rapid precipitation of calcite (Sibson et al., 1988). Reduced gases such as H2 S, CH4 , and H2 in addition to CO2 are more volatile and are lost during phase separation resulting in oxidation of the fluid and precipitation of calcite with most of the calcite precipitation in initial stages of CO2 loss (Bowers, 1991). 3.2.1. Sigma and Hollinger-McIntyre Detailed fluid-inclusion measurements for the Sigma mine (Robert and Kelly, 1987) and the Hollinger-McIntyre mine (Smith et al., 1984) have been reinterpreted using the methods of Parry (1986). At the Sigma mine, pressures of 160 to 270 MPa at temperatures of 215º to 395ºC and 6 to 18 mole% CO2 are calculated and plotted on Fig. 1. Estimated pressures at Hollinger-McIntyre are 60 to 150 MPa at 243º to 308ºC and 8 to 11 mole% CO2 . These pressure–temperature points, plotted on Fig. 1 represent actual entrapment conditions because the fluids are trapped on the solvus (Spooner et al., 1987; Parry et al., 1988). Reinterpretation of the data of Smith et al. (1984) by Brown and Lamb (1986) suggests that for the higher temperatures inferred by Smith et al. (1984),

pressures could range from 400 to 800 MPa and reinterpretation of data for other deposits are consistent with trapping at pressures of 300–500 MPa for temperatures of 350º to 450ºC. 3.2.2. Klamath Mountains Lode gold mines in the central Klamath Mountains cluster along a major transcrustal thrust fault (Soap Creek Ridge fault) (Elder and Cashman, 1992). Veins occupy a reactivated, moderately southeast-dipping fault that originally formed synchronously with the transcrustal Soap Creek Ridge fault. Stage I veins represent multiple generations of fluid injection becoming younger inward with renewed fracture dilations; stage II veins are gold– quartz–carbonate veins; stage III veins are discontinuous lenticular bodies. Fluid inclusions contain 0–6% NaCl, 0–9 mole% CO2 and homogenize at 170º to 270ºC; high-salinity aqueous inclusions containing 11–16% NaCl homogenize above 285ºC; CO2 -rich inclusions with 13–51% CO2 , 1.4% NaCl and possibly CH4 also homogenize above 285ºC. Entrapment pressures of the CO2 -containing fluids, shown in Fig. 1, vary from 77 to 138 MPa. 3.3. Normal faults Footwall rocks of exhumed normal faults provide access for direct examination of rocks subjected to the action of heated fluids at elevated temperature and pressure at depth on the fault because of exposure of deeper rocks during progressive normal fault displacement. The history of chemical reactions in the footwall rocks is indicated by superposition of younger, lower temperature and pressure effects on the older, high temperature and pressure effects. We describe here the characteristics of fault fluids for the Wasatch, West Stansbury, and Corral Canyon faults, Utah; the Dixie Valley fault, Nevada, the Teton fault, Wyoming, and the White Wolf fault, California. Fluid characteristics are shown in Table 1 and Figs. 1 and 2. 3.3.1. Wasatch fault, Utah Fluid inclusions and alteration minerals have been examined in the southern portion of the Salt Lake segment of the Wasatch fault (Parry and Bruhn, 1986; Parry et al., 1988). Fault displacement in this

W.T. Parry / Tectonophysics 290 (1998) 1–26

area is at least 11 km since about 18 Ma (Parry and Bruhn, 1987). Vein filling and pervasive alteration minerals are biotite, K-feldspar, chlorite, epidote, muscovite, laumontite, prehnite, calcite, and clay minerals. Secondary fluid inclusions are present in healed microfractures in igneous quartz. Carbon dioxide bearing fluid inclusions associated with the chlorite and epidote alteration for which a full set of measurements have been obtained have salinities from 4.5 to 17.3 equivalent wt.% NaCl (Fig. 2) and contain 3 to 32 mole% CO2 shown in Fig. 1. A few fluid inclusions contain up to 100 mole% CO2 . Homogenization temperatures of these fluid inclusions range from 176º to 353ºC (Fig. 2). Minimum fluid pressures (Fig. 1) are estimated from the two-phase boundary curves for the system H2 O–CO2 –NaCl using the methods presented in Parry (1986), and vary from 60 to 295 MPa. Fluid temperature, pressure, and composition evolved along a path from 350º to 150ºC, from lithostatic to hydrostatic pressure, and from more than 30 mole% CO2 to less than 3 mole% CO2 with continued displacement of the fault. Moderate-salinity fluid inclusions are associated with laumontite, prehnite, and clay, and have salinities shown in Fig. 2 that range from 2 to 16 equivalent wt.% NaCl and no CO2 could be detected. 3.3.2. Dixie Valley fault, Nevada Fluid inclusions and alteration minerals have been examined in a portion of the footwall of the Dixie Valley fault, Nevada (Parry et al., 1991). Temporally and spatially overlapping hydrothermal alteration mineral assemblages occur as a narrow band near the 1954 rupture in the fault footwall. Early hydrothermal biotite and K-feldspar are followed by later Fe-chlorite and epidote then sericite and calcite. The latest hydrothermal minerals are stilbite, laumontite, kaolinite, alunite, smectite, illite and pervasive replacement of rock with fine-grained quartz, chalcedony, and opal. The displacement history of the fault is constrained by fluid temperature, pressure, and composition and alteration minerals. The age of the fault is 20 to 25 Ma with total displacement of 6 km (Parry et al., 1991). Secondary fluid inclusions are trapped in microfractures in igneous quartz. Moderate-salinity, CO2 -bearing, low-eutectic, and halite-bearing fluid inclusions are present. Moderate-salinity fluid-inclu-

9

sion salinity ranges from 0.1 to 16.9 equivalent wt.% NaCl (Fig. 2), CO2 -bearing fluid inclusions have salinities that range from 0.62 to 6.29 equivalent wt.% NaCl (Fig. 2). Ice melting temperatures in low-eutectic fluid inclusions are 10:1ºC to 26:0ºC and initial melting temperatures are as low as 45ºC and ice crystal nucleation temperatures are as low as 60ºC to 80ºC: These characteristics suggest the presence of a significant CaCl2 component. Low-eutectic inclusion salinities are 12.9 to 25.3 wt.% NaCl C CaCl2 with 30 to 55% of the total being NaCl. Halite bearing salinities are 30.1 to 39.2 equivalent wt.% NaCl. Carbon dioxide bearing fluid inclusions for which pressure estimates could be made contain 3 to 16.7 mole% CO2 and were trapped at minimum pressures up to 157 MPa at minimum temperature of up to 340ºC (Fig. 1). A few inclusions contain up to 51 mole% CO2 . 3.3.3. Teton fault, Wyoming The eastward-dipping Teton fault in west-central Wyoming is the easternmost normal fault in the northern Basin and Range province. The exposed footwall consists of Precambrian metamorphic and plutonic rocks. Displacement estimates range from 2 to 11 km, and age estimates range from 2 to 13 Ma, and Holocene fault scarps 3 to 52 m high indicate that the fault is active (Byrd et al., 1994). Precambrian quartz monzonite in the footwall of the Teton fault is hydrothermally altered. Muscovite occurs as an alteration product mineral as fine grains that are disseminated throughout the feldspars or as coarser grains in veinlets. Epidote and chlorite commonly replace biotite and hornblende. Prehnite rarely occurs as a replacement of biotite crystals. Fluid inclusions are present in quartz in the quartz monzonite and in quartz veins as planar arrays of secondary fluid inclusions. Carbon dioxide bearing fluid inclusions contain 1.8 to 9.2 equivalent wt.% NaCl and 7 to 30 mole% CO2 (Fig. 2). Homogenization temperatures are 263ºC to 347ºC, and estimated minimum entrapment pressures are 80 to 160 MPa (Fig. 1). Low-eutectic fluid-inclusions fluids contain 12 to 24 equivalent wt.% NaCl, but the dissolved salt is dominantly CaCl2 (20% NaCl, 80% CaCl2 ). 3.3.4. White Wolf fault, California The White Wolf fault separates the steep northwest slope of the Tehachapi Mountains from the

10

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W.T. Parry / Tectonophysics 290 (1998) 1–26

11

southern San Joaquin Valley southeast of Bakersfield, California. Exposed rocks on the hanging wall of the fault are predominantly the crystalline igneous rocks of the Sierra Nevada granitic batholith, and the footwall is composed of the Tertiary and Quaternary sediments of the San Joaquin valley. The historic uplift rate near the White Wolf fault has been 5 to 10 mm per year since 1926, and the post 0.6 to 1.2 m.y. net throw is 3.6 to 5.1 km (Stein and Thatcher, 1981; Stein et al., 1988). Uplift and erosion of the hanging wall of the White Wolf fault has exhumed granitic rock from depth that contains evidence of episodic fluid flow. Preliminary work on exhumed fault rock (Hitchcock, 1993) has identified extensive hydrothermal alteration, abundant quartz and calcite mineral veins, and secondary fluid inclusions that provide a record of fluid migration along the fault. Alteration minerals in exhumed granitic rock include K-feldspar, chlorite, sericite, laumontite, prehnite, and calcite. Secondary fluid inclusions were found in quartz in two samples of altered granitic rocks. These fluid inclusions have salinities that range from 2.2 to 6 equivalent wt.% NaCl. Homogenization temperatures are 137ºC to 272ºC shown in Fig. 2. Fluid inclusions showed no evidence of CO2 , but abundant calcite in the fault zone demonstrates the presence of CO2 in the fault fluids.

Fig. 3. Histograms of homogenization temperatures of fluid inclusions containing a tiny halite crystal from the Corral Canyon shear zone. Each fill pattern represents fluid inclusions from a separate healed fracture. Fluid-inclusion data from Cady (1982) and Parry (1994).

3.3.5. Corral Canyon shear zone, Utah The Corral Canyon shear zone is one of several high-level denudation faults resulting from extension in the eastern Great Basin that cuts Tertiary granite on the west flank of the Mineral Mountains in southwestern Utah. Alteration and studies of halitebearing fluid inclusions are reported by Cady (1982). In less deformed but highly sheared granite, sericite, chlorite, and epidote replace plagioclase and biotite. Greater abundance of sericite, chlorite and epidote with calcite are observed in more intensely deformed cataclasite. Fluid inclusions occur in vein quartz or in healed microfractures in primary igneous quartz.

The veins and microfractures are aligned with the major planes of failure within the shear zone. Salinities of NaCl–H2 O fluid inclusions range from 0 to 25.8 equivalent wt.% NaCl and homogenization temperatures are 150ºC to 335ºC (Fig. 2). Halite-bearing fluid inclusions contained a small halite crystal that was too small to observe melting behavior in the heating stage. Salinities of these inclusions calculated from an estimate of the size of the halite crystal were from 26.6 to 27.8% NaCl. The inclusions homogenized by vapor disappearance at temperatures of from 155ºC to 335ºC (Fig. 3). Several fluid in-

Fig. 2. Salinities and homogenization temperatures of fluid inclusions from the West Stansbury fault, Utah, the Corral Canyon shear zone, Utah, the Wasatch fault, Utah, the Dixie Valley fault, Nevada, the Teton fault, Wyoming, and the White Wolf fault, California. Fluid-inclusion data taken from Parry et al. (1988, 1991), Parry (1994). Type I inclusions are moderate-salinity NaCl–H2 O inclusions; type II are NaCl–H2 O–CO2 inclusions.

12

W.T. Parry / Tectonophysics 290 (1998) 1–26

clusions were measured by Cady (1982) on each of eight different healed fractures with the results shown in Fig. 3. Fluid inclusions on some fractures consistently homogenized at temperatures near 200ºC and others homogenized at temperatures as high as 300ºC. Homogenization temperatures of fluids with very low salinities also homogenized at temperatures from 150ºC to nearly 350ºC. We were not able to establish a fluid-inclusion chronology that would indicate the relative ages of low- and high-salinity fluids, but similar homogenization temperatures with different salinities suggests that highand low-salinity fluids were present at similar depths on the fault. 3.3.6. West Stansbury fault, Utah The West Stansbury fault forms the boundary between Skull Valley and the Stansbury Mountains in north-central Utah. Footwall rocks contain abundant quartz and calcite veins associated with faulting that are host for secondary fluid inclusions. Fluid salinities and homogenization temperatures are shown in Fig. 2. Fluid inclusions in calcite have relatively low salinities from 0 to 3 equivalent wt.% NaCl and homogenization temperatures from 120ºC to 340ºC. Fluid inclusions in vein quartz have higher salinities from 7 to 18 equivalent wt.% NaCl and homogenization temperatures from 100ºC to 230ºC. The homogenization temperatures for fluid inclusions in calcite and quartz overlap; the salinities do not, suggesting the presence of two types of fluids on the fault. 3.4. Summary of fluid-inclusion data Fluid inclusions and hydrothermal alteration minerals record the presence of aqueous fluids at depth in faults in a variety of host rocks. The major constituents identified in the fluids in fluid-inclusion studies are H2 O, NaCl, CaCl2 , CO2 , and CH4 . A survey of fluid-inclusion studies of fault-related veins (Table 1) shows that several types of fluid inclusions have been observed in fault systems: (1) moderate-salinity inclusions with salinities from 0 to 20 equivalent wt.% NaCl; (2) carbon dioxide bearing inclusions that contain up to 30 mole% CO2 and may contain CH4 in addition to a saline aqueous fluid;

(3) low eutectic temperature inclusions with initial melting temperatures as low as 45ºC suggesting a significant CaCl2 component in the fluid; and (4) inclusions containing a daughter halite crystal in addition to a vapor bubble. Fluid composition is extremely variable from near fresh water to saline brines that includes CaCl2 and from 0 to 100% dissolved gasses CO2 and CH4 . Fluid compositions vary at similar minimum temperatures and pressures on individual faults. NaCl concentrations vary from 0 to 39 wt.%, CaCl2 concentrations range up to 19 wt.% and CO2 concentrations range up to 32 mole% in fluid inclusions that also contain an aqueous salt solution, but some inclusions are present that are 100 mole% CO2 . Fluid composition, temperature, and pressure variability is as great within footwall rocks of an individual fault as among all of the faults studied. The source of dissolved constituents is not known, but the present-day Dixie Valley fault fluid involves evaporites from the nearby basin, and evaporites are the source of salinity in the northern Apennines, Italy (Hodgkins and Stewart, 1994) and in Alpine thrusting in the central Pyrenees (Banks et al., 1991; McCaig et al., 1995). The most common explanation of variable salinities in fluid inclusions is mixing of different fluid reservoirs with differing salinities, a process that is to be expected in tectonically active areas where repeated pulsing of fluid is likely to result in rapid compositional changes (O’Hara, 1995). Constituents in the fluid such as Ca, Na, and K may result from reactions with the host rock, and if the reactions are hydration reactions, salinity may increase (Crawford et al., 1979; Bennett and Barker, 1992; O’Hara, 1995). Variations in fluid density and CO2 content are accounted for by fluctuations in fluid pressure and by evolution of fluid pressure during slip history. Such fluctuations permit entrapment of fluid-inclusion fluids at high pressure (high density corresponding to low homogenization temperature) and during lower-pressure intervals (low density corresponding to higher homogenization temperature). Evidence from H2 O–CO2 –NaCl fluid inclusions in the rocks of thrust faults, normal faults, and high-angle reverse faults indicates that the fluid pressure at depth varies from lithostatic pressure exerted by a column of rock to a minimum pressure exerted by a column of fluid

W.T. Parry / Tectonophysics 290 (1998) 1–26

(Sibson et al., 1988; Parry and Bruhn, 1990; Robert et al., 1995; this study). Pressure fluctuations permit effervescence of CO2 and other volatiles so that a range in volatile content is observed in a series of fluid-inclusion observations. 4. Mineral solubilities in fault fluids The general model of episodic fracturing, fluid flow and fracture filling is consistent with observations of fluid-inclusion characteristics, alteration mineral abundance patterns, and veins present on thrust faults, reverse faults, and normal faults. Decompression of the fluid, coupled with loss of volatiles, results in precipitation of calcite and quartz sealing the fractures. Additional fracture-sealing minerals commonly observed are K-feldspar, laumontite, and muscovite (Fig. 4). The nature of the vein fillings also suggest elevated fluid pressures. For example, Fig. 4D shows a vein with no recognizable shear displacement of the vein walls filled with comminuted mineral fragments interpreted to have been emplaced by high-pressure fluids. Solubilities of vein minerals and stabilities of minerals that occur in pervasive alteration of fault wall-rocks are a function of fluid temperature, pressure and composition. The solubility of quartz increases with increasing temperature and increasing pressure so that both pressure reduction and cooling can result in quartz precipitation. The solubility of calcite in open systems with externally fixed CO2 pressure decreases drastically with increasing temperature and increases with increasing CO2 pressure. Fluid pressure not pressure-supported by solid phases is responsible for pressure contributions to solubilities (Bruton and Helgeson, 1983). We examine the effect of fluid decompression, temperature and salinity changes on the solubility of quartz and calcite in fault rocks and on equilibrium among alteration minerals and fault fluids. 4.1. Calcite Calculated calcite solubilities are comparable with solubilities measured in laboratory experiments of Ellis (1963) shown in Fig. 5. Calcite solubility decreases with increasing temperature and increases with increasing salinity as shown in Fig. 5. Fig. 5a

13

shows calcite solubilities from 100º to 300ºC at PCO2 of 0.12 MPa at salinities of 0.2 to 1 m NaCl. Salinities and temperatures of fault fluids vary over a wider range than the experiments of Ellis (1963) resulting in a wide range of solubilities of calcite. Representative NaCl–H2 O fault fluids are plotted in Fig. 5b together with isopleths of calcite solubility indicating calcite solubilities from 1 to more than 5 mmoles per kg of water at PCO2 of 0.1 MPa. Calcite solubility is plotted versus salinity in Fig. 5c showing solubility curves at constant temperature that are convex upward. The major effect of increasing salinity takes place between 0 and 1 m NaCl at each temperature. Increasing PCO2 increases calcite solubility as shown in Fig. 6; the effect is greater at lower temperatures and lower CO2 pressures. Cooling fluids could not precipitate calcite without significant changes in fluid chemistry such as loss of CO2 or decreasing salinity that overwhelmed the effect of increasing solubility with decreasing temperature. Fluids rising from depth and simply cooling could not be responsible for the nearly ubiquitous calcite veins observed in fault rocks. Loss of CO2 from decompression of the fluid or fluid heating are the likely mechanisms for calcite precipitation. Mass and volume of calcite precipitated from saturated solutions as a consequence of changes in temperature, salinity, and CO2 pressure are shown in Table 2. Increasing temperature from 100º to 150ºC in a 1 m NaCl (5.5 wt.%) solution saturated with calcite would precipitate 6.2 mmoles (0.23 cm3 ) of calcite per kg of fluid. Dilution of a 2 m NaCl solution to 0.2 m NaCl at 100ºC would decrease the solubility by 3.2 mmoles (0.12 cm3 ) of calcite per kg of fluid. Decompression of fluid that is saturated with dissolved carbon dioxide will result in effervescence of CO2 as a separate vapor phase lowering the dissolved CO2 content of the fluid. The fluid pH will rise as a consequence of CO2 effervescence and the solubility of calcite will decrease causing calcite precipitation. PCO2 may commonly be 10 MPa and higher and decrease to 1 MPa during pressure cycling (Table 1). The lower limit of detection of CO2 in fluid inclusions is about 1 MPa, but alteration minerals such as laumontite indicate that CO2 must reach values much below that. At 300ºC a decrease in PCO2 from 10 to 1 and then to 0.1 MPa results in increasing pH from

14 W.T. Parry / Tectonophysics 290 (1998) 1–26 Fig. 4. Photomicrographs of crack sealing and filling minerals from the Wasatch fault, Utah. (A) K-feldspar filling a series of fractures in a larger feldspar phenocryst. The K-feldspar is at extinction with polarizers crossed and is cut by a later calcite vein. Sample is ChC-14 (Parry et al., 1988). (B) Laumontite vein with small amount of calcite in larger plagioclase crystal. Sample is CC-10. (C) Muscovite vein in feldspar from sample ChC-14. (D) Cataclastic vein of comminuted feldspar and quartz crystals in a feldspar grain in sample CC-4.

W.T. Parry / Tectonophysics 290 (1998) 1–26

15

Fig. 5. Solubility of calcite in mmoles per kg of water as a function of temperature and salinity. (a) Calcite solubility measurements of Ellis (1963) at various salinities compared with calculated solubilities (solid lines). CO2 pressure is 1.2 MPa and total pressure is for liquid–vapor equilibria at each temperature. (b) Temperature–salinity projection contoured in terms of calcite solubility in mmoles per kg of water. Fugacity of CO2 is 1 MPa and total pressure is 100 MPa. Contours are 10, 5, 3, 2, and 1 mmole of calcite per kg of water, respectively. Fluid-inclusion temperatures and salinities plotted are as follows. Squares, NaCl–H2 O fluid inclusions from the Wasatch fault, Utah (Parry et al., 1988). Shaded rectangles are: A D fault bend fold (Srivastava and Engelder, 1990); B D Rector Branch thrust (O’Hara and Haak, 1992); C D Gavarnie thrust (Grant et al., 1990; Banks et al., 1991); D D White Oak Mountains thrust (Foreman and Dunne, 1991).

Fig. 5 (continued). (c) Calcite solubility as a function of molality of NaCl; contours are temperature in ºC, CO2 pressure is 1 MPa, and total pressure is 100 MPa.

Fig. 6. Solubility of calcite in mmoles per kg of water as a function of carbon dioxide fugacity and temperature. Total pressure is 100 MPa and salinity is 1 m NaCl.

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Table 2 Calcite solubility changes from fluid characteristic variations Temperature (ºC)

Salinity (molality, m)

Increasing temperature 100 to 150 150 to 200 200 to 250 250 to 300

(ºC) 1 1 1 1

Decreasing salinity (molality) 100 4 to 2 100 2 to 0.2 200 4 to 2 200 2 to 0.2 Decreasing PCO2 (MPa) 100 1 100 1 200 1 200 1 300 1 300 1

PCO2 (MPa)

Solubility (mmoles=kg)

Change (mmoles=kg)

Pore volume (cm3 =kg)

1 1 1 1

13.6 7.4 4.3 2.6

7.4 4.3 2.6 1.7

6.2 3.1 1.7 0.9

0.23 0.11 0.06 0.03

1 1 1 1

15 14.4 6.1 5.0

14.4 11.2 5.0 2.8

0.6 3.2 1.1 2.2

0.02 0.12 0.04 0.08

30.3 13.6 9.4 4.3 3.8 1.7

13.6 6.2 4.3 2. 1.7 0.9

16.7 7.4 5.1 2.3 2.1 0.8

0.62 0.27 0.19 0.08 0.08 0.03

10 to 1 1 to 0.1 10 to 1 1 to 0.1 10 to 1 1 to 0.1

5.4 to 6.1 then to 6.7 and decreasing calcite solubility from 3.8 mmoles per kg to 1.7 mmoles and then to 0.9 mmoles per kg of fluid and precipitation of 2.1 mmoles (0.08 cm3 ) then 0.8 mmoles (0.03 cm3 ) of calcite (Table 2). Significantly the precipitation of calcite cannot be produced by decrease in temperature with uplift of the footwall, because decreasing temperature increases calcite solubility. The increase in solubility of calcite with increasing salinity, Fig. 5c, shows isothermal curves that are convex upwards. As a consequence, isothermal mixing of high- and low-salinity NaCl solutions will always produce a mixture that lies below the saturation curve and would be undersaturated with calcite. However, mixing solutions of different temperatures and salinities could result in calcite precipitation depending on the combination of salinity, temperature and mixing fraction. Fluid-inclusion observations show that fault fluids sometimes include a significant CaCl2 component, and NaCl solutions with dissolved CO2 contain HCO3 as a consequence of hydrolysis of the CO2 . Mixing of these two types of fluids would result in calcite precipitation. For example, NaCl–CaCl2 brine observed often in fluid inclusions that mixed with NaCl–HCO3 also observed in fluid inclusions would be expected to precipitate calcite. Quantities

of calcite precipitated upon mixing would be determined by chemical properties of each solution and mixing fractions. Example calculations using the computer programs SOLVEQ and CHILLER (Reed, 1982; Spycher and Reed, 1989) at 100ºC and a pH of 6 show that mixing Na–Ca–Cl brine containing 0.3 m NaCl and 0.3 m CaCl2 with a Na–Cl–HCO3 solution containing 0.4 m Na, 0.36 m Cl and 0.1 m HCO3 .PCO2 D 0:64 MPa) would precipitate 4.4 g of calcite (1.6 cm3 ) per kg of solution with 0.2 mass fraction of Na–Ca–Cl brine up to a maximum of 6.1 g of calcite (2.3 cm3 ) per kg of solution with 0.4 mass fraction of Na–Ca–Cl brine (Table 3). Mixing a Na–Cl–HCO3 solution containing 0.3 m HCO3 from equilibration at higher PCO2 of 1.9 MPa with the Na– Ca–Cl brine results in precipitation of 13.3 g of calcite (4.9 cm3 ) per kg of solution at equal fractions of each solution (Table 3). Other solution compositions and mixing fractions would result in substantially different quantities of calcite precipitation. Calcite in veins that seal fractures could have precipitated by several mechanisms. A rock-fracturing event would open new porosity, lower fluid pressure, and cause effervescence of dissolved CO2 . The CO2 loss would result in a decrease in calcite solubility and resultant precipitation. Calcite could also precipitate from cool solutions that are circulating

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Table 3 Calcite precipitated on mixing Na–Cl–HCO3 solution with Na–Ca–Cl solution at 100ºC Species

Solution 1 molality

Solution 2 molality

Solution 2 fraction

Na Ca Cl HCO3 pH PCO2

0.4 1 ð 10 6 0.36 0.1 6 0.64 MPa

0.3 0.3 0.9 1 ð 10 6

0.2 0.4 0.6 0.8 1.0

4.4 6.1 5.7 5.1 4.6

1.6 2.3 2.1 1.9 1.7

Na Ca Cl HCO3 pH PCO2

0.3 0.3 0.9 1 ð 10 6

0.4 0 0.1 0.3 6 1.9 MPa

0.2 0.4 0.6 0.8 1.0

4.6 8.1 10.7 12.5 13.3

1.7 3.0 3.9 4.6 4.9

6

6

Calcite precipitated=kg mix g

cm3

Speciation and mixing calculations using the computer programs SOLVEQ and CHILLER (Reed, 1982; Spycher and Reed, 1989).

into warmer regions of a fault, or mixing of diverse solutions of different compositions. 4.2. Quartz The effects of temperature and salinity on quartz solubility are shown in Fig. 7 and the effects of temperature and pressure are shown in Fig. 8. Precipitation of quartz is favored by decreasing temperature, salinity, and pressure. Quartz solubility is unaffected by salinity up to about 300ºC. Above 300ºC solubility is increased by increases in salinity as shown in Fig. 7 (Fournier, 1985a) affecting the sealing of fractures by precipitated quartz; for example, at 350ºC quartz solubility is increased from 800 mg=kg to 1000 mg=kg as a consequence of increasing salinity from 1 to 10 wt.% NaCl at liquid–vapor equilibrium pressure for pure water (16.5 MPa). A temperature change of approximately 40ºC results in a similar change in solubility at 350ºC. Typical fluid-inclusion salinities and homogenization temperatures for the Wasatch fault, Utah are also plotted on Fig. 7, and they indicate that fluid-inclusion characteristics fall within the range of quartz solubilities that are affected by temperature and salinity. Mass and volume of quartz precipitated as a consequence of decompression of fault fluids is shown in Table 4. Decompression of fluids from lithostatic to hydrostatic pressure at 300ºC decreases quartz

solubility from 1250 mg=kg to 800 mg=kg (Fig. 8). Precipitation of 450 mg of quartz would seal 0.17 cm3 of pore space in the rocks per kg of fluid. Fig. 8 also shows the fluid pressure variations indicated for the Wasatch fault, Utah at constant temperature (Parry and Bruhn, 1990). Pressure variations at 285º, 295º, 305º, and 325ºC result in quartz solubility decreases of 220, 350, 336, and 259 mg=kg of fluid, respectively (Table 4). Precipitation of this dissolved silica as quartz in fractures would seal 0.084, 0.132, 0.127, and 0.098 cm3 of pore space in the rocks per kg of fluid. Fluid decompression is thus seen to result in precipitation of significant quantities of quartz and sealing of significant porosity. 4.3. Mineral equilibria Temperature and pressure changes in fault fluids affects equilibria among minerals and fault fluid. On the Dixie Valley fault, Nevada (Parry et al., 1991), hydrothermal alteration of granitic host rocks consists of temporally and spatially overlapping mineral assemblages. Chlorite and epidote are replaced by hydrothermal muscovite (sericite); laumontite, stilbite, and kaolinite are later than the muscovite. These alteration assemblages are distributed discontinuously laterally along the fault (Parry et al., 1991). K-feldspar is a common alteration and vein-filling mineral on both Wasatch and Dixie Valley

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Fig. 7. Solubility of quartz as a function of temperature and salinity (heavy lines) and fluid densities in g=cm3 of fault fluids (light lines). Squares are salinity and homogenization temperature of NaCl–H2 O fluid inclusions on the Wasatch fault, Utah.

Fig. 8. Calculated solubility of quartz in mg per kg of water as a function of temperature and pressure. Also shown are lithostatic and hydrostatic pressure gradients and fluid pressure cycles observed on the Wasatch fault, Utah, by Parry et al. (1988) and Parry and Bruhn (1990) in NaCl–H2 O–CO2 fluid inclusions.

W.T. Parry / Tectonophysics 290 (1998) 1–26

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Table 4 Quartz solubility changes from fluid pressure variations Temperature (ºC) 285 295 305 325

Pressure (MPa)

Solubility (mg/kg)

min.

max.

min. pres.

max. pres.

120 145 170 210

210 265 272 274

746 877 1025 1354

966 1227 1361 1613

faults. Fig. 4A shows fine veinlets of precipitated K-feldspar precipitated in fractures in footwall rocks of the Wasatch fault, Utah. The equilibrium stability of K-feldspar with respect to other alteration product minerals is shown in Fig. 9. The stability of K-feldspar is dependent on solution composition. Muscovite (sericite), clinozoisite (epidote) and K-feldspar are commonly observed minerals in fault rocks. Calcite saturation is also shown on the diagram at two fixed values of PCO2 , 1 and 10 MPa. If PCO2 is reduced due to fluid decompression and effervescence of CO2 during fault rupture, the fluid

Change (mg/kg)

Pore volume (cm3 )

220 350 336 259

0.084 0.132 0.127 0.098

pH increases. The rise in pH increases the value of log aKC =aHC , log aCa2C =.aHC /2 ; and log aNaC =aHC so that solution composition changes in the direction of the arrow in Fig. 9 resulting in precipitation of K-feldspar or albite depending on the relative values of log aKC =aHC and log aNaC =aHC . CHILLER simulations of decompression and decrease in PCO2 from 10 to 1 MPa at 300ºC in a Na–K–Cl–HCO3 solution containing 1.5 m of Cl, 0.7 m of K, 0.8 m of Na, and 1.4 m of HCO3 at a pH of 5.4 (approximately 15 wt.% dissolved solids) initially saturated with calcite and muscovite results in precipitation of 0.123 cm3 of calcite with 0:81 ð 10 4 cm3 of K-feldspar per kg of water on decompression as the pH rises to 6.1. A 1 m NaCl (5.5 wt.%) solution initially saturated with calcite and muscovite precipitates 0.08 cm3 of calcite (Table 2) along with 2 ð 10 4 cm3 of albite. Precipitation of K-feldspar from acid solutions at temperatures below 300ºC is also favored by increasing temperature (Helgeson, 1992). Increasing temperature is unlikely in the footwall of a normal fault so fluid decompression is the likely mechanism for K-feldspar precipitation. Fig. 10 illustrates additional representative temperature- and pressure-dependent mineral equilibria observed in altered rocks on faults (Parry et al., 1988, 1991). Univariant chemical reactions illustrated include: Kaolinite C 2 Quartz D Pyrophyllite C Water

(13)

2 Laumontite D Prehnite C Kaolinite Fig. 9. Stability relations among K2 O–CaO–Al2 O3 –SiO2 –H2 O minerals as a function of aqueous species activities at 300ºC and 100 MPa. The consequences of depressurization of the fluid lowering PCO2 from 10 MPa to 1 MPa is shown by the circles connected by the shaded arrow. The position of each circle corresponds to calcite saturation. Equilibrium constants used in constructing the diagram were calculated using the computer program SUPCRT92 (Johnson et al., 1992).

C 3 Quartz C 5 Water

(14)

Heulandite D Laumontite C 3 Quartz C 2 Water (15) Stilbite D Heulandite C Water

(16)

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Fig. 10. Univariant mineral equilibria as a function of temperature and pressure. Data for reactions: Kaolinite C Quartz D Pyrophyllite C Water from Hemley et al. (1980); 2Laumontite D Prehnite C Kaolinite C 3Quartz C 5Water from Bird and Helgeson (1981); Stilbite D Heulandite C Water and Heulandite D Laumontite C 3Quartz C 2Water from Cho et al. (1987); Stilbite D Laumontite C 3Quartz C 3Water from Liou (1971); and Analcime C Quartz D Albite C Water from Liou et al. (1985). Also shown are typical lithostatic and hydrostatic pressure gradients for a thermal gradient of 35ºC per km and the liquid–vapor (boiling) curve for a 1% NaCl solution.

Stilbite D Laumontite C 3 Quartz C 3 Water

(17)

Analcime C Quartz D Albite C Water

(18)

Quartz-rich rocks could not contain kaolinite above about 270º to 290ºC, pyrophyllite or muscovite would be stable depending on solution composition. The presence of kaolinite with quartz in fault rocks thus indicates temperatures below 270º to 290ºC. Laumontite is not stable relative to prehnite, kaolinite, quartz and water above about 220ºC. Heulandite is stable only above 60 MPa in the temperature range 140º to 190ºC so that equilibria among stilbite, laumontite, and solution is possible below 60 MPa from 125º to 140ºC. Spotty and discontinuous distribution of these minerals along faults such as shown by Parry et al. (1991) suggests that isolated fluid reservoirs are present at the appropriate depths on the faults for each mineral assemblage to form. Fault fluids that contain substantial Ca favor the precipitation of laumontite and prehnite. CHILLER simulations of reaction of a fault fluid containing 0.2 m Na and 0.07 m Ca (2 wt.% NaCl equivalent)

with feldspathic wall rocks indicates precipitation of approximately 0.10 cm3 of prehnite and 0.13 cm3 of laumontite following reaction with 4 g of wall rock per kg of water. The pH of solutions in equilibrium with albite, K-feldspar, and muscovite is a function of solution salinity because of the effect of salinity on activity coefficients of aqueous species in the reaction (Henley and Brown, 1985). The pH of fluids in equilibrium with quartz, K-feldspar, albite, and muscovite was calculated using equilibrium constants calculated for reactions 19 and 20 below with SUPCRT92 (Johnson et al., 1992) and activity coefficients calculated using equation 298 in Helgeson et al. (1981). The calculations are illustrated in Fig. 11 in the temperature range 100ºC to 400ºC at 1 kbar fluid pressure for salinities in wt.% NaCl. 3 KAlSi3 O8 C 2 HC D KAlSi3 O10 .OH/2 C 6 SiO2 C 2 KC KAlSi3 O8 C NaC D NaAlSi3 O8 C KC

(19) (20)

W.T. Parry / Tectonophysics 290 (1998) 1–26

Fig. 11. Calculated pH as a function of fluid salinity for equilibrium among albite, K-feldspar, and muscovite. Salinity and homogenization temperature of NaCl–H2 O–CO2 fluid inclusions from the Wasatch fault, Utah, are plotted as squares.

Fig. 11 shows that pH is lower for higher salinities and higher temperatures. Fluid-inclusion temperatures and salinities for NaCl–CO2 –H2 O fluid inclusions on the Wasatch fault are shown for comparison. These fluids would lie in the pH range 5 to 6 in equilibrium with feldspars and muscovite consistent with pH calculated for equilibrium of these fluids with calcite. Muscovite, K-feldspar, with calcite veins are commonly observed alteration products on the Wasatch fault, Utah (Parry et al., 1988, 1991). 5. Discussion and conclusions Fluid inclusions and hydrothermal alteration minerals record the temperature, pressure and composition of aqueous fluids at depth on faults in a variety of host rocks. Fluid composition is extremely variable from near fresh water to high salinities of 39.2 wt.% NaCl and include CaCl2 and CO2 . A typical fault fluid does not exist. Fluid variability is as great

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within footwall rocks of an individual fault as among all of the faults studied. Also, fluid compositions vary at similar minimum temperatures and pressures on individual faults. Solubilities of vein minerals and hence fracture sealing in the faults is dependent on fluid temperature, pressure, and composition. Quartz solubilities increase with temperature and pressure and at elevated temperature with salinity suggesting that sealing fractures by quartz precipitation is accomplished by cooling or pressure reduction. For example, the footwall damage zone on the Wasatch fault is 20 to 200 m thick and might typically contain fractures 1 m long with cm to dm spacing (Bruhn et al., 1994). Assuming 1 m3 of rock contains 30 of these fractures with an average aperture of 330 µm, then the fracture porosity is 1%. Assuming the 104 cm3 of fracture pore volume is filled with quartz-saturated saline fluid at 300ºC, and the fluid density is 0.82 g=cm3 , then decompression of the fluid from 272 to 170 MPa (Table 4) would precipitate 0.127 cm3 of quartz per kg of fluid. The precipitated quartz would fill 1.04 cm 3 of fracture porosity in 1 m3 of fractured rock, and 104 pore volumes or 8 ð 104 kg of fluid would be required for complete sealing of all of the fractures. Calcite solubilities decrease with increasing temperature, increase with increasing PCO2 and salinity suggesting that calcite fracture sealing is accomplished by loss of CO2 or decreases in salinity. Decreasing PCO2 from 10 MPa to 1 MPa lowering calcite solubility and precipitating 0.62 cm3 of calcite from 1 kg of solution would fill 5.1 cm3 of fracture porosity in 1 m3 of the fractured rock described above. Approximately 2000 pore volumes or 1:6 ð 104 kg of fluid would completely seal the fractures. A decrease in pressure can decrease the solubility of both quartz and calcite, thus, coprecipitation of calcite and quartz suggests de pressurization of the fluid. Loss of CO2 also results in increase in pH which can cause precipitation of K-feldspar or albite depending on solution composition. However, the volume of feldspar precipitated is small compared to the volume of calcite precipitated. Fault solutions that contain substantial Ca are common (Table 1) and favor precipitation of laumontite (Fig. 4B) and prehnite in volumes that are comparable to calcite precipitation.

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The characteristics of fault fluids determined from fluid-inclusion observations support the fault models of Sibson et al. (1988); Boullier and Robert (1992); Byerlee (1993), and Robert et al. (1995) in three ways. First, fluid-inclusion characteristics are consistent with episodic decompression, effervescence of CO2 , and precipitation of calcite and quartz filling fractures associated with repeated fault rupture. Second, wall rocks in reverse faults hosting gold– quartz veins are depleted in SiO2 and fault rocks are enriched in SiO2 as a result of pressure solution of quartz and albite and diffusional mass transport during shear zone deformation, but most vein quartz is derived by fluid movement into lower-pressure reservoirs. Chemical composition and densities of fluid-inclusion fluids are consistent with cyclical fluid pressure fluctuations. In other faults, the damage and core zones are both depleted in silica and other mobile elements and enriched in immobile elements due to fluid-assisted mass transport (Goddard and Evans, 1995; Evans and Chester, 1995). Third, fluids with diverse compositions are present on individual faults without apparent mixing. Isolated fluid reservoirs may have been present, although different fluids may not have been present at the same time. Implicit in the fault-valve model of Sibson et al. (1988); Boullier and Robert (1992), and Robert et al. (1995) is episodic flow of fluids into the fault zone from surrounding rocks, fracture sealing, fluid pressure increase, fracturing, decompression and effervescence of fluids within the fault zone and flow of fluids into surrounding rocks. Acknowledgements Financial support was provided by NSF grants EAR-8618250 and EAR-9104677 and the U.S. Geological Survey Earthquake Hazards Reduction Program Grant 14-08-0001-G-886. We are grateful to J.P. Evans and two anonymous Tectonophysics reviewers for careful review and many helpful suggestions. References Banks, D.A., Davies, G.R., Yardley, B.W.D., McCaig, A.M., Grant, N.T., 1991. The chemistry of brines from an Alpine thrust system in the Central Pyrenees: An application of fluid

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