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4 Fluid Dynamics During Progressive Regional Metamorphism JOHN V. WALTHER Northwestern University INTRODUCTION A major part of active tectonic processes is the evolu- tion of continental margins with the progressive burial and metamorphism of lavas and sedimentary strata to deep crustal levels. The nature of fluid components in these rock sequences buried to mid or lower crustal depths (10 to 40 kin) is poorly known but of utmost importance. The presence of fluid changes the theological properties of the rock and thus its response to stress. The fluid also has the potential to carry heat and dissolved mineral components and can therefore dictate the style of metamorphism, the extent of chemical equilibrium, and much about the tex- tural fabric observed in rocks. It has become abundantly clear that the majority of sediments and rocks in active tectonic processes in the upper few kilometers of the Earth's surface have experienced large fluid fluxes. The nature of the flow and its chemical and physical consequences are addressed in other chapters in this volume. This chapter addresses problems regarding the origin, flow dynamics, and chemical consequences of fluid at mid-to-lower crus- tal depths. Classically, to those who study the mineralogi- cal changes in such rocks, the general process is termed progressive regional metamorphism. All mountain fold belts contain lavas and sediments that have been region- ally metamorphosed. 64 FLUID-ROCK RATIOS AND FLUID FLUX It is not surprising that large fluid fluxes can be re- corded within the upper few kilometers of the Earth's surface where the porosity and/or permeability of many rocks are large and dramatic temperature and topographic gradients can occur. Perhaps more surprising is the evi- dence for large amounts of fluid flow occurring in deeply buried rocks. Essentially this evidence consists of docu- menting chemical changes in mineral assemblages during metamorphism and calculating the minimum fluid volume necessary to account for the observed changes, as shown in Figure 4.1. Perhaps the most straightforward calculation of this type is documentation of the extent of reaction for a simple decarbonation such as the production of wollastonite from calcite and quartz: CaCO3 + SiO2 CaSiO3 + CO 2 2. At fixed pressure and temperature this reaction fixes the fugacity of CO2 in the fluid. Thus, the concentration of H2O or other components in the fluid needed to maintain the fugacity of CO2 for the extent of production of wollas- tonite can be calculated (e.g., Rumble et al., 1982). The calculated minimum fluid necessary to produce the observed chemical changes seen during metamorphism is
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FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM ' t /~/ ~- :~? ;~1 - Fluid flowing along grain boundaries influences phase equilibria Instantaneous FRR<~0.01 1 1 Mineral assemblage after meta morphism t Mineral assemblage before meta morphism termed the time-integrated fluid-rock ratio or more com- monly the fluid-rock ratio (ERR). FRRs, calculated by a number of approaches from mineral assemblages in basal- tic or carbonate units, are as high as 2 to 20 (e.g., Ferry, 1976, 1980;Grahametal.,1983;Tracyetal.,1983~. That is, 2 to 20 volumes of fluid have reacted with each volume of rock during progressive metamorphism. Note that these are minimum values since fluids in equilibrium with the mineral assemblage are unrecorded by this technique. Additionally, there is no guarantee that the fluid has reached complete chemical equilibrium with the mineral assem- blage, an assumption inherent in the calculation. How- ever, given the high temperatures involved and the large surface area to fluid volume in metamorphic rocks, the equilibrium assumption is probably reasonable (Walther and Wood, 1986~. Because the actual grain boundary porosity of metamorphic rock is probably less than 0.1 percent, these high Fl(Rs indicate that fluid within the flow porosity must have been replenished thousands of times if the introduction of fluid does not expand the rock. Fluid-rock ratios are often calculated based on changes in the isotopic composition of oxygen in minerals. Unfor- tunately, these changes are sensitive to fluid-rock interac- tions at any point in the rock's history, that is, during its burial to peak metamorphic conditions, at the peak of metamorphism, or during its uplift to the Earth's surface. In some cases isotopic exchange indicating substantial fluid-rock interaction may be recording exchange that has occurred near the Earth's surface either during burial or uplift. Thus, they may not be recording FRRs during the peak of metamorphism. Wickham and Taylor (1985; Chapter 6, this volume) have suggested that high-grade metamorphic rocks from the Pyrenees appear to require 65 Total volume of fluid needed FIGURE 4.1 A comparison of the amount of fluid within a rock during metamor- phism (left) and the fluid-rock ratio (FFR) calculated from mineral equilibria (right) (from Wood and Walther, 1986~. the introduction of a large volume of seawater deep into the metamorphic pile. Such a flux of fluid to great depths in unlikely (Walther and Orville, 1982; and as discussed below). The data are, however, also consistent with oxy- gen isotope exchange nearer the Earth's surface. Fluid-rock ratios are not a measure of the total time integrated fluid flux (IFF) during metamorphism but must be considered in the context of their ability to place limits on the IFF. First, the volume of rock affected must be known, but unfortunately it is often difficult to determine. It is obvious, however, that the same total fluid volume reacting toward equilibrium with a thin rock unit will record a higher FRR than an identical unit that is thicker, because the extent of reaction in the thinner unit is greater per unit rock volume. These problems have been discussed in more detail elsewhere (Wood and Graham, 1986; Wood and Walther, 1986~. - Fluid-rock ratios do not record the passage of fluid in equilibrium with a rock. This has important ramifications. Consider a typical divariant (sliding) reaction in a politic rock undergoing metamorphism: 2HCl + CaCO3 + 2NaAlSi3O~ (calcite) (in play.) CaAl2Si2O6 + 2NaCl + 4SiO2 + CO2 + H2O. (in play.) (quartz) Assuming that calcite and quartz are pure phases, at fixed pressure, temperature, and chloride activity the composi- tion of the plagioclase will dictate the equilibrium ratio of CO2 and H2O in the fluid. Imagine a variety of composi- tional layers, perhaps of sedimentary origin, with varying plagioclase composition. A fluid passing through such
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66 layering will react to adjust its CO2 to H2O ratio to ap- proach equilibrium with each plagioclase crystal it makes contact with. The.extent of reaction will depend then on how far the plagioclase composition is from equilibrium with the fluid. For the same IFF the recorded FRR will be different depending on the variability of plagioclase com- position. One may then be drawn to the wrong conclusion that some beds experienced large IFFs while others did not. Thus, it would erroneously appear that some beds acted as metamorphic fluid aquifers while others were aquitards. This could then lead to a model of the major flow occurring along bedding when in fact it may not (Ferry, 1987~. However, a number of stable isotopic stud- ies have apparently demonstrated that bed-parallel fluid flow occurs. What will be argued here is that it is not differences in intrinsic permeabilities but differences in the theological properties of the different layers that con- trol fluid flow. FLUID FLOW AT MID- TO LOWER-CRUSTAL LEVELS Fluid-rock ratios at depth are apparently similar to those determined near the Earth's surface, where fluid convec- tion is known to operate. It has, therefore, been proposed that fluid convects to deep crustal levels (e.g., Etheridge et al., 1983; Wickham and Taylor, 1985; Ferry, 1986~. Arguments to the contrary have also been presented (Walther and Orville, 1982; Walther and Wood, 1984; Wood and Walther, 1986~. Let us assess some of the evidence for the state and transport of fluid at mid or lower crustal depths. Fluid flow is obviously highly dependent on permeabil- ity. Permeabilities of metamorphic rocks are not well known. What is clear is that they increase dramatically in laboratory experiments when fluid pressure equals rock pressure (Brace, 1980~. Since most of the experimental studies do not concern themselves with large fractures, such permeability estimates are considered by many to be minimal for the crust as a whole. Use of these permeabil- ity values suggests that fluid flow during metamorphism should occur under conditions of fluid pressure signifi- cantly below rock pressure. That is to say that for the anticipated fluid flux the permeabilities during metamor- phism are great enough to allow fluid pressure to drop below the rock pressure toward a more hydrostatic gradi- ent. A hydrostatic gradient is required to promote convec- tion. By hydrostatic pressure what is meant is the pressure resulting from the density of an overlying column of fluid as opposed to the much greater pressure resulting from the density of an overlying column of rock (lithostatic pres- sure). Fluid pressure less than lithostatic does not seem to be the case when metamorphic rocks are closely examined JOHN V. WALTHER (Norris and Henley, 1976; Fyfe et al., 1978). Phase equi- librium studies of devolatilization reactions observed in metamorphic rocks generally seem to require fluid pres- sure to be near lithostatic pressure. This observation is confirmed by fluid inclusion studies where the trapped fluid has the appropriate density for fluid pressure equal to lithostatic pressure at the temperature of metamorphism. These two divergent observations could be rectified if the properties of the fluid phase were not those of the bulk fluid. It has been suggested, for example, that the fluid phase present during deep crustal metamorphism is not a discrete fluid phase but rather an absorbed phase on min- eral surfaces with presumably greater viscosity and lower fugacity (Elliott, 1973; Rutter, 1976~. Studies at room temperature indicate that multimolecu- lar layers of H2O are absorbed on mineral surfaces. At metamorphic temperatures and pressures where even struc- turally bound water becomes unstable in hydrous miner- als, it is reasonable to assume that absorption of H2O is no more than a monolayer in thickness. If we examine the amount of fluid trapped as fluid inclusions along healed microcracks during metamorphism and redistribute it evenly along the entire crack, the width of fluid in the crack is generally about 200 ~ or 10 times the thickness of a double monomolecular layer (Walther and Orville, 1982~. Because this is the minimum thickness of the fluid film at the time the crack healed, we may reasonably assume that the properties of fluid flowing through such cracks are not significantly modified by absorptive properties of mineral surfaces, and hence the fluid phase can be considered to have the properties of a bulk fluid. Near the Earth's surface fluid pressure is controlled by the extent of the overlying fluid column (hydrostatic pres- sure) while the pressure exerted on mineral grains (lithostatic pressure) is considerably greater owing to the much greater density of the- minerals. This difference between fluid pressure and rock pressure is maintained by the effective crushing strength of the rock. At some depth the closure of pores and the resultant decrease in permea- bility causes the fluid pressure gradient to increase dra- matically, so that fluid pressure equals lithostatic pressure. Note that fluid pressure equal to rock pressure does not imply that fluid is trapped in a static state but that it is possible that the flow of fluid is balanced by permeability changes near fluid pressure equal to rock pressure. Figure 4.2 shows fluid pressure as a function of depth determined from well bottom hole fluid pressure measure- ments from a number of wells in the U.S. Gulf Coast. Note that at a depth just below 3 km fluid pressure begins to depart from hydrostatic, and at 5.5 km or at a lithostatic pressure of 1.5 kbar the fluid pressure is very close to lithostatic. The fluid pressure below 3 km is greater than that exerted by an overlying column of fluid and is there
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FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM y '_ 3 1 1 ~ BRAZORIA CO.. TEXAS ~ W 1 _\ '\ LITHOSTATIC \ ~- GRADIENT ~ BOTTOM HOLE 2 \ \ FLUID PRESSURE ~- -x iFLUID PRESSURE =._. ...' ._. . . . 4 _ 5 _ HYDROSTAT C',\ GRADIENT \ i\ .w '\ \ ~ _ -do 6 _ 0 0.2 0.4 0.6 0.8 1.0 1.2 PRESSURE, KBAR FIGURE 4.2 Fluid pressure as a function of depth in a sedimen- tary basin of the U.S. Gulf Coast. Note the crossover from hydrostatic to lithostatic fluid pressure and its implications for fluid convection (from Wood and Walther, 1986~. fore considered "geopressured." There are various chemi- cal and physical factors that influence the depth at which geopressuring of the fluid occurs. The decrease of hydrau- lic conductivity in clay layers is often considered the pri- mary factor in determining the characteristics of the geo- pressured zone (Bredehoeft and Hanshaw, 1968; Chapman, 1972). The depth of onset of geopressuring has been observed from as little as 45 m to depths greater than 8 km, although in most sedimentary basins it is above 6 km. We might imagine that in crystalline rocks in extensional environ- ments hydrostatic gradients in the fluid may be maintained to depths greater than 8 km. However, at a depth of 11 km the difference between hydrostatic pressure and lithostatic 67 pressure is about 2 kbar. Although the crushing strengths of rocks at the temperatures, pressures, and strain rates ,L ~1` appropriate for this depth are not well known, it seems it. ~reasonable to conclude that hydraulic conductivity would ~ en be so greatly reduced by mechanical compaction that fluid _ ~ ° below this depth can in general be considered to be close O to lithostatic pressure. As mentioned above, this observa lion is consistent with the evidence from fluid inclusion studies that indicates that fluid pressure equals rock pres . _ ._ sure during metamorphism. ° It would, therefore, seem that laboratory measurements _ ~of the permeabilities of metamorphic rocks do not charac ~n ~terize the permeability of these rocks at the time they are undergoing metamorphism, but overestimate them. It may ~be that the permeability is reduced by the inevitable re <' crystallization that occurs when these rocks are subject to high temperatures and pressures for time periods on the order of millions of years during metamorphism. In any event it seems that fluid pressure must be close to lithostatic at mid- to lower-crustal depths. The imposition of a pres sure gradient significantly greater than hydrostatic on the fluid during progressive regional metamorphism means that there is no reasonable way to transport the less-dense fluid by flow downward (i.e., no convection of fluid can occur). This means that the large integrated FRRs re corded by some rocks must be recording fluid generated at some greater depth. FLOW MECHANISM WHERE FLUID PRESSURE EQUALS LITHOSTATIC Consider a rock initially devoid of fluid undergoing a devolatilization reaction in response to increased tempera ture during progressive metamorphism at mid-crustal lev els. Fluid will be produced by each volatile-containing mineral undergoing destruction by reaction. Depending on the wetting characteristics of the fluid, it will either coat the mineral surfaces or begin to collect in isolated pores at mineral triple junctions (White and White, 1981; Watson and Brenan, 1987), as shown on the left side of Figure 4.3. Z zap FIGURE 4.3 Fluid production at reacting volatile containing minerals collecting in isolated pores at mineral triple junctions (left). On a larger scale with increased fluid production, these must interconnect, producing a fluid phase of some vertical extent, z (right).
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Representative terms from entire chapter:
68 In any event, continued devolatilization will eventually build an interconnected three-dimensional fluid network of some vertical extent, as shown schematically on the right side of Figure 4.3. A static fluid at lithostatic pres- sure would rise due to its lower density and therefore to its buoyancy relative to the surrounding rocks if it was not held by the tensile strength of the rock. Because rocks at metamorphic conditions have low tensile strengths, the fluid would become mechanically unstable and hydrofrac- ture its way toward the Earth's surface. Similar arguments have been used to explain the ascent of magmas from depth even though the buoyancy forces are less. As the left side of Figure 4.4 shows, the greater the amount of inter- connected fluid space, the greater will be the difference between the pressure on the fluid and surrounding rock and thus the ability of the fluid to hydrofracture the rock. The pressure difference between static fluid and rock is given by UP = gZ(prock - pfluid) ~ (4.1) where g stands for the acceleration due to gravity, z is the vertical distance of interconnectivity of the fluid, and PrOck and Pfluic are the rock and fluid densities, respectively. While the tensile strength of a rock acting to prevent propagation of a fracture at a crack tip is poorly known, we might imagine that under metamorphic conditions, with the interplay of deviatoric stresses during active tecton- ism, it is no more than 10 bars or so, particularly for subcritical crack growth. This means that no intercon- nected fluid network can exist statically that extends in the vertical direction more than about 60 m. This severely limits the possible vertical extent of fluid convection be- fore the tensile strength is exceeded and fluid migrates upward only. Over this short distance no convection is possible for any reasonable temperature gradient and per- meability. The potential to produce an interconnected fluid net- work is great. The approximately 2 moles of fluid on average given off per kilogram of politic composition rock during medium- to high-grade metamorphism has the potential to hydrofracture the rock many times. This amount of fluid would occupy a volume of approximately 12 per- cent under metamorphic conditions if it did not escape. While we do not know the volume of fluid necessary to produce an interconnected fluid pathway, it is probably no more than a few tenths of a percent of the volume of the rock. Thus, we might imagine that even on a local scale hydrofracturing may have occurred a large number of times. If we consider the passage of fluid released from minerals lower in the metamorphic pile, it is not surprising that we observe large numbers of healed fluid-filled microcracks in metamorphic rocks. JOHN V. WALTHER The predominance of divariant (sliding equilibria) devolatilization reactions during metamorphism suggests that fluid is released from hydrous and carbonate minerals more or less continuously in the metamorphic pile during prograde metamorphism. It is possible that some fractures stay open for extended periods of time fed by a continuous supply of fluid. This situation is shown on the right side of Figure 4.4, where we require a viscous pressure gradi- ent, dPViS, to compensate for the difference in the pressure gradient between the static fluid and the surrounding rock. That is, the loss in hydraulic head by the upwardly flowing fluid due to its viscosity compensates for the difference between the hydrostatic and lithostatic gradients so that fluid pressure in the flowing fluid equals lithostatic pres- sure along the walls of the fracture, which in turn allows the fracture to remain open: dPvis = g ~ Ink - pfluid ~ · (4. 2) Assuming fluid and rock densities of 0.9 and 2.8 g/cm3, the viscous pressure gradient required as calculated from Eq. (4.2) is about 2.0 X 103 dyne/cm3. We can model the fluid flow in a microcrack as steady-state incompressible viscous laminar flow through two parallel plates that ex- tend in all directions to a very much greater extent than d, the distance they are apart. Due to the viscous nature of the fluid, the profile of the flow velocity across the crack is parabolic, zero at the wall, and a maximum in the center. The solution to this fluid mechanical problem is 2vmax d dPvis ,, ~ ~ A ~ 4 . 3 3 12v where v and vmax are the average and maximum fluid velocities, respectively, d is the fracture width, and v is the viscosity of the fluid. Noting that the fluid flux, q, is equal Pf>Priz+ / Pf=Pr Ny I Pf
FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM to the cross-sectional area of the fracture opening times the average velocity, we have q= vdl = HI d Pvis (4.4) 12v where I is the length of the fracture opening. Viscosities of supercritical H2O-CO2 mixtures are gen- erally between 0.1 and 0.2 centipoise. For a given flux of fluid we can calculate the length of cracks perpendicular to flow versus their width per square centimeter. Such calculations have been done (Walther and Orville, 1982; Walther and Wood, 1984) and indicate that the widths of the fractures are in most cases less than 105 A. Because of the cubic dependence of fluid flux on crack width, it seems likely that changes in the flux of fluid are accommodated by small changes in fracture width to maintain the viscous pressure gradient at the value neces- sary to keep the fracture open. Apparently, judging from the extent of fluid inclusions along sealed fractures in metamorphic minerals, if the fracture width falls below about 200 A the crack will seal. Laboratory investigations indicate that these cracks seal in a matter of days, at least in quartz. It stands to reason that, if the fluid flux through the metamorphic pile is not continuous, many generations of these microcracks will form. If such a mechanism of fluid flow operates, the useful- ness of the concept of intrinsic permeability of a rock is questionable. That is, the permeability is a dynamic func- tion of the fluid flux through the metamorphic rocks. The permeability of the rock is adjusted by the fluid phase to accommodate the flux of fluid, so the fluid pressure is near lithostatic during fluid flow. This is different from the concept of reaction-enhanced permeability due to volume loss of solids because of reaction- (Rumble and Spear, 1983~. What is argued here is that as a general approxima- tion the permeability adjusts itself, so fluid pressure is always close to rock pressure irrespective of the extent of reaction. With the deviatoric stresses that operate during meta- morphism and the differences in tensile strength of differ- ent layers of rock, fracture production and, therefore, fluid channels may develop along preferred layers that are at some angle to the Earth's gravitational field. This would give rise to layers that appear as "metamorphic aquifers" and others that appear as aquitards, at least over limited distances. CHANNELIZATION OF FLUID FLOW The interconnectivity, bifurcation, coalition, or distri- bution of cracks in metamorphic rocks are largely un- known. We do know that some fractures or fracture net 69 works have remained open long enough to experience the passage of a considerable amount of fluid. In politic rocks these fractures are often marked by quartz veins. At lower temperatures calcite veins are dominant because calcite has a higher solubility at low temperatures than does quartz. The thickness of the vein represents the accumulated ef- fects of mineral precipitation from the passing fluid. If such major channelways of fluid escape are present, the flow paths of these fluids must to some extent coalesce as fluid is produced at each mineral in the rock that under- goes devolatilization. Thus, there must be some flow of fluid in intimate contact with minerals before fluid enters a major channelway. Obviously, the extent of grain bound- ary flow versus flow in major channelways controls to a large extent the amount of fluid a particular rock unit may experience during metamorphism and hence IFF. This in turn dictates much about the textural fabric of the rock and the approach to chemical equilibrium that may be ex- pected. Figure 4.5 shows a quartz segregation/vein in the Bund- nerschiefer formation of Switzerland that is considered to have formed during Lepontine metamorphism. If this quartz segregation represents the cross section of quartz deposited along a major conduit for fluid flow, we can calculate the amount of fluid that must have passed to cause the extent of quartz precipitation seen. Because quartz is present in most of the mineral assemblages in the Bundnerschiefer, it is anticipated that fluids responsible for the deposition of quartz along the segregation/vein are at quartz saturation. Let us calculate the fluid necessary to precipitate a quartz vein 50 cm in diameter, much like the one shown in FIGURE 4.5 Quartz segregation thought to represent the cross section of quartz precipitated during the lifetime of a major fluid conduit during metamorphism. While the width of the fracture was probably less than 10 ~m, the large amount of flow precipi- tated substantial quantities of quartz.
70 Figure 4.5. For a distance of 1 cm along the vein this amounts to 2000 cm3 of quartz or about 87 moles. At 500°C and 4 kbar with a geothermal gradient of 20° to 30°C~n, this requires 4 x 108 moles of H2O or the dehy- dration of 7.6 x 10~° cm3 of average politic rock during metamorphism if 2 moles of H2O are released per kilo- gram of rock. This corresponds to a cube of petite 42 m on a side. Given the large volumes of politic and psarnmitic rock in most metamorphic terrains, the fluid fluxes must be high. For the whole of the metamorphic pile, fluid fluxes of 1 x 1~° to 1 x 10~9 g cm-2 s~i have been calculated (Walther and Orville, 1982), which means that an average of 3 to 30 kg of fluid must pass through each square centimeter of crust overlying the 400°C isotherm in each million years of progressive metamorphism. The upward flow of fluid may be even greater if significant fluid is released from a subducting slab during metamorphism (perhaps through a magma intermediary) or if mantle fluids (Dawson, 1980) are significant. Until it is determined to what extent fluid is channeled along structural features such as faults, fold hinges, pri- mary lithological contacts, and layers with low tensile strength or, alternatively, flows at the scale of grain bounda- r~es, the extent of heat and material transport and the very nature of metamorphism will not be understood. ACKNOWLEDGMENTS This chapter was written while I was on sabbatical leave at Universite Paul Sabatier, Toulouse, France. I would like to thank Jacques Scholl for his warm hospital- ity. Comments by J. M. Ferry and B. J. Wood led to substantial improvement. REFERENCES Brace, W. F. (1980~. Permeability of crystalline and argillaceous rocks, International Journal of Rock Mechanics and Mineral Sciences 17, 241-251. Bredehoeft, J. D., and B. B. Hanshaw (1968~. On the mainte nance of anomalous fluid pressures: I. Thick sedimentary sequences, Geological Society of America Bulletin 79, 1097- 1106. Chapman, R. E. (1972~. Clays with abnormal interstitial fluid pressures. American Association of Petroleum Geologists Bulletin 56, 790-795. Dawson, J. B. (1980~. Kimberlites and Their Xenoliths, Sprin ger-Verlag, New York. Elliott (1973~. Diffusion flow laws in metamorphic rocks, Geo logical Society of America Bulletin 84, 2645-2664. Etheridge, M. A., V. J. Wall, and R. H. Vernon (1983~. The role JOHN V. WALTHER of the fluid phase during regional metamorphism and deformation, Journal of Metamorphic Petrology l, 205-226. Ferry, J. M. (1976~. P. T. fCO2 and fH2O during metamorphism of calcareous sediments in the Waterville-Vassalboro area, south central Maine, Contributions to Mineralogy and Petrol- ogyS7,119-143. Ferry, J. M. (1980~. A case study of the amount and distribution of heat and fluid during metamorphism, Contributions to Mineralogy and Petrology 71, 373-385. Ferry, J. M. (1986~. Reaction progress: A monitor of fluid-rock interaction during metamorphic and hydrothermal events, in Fluid-Rock Interactions During Metamorphism, J. V. Walther and B. J. Wood, eds., Springer-Verlag, New York, pp. 60-88. Ferry, J. M. (1987~. Metamorphic hydrology at 13 kilometers depth and 500-550°C, American Mineralogist 72, 39-58. Fyfe, W. S., N. J. Price, and A. B. Thompson (1978~. Fluids in the Earth's Crust, Elsevier, Amsterdam, 383 pp. Graham, C. M., K. M. Greig, S. M. F. Shepherd, and B. Turi (1983~. Genesis and mobility of the H2O-CO2 fluid phase during regional greenschist and epidote amphibolite facies metamorphism: A petrological and stable isotope study in the Scottish Dalradian, Journal of the Geological Society of Lon- don 140, 577-599. Norris, R. J., and R. W. Henley (1976~. Dewatering of a meta- morphic pile, Geology 4, 333-336. Rumble, D., and R. S. Spear (1983~. Oxygen-isotope equilibra- tion and permeability enhancement during regional metamorphism, Journal of the Geological Society of London 140, 619-628. Rumble, D., J. M. Ferry, T. C. Hoering, and A. J. Boucot (1982~. Fluid flow during metamorphism at the Beaver Brook fossil locality, New Hampshire, American Journal of Science 282, 886-919. Rutter, E. H. (1976~. The kinetic of rock deformation by pres- sure solution, Philosophical Transactions of the Royal Society of London A283, 203-219. Tracy, R. J., D. M. Rye, D. A. Hewitt, and C. M. Schiffries (1983~. Petrologic and stable-isotopic studies of fluid-rock interactions, south-central Connecticut. I. The role of infiltra- lion in producing reaction assemblages in impure marbles, American Journal of Science 283A, 589-616. Walther, J. V., and P. M. Orville (1982~. Rates of metamorphism and volatile production and transport in regional metamorphism, Contributions to Mineralogy and Petrology 79, 252-257. Walther, J. V., and B. J. Wood (1984~. Rate and mechanism in prograde metamorphism, Contributions to Mineralogy and Petrology 88, 246-259. Walther, J. V., and B. J. Wood (1986~. Mineral-fluid reaction rates, in Fluid-Rock Interactions During Metamorphism, J. V. Walther and B. J. Wood, eds., Springer-Verlag, New York, pp. 194-211. Watson, E. B., and J. M. Brenan (1987~. Fluids in the litho- sphere, 1. Experimentally-determined wetting characteristics of CO2-H2O fluids and their implications for fluid transport, host-rock physical properties, and fluid inclusion formation, Earth and Planetary Science Letters 85, 497-515.
FLUID DYNAMICS DURING PROGRESSIVE REGIONAL METAMORPHISM White, J. C., and S. H. White (1981~. On the structure of grain boundaries in tectonites, Tectonophysics 78, 613-628. Wickham, S. M., and H. P. Taylor, Jr. (1985~. Stable isotopic evidence for large-scale seawater infiltration in a regional metamorphic terrane; The Trois Seigneurs Massif, Pyrenees, France, Contributions to Mineralogy and Petrology 91, 122- 137. An.. 71 Wood, B. J., and C. M. Graham (1986~. Infiltration of aqueous fluid and high fluid-rock ratios during greenschist facies meta- morphism: A discussion, Journal of Petrology 27, 751-761. Wood, B. J., and J. V. Walther (1986~. Fluid flow during meta- morphism and its implications for fluid-rock ratios, in Fluid- Rock Interactions During Metamorphism, J. V. Walther and B. J. Wood, eds., Springer-Verlag, New York, pp. 89-108. .
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Representative terms from entire chapter: