FOURNIER 1999_hydrothermal Proc & Movement Fluids-1

Economic Geology BULLETIN OF THE SOCIETY OF ECONOMIC GEOLOGISTS

VOL. 94 December 1999 No. 8

Hydrothermal Processes Related to Movement of Fluid From Plastic into Brittle Rock in the Magmatic-Epithermal Environment

ROBERT O. FOURNIER

108 Paloma Road, Portola Valley, California 94028

Abstract

Repeated intrusions of magma to depths of 2 to 6 km beneath a volcanic complex result in a large body of surrounding rock attaining temperatures sufficiently hot (>400C) for crustal rocks to behave in a plastic manner at normal strain rates, --10 -14 sec -1. Where the least principal stress is the lithostatic load, hypersaline brine and gas exsolved from crystallizing magma accumulate at lithostatic pressure within this plastic rock in horizontal lenses. A narrow, self-sealed zone of relatively impermeable material separates the lithostatically pressured region from a region where meteoric-derived hydrothermal fluids circulate through brittle rock at hydrostatic pressure. Episodically, major breaches of the self-sealed zone occur, and several different mechanisms may play a role in triggering such a breach. A particularly likely mechanism is an upward surge in magma that temporarily increases the local strain rate to such a degree that previously plastic material undergoes shear failure in response to a very small stress difference. This allows hypersaline brine and gas to be expelled quickly from the normally plastic region into the brittle, lower pressure and lower temperature domain, where epithermal veins are deposited as a result of decompression and cooling of the magmatic fluid. The resulting increase in fluid pressure and temperature within the brittle domain leads to faulting and brecciation that increase permeability and allow an increase in the rate of discharge of hydrothermal fluid. The enthalpy of the relatively dilute supercritical fluid that separates from boiling brine as a result of decompression greatly influ- ences the type of mineralization that occurs when that fluid moves upward through cooler, brittle rock. Re- peated cycles of breaching of the self-sealed zone, healing by mineral deposition and plastic flow, and rebreaching result in banded epithermal veins extending long distances into brittle rock. Eventually, intrusive activity stops or declines and the 400C isotherm moves downward, accompanied by hydrostatically pressured hydrothermal activity dominated by meteoric water.

Introduction

SEISMOLOGISTS and petrophysicists have long recognized the importance of the transition from brittle to plastic behavior in the lithosphere in regard to the bottoming of earthquakes that result from shear failure (e.g., MacElwane, 1936; Byerlee, 1968). They also have recognized that the bottoming of seis- micity (implying the onset of quasiplastic behavior) occurs at very shallow depths beneath large and hot presently active geothermal fields, such as The Geysers and Clearlake Highlands (Majer and McEvilly, 1979, Sibson, 1982), the Imperial Valley (Gilpin and Lee, 1978), and Yellowstone National Park (Smith and Braile, 1984, 1994). In addition to limiting seismic activity, the onset of plastic flow is important because it closes interconnected pore spaces and fractures, thereby controlling the depth of circulation of meteoric water into the crust at hydrostatic pressure (pressure imparted by the weight of an overlying column of water; Brace, 1972).

Meteoric-derived fluids at hydrostatic pressure have been encountered by many deep geothermal exploration wells in continental crystalline rocks at temperatures up to about 350 to 360C. To date, the few deep wells drilled to temperatures

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greater than 370 to 400C either have encountered little permeability or have produced gas-rich brines at greater than hydrostatic pressures (Fournier, 1991). These observations indicate that the brittle-plastic transition commonly occurs at about 370 to 400C within presently active continental hydrothermal systems. The well data also show that a narrow zone or shell of relatively impermeable material commonly separates the hydrostatically pressured domain and a domain in which fluid at greater than hydrostatic pressure may accumulate in apparently quasiplastic rock. Narrow zones separating high- and low-pressure environments also have been found in epithermal ore deposits. For example, at the Far Southeast porphyry Cu-Au deposit in the Philippines there is fluid inclusion evidence that the transition from hydrostatic to lithostatic pore fluid pressure (Pt) occurred over a vertical interval as small as 100 m (Hedenquist et aI., 1998).

Many investigators have recognized that brittle-plastic phenomena and a transition from lithostatic to hydrostatic pressure conditions characterize porphyry Cu deposit mineralization (e.g., Fournier, 1967, Gustafson and Hunt, 1975; Cunningham, 1978). More recently, several papers have related epithermal and skarn mineralization to the escape of lithostatically pressured magmatic fluids across brittle-plastic

Vol. 94, 1999, pp. 1193-1212 0361-0128/99/3007/1193-20 $6.00 1193

1194 ROBERT O. FOURNIER

transition zones into overlying hydrostatically pressured hydrothermal systems (e.g., Fournier, 1987, 1992; Rye, 1993; Deenet al., 1994; Meinert et al., 1997; Hedenquist et al., 1998). In this presentation, mechanisms for accumulating those fluids in plastic rock at lithostatic pressure and for allowing their episodic discharge into the hydrostatically pressured brittle domain are stressed. Also stressed are the con- sequences of accumulating magmatic fluids at different depths and temperatures in plastic rock. A few well-studied ore deposits will be used for illustrative purposes.

The Brittle-Plastic Transition

General considerations

The general relationships for initiating failure of materials in the brittle and plastic regions of the earth's crust are shown schematically in Figure 1. In the brittle region the stress difference required to cause shear failure of a preexisting open crack increases with increasing depth and is relatively independent of temperature, rock type, and strain rate (Jaeger and Cook, 1979). It is, however, highly dependent on the coefficient of friction, on Pf, and on the orientation of the fracture with respect to the stress field (Sibson, 1984). The coefficient of friction for most rocks is about 0.6 to 0.8, according to measurements by Byedee (1978). In Figure i the line from

the origin at zero depth through point A shows the stress difference (l-cr 3) required to activate fault movement on an existing open crack (a crack having no cohesive strength) that is oriented at an optimum angle of about 26 to the direction of application of the maximum principal stress, when Pf is assumed to equal the hydrostatic pressure. Here l is the maximum principal stress and ,3 is the least principal stress. Commonly, the symbol )t is used to express the magnitude of Pt in the crust relative to Sv, the vertical stress or lithostatic load. Therefore, by definition, )t = Pf/Sv. For an average fluid density of i and average rock density of 2.6, )t = 0.38 where hydrostatic Pf prevails. A somewhat greater stress difference would be required to cause shear failure on an otherwise open fracture that had regained cohesive strength as a result of cementation by vein formation. For example, for a fracture that has cohesive strength given by point N in Figure 1, the dashed line from point N through point A' shows the stress differences required to renew shear failure on that fracture as confining pressure increases with depth, provided h = 0.38.

In contrast to the conditions for brittle failure, the stress difference required to initiate plastic deformation is highly dependent on temperature, strain rate, and rock type (material constants), and it is little affected by confining pressure. However, compared to dry rock, the presence of water allows plastic behavior at lower temperatures (Carter and Tsenn, 1986). The general law governing steady-state plastic or ductile

Stress difference (o-o3) N

1 gg edneePr ldy s0. o g!fsi iaenn 1 fu if ;tjes u r e \ 0.38

' "",, Decreasing strain rate \ ' , and/or more silicic

\ A 'A ' Brittle = 0 38 .' Plytic Brittle = 0.7 cmasing stroh rme l / or more mc C ............. Brittle = 1.0

...... _7'-- ..... t - Lo tat' fim cohesive streng

0c= 1)

Fig. 1. Stress difference (rrl-r a) versus depth, showing schematically conditions for brittle shear failure along a fracture in the upper crust at selected ratios of hydrostatic pressure to the lithostatic load (k values) and plastic deformation at deeper levels. See text for discussion.

HYDROTHERMAL PROCESSES IN THE MAGMATIC-EPITHERMAL ENVIRONMENT 1195

deformation can be expressed approximately by the following equation:

E =A (cr-3) exp(-Q/RT), (1) where E is the strain rate, R is the gas constant, T is absolute temperature, (cr- 3) is the stress difference acting on the material, and A, Q, and n are material coefficients that change with rock type (Turcotte and Schubert, 1982). For the normal situation in which temperature increases in direct proportion to depth, the stress difference required to cause plastic deformation at a given strain rate decreases exponentially with depth (curve A'ABC in Fig. 1). For a given externally imposed strain rate, shales and rocks rich in quartz exhibit plastic behavior at lower temperatures (less deep in the crust) than more mafic rocks, such as basalts and gabbros. Even more important, according to equation 1, an increase in strain rate allows brittle behavior to occur at higher temperatures (greater depths) in the crust. An increase in strain rate also can result in an increase in temperature by frictional heating (eq. 1).

In the brittle region, increasing the stress difference causes shear failure to occur before the stress difference becomes great enough to cause plastic deformation at a given externally imposed strain rate. In contrast, in the plastic region, at a given externally imposed strain rate, increasing the stress difference causes plastic deformation before the stress difference becomes great enough to cause shear failure. It follows that the brittle to plastic transition occurs where the brittle and plastic deformation curves intersect. In Figure 1, point A marks the depth of the brittle to plastic transition at a speci- fied strain rate when hydrostatic pressure prevails, or k = 0.38. In reality, because of rock inhomogeneities the brittle to plastic transition may occur within a depth interval, rather than at a specific depth. Also, an increase in Pf, so that k is >0.38, will move the brittle-plastic transition to a greater depth along the plastic deformation curve (curve ABC in Fig. 1).

The line from the origin (at zero depth) through point B in Figure 1 shows the stress difference required to activate fault movement (brittle shear failure) when k: 0.7. In this situation the depth of the brittle-plastic transition is moved downward from point A to point B, and it occurs at a correspondingly lower stress difference. Theoretically, when Pf = Sv shear failure can occur on open fractures having no cohesive strength in response to vanishingly small stress differences. In this situation the brittle-plastic boundary would occur very deep in the crust at a very high temperature. However, as temperature increases with depth, open fractures quickly become chemically healed or cemented by vein deposition and sintering of adjacent grains, so that a finite stress difference is required to activate shear failure of faults even where Pf = Sv (for example, the dashed vertical line passing through point C in Fig. 1; Fournier, 1996). Brittle behavior of normally plastic rock at high strain rates

Consider a shallow plutonic environment where a magmatic body is crystallizing at a depth of about 3.5 to 4 km and where an overlying hydrothermal system at hydrostatic pressure circulates from the land surface to a depth of about 3 kin. With temperature gradient information provided by presently active hydrothermal systems in mind, Figure 2 was constructed

so that the maximum temperature attained by the circulating hydrothermal system is in the range 360 to 375C, the average temperature gradient from the surface to a depth of 3 km is 125C km -, and the thermal gradient below 3 km, where heat is transferred mainly by conduction from the magmatic heat source to the base of the hydrothermal system, is 500C km -. A coefficient of friction of 0.75 was used for shear failure in the brittle region (Byerice, 1978; Sibson, 1985), and the rheologic properties of wet quartz diorite (Carter and Tsenn, 1986) were used for the plastic region.

Given the above thermal profile and material properties, Figure 2 shows conditions for brittle and plastic behavior when k = 0.38, 0.6, and 1.0 at selected strain rates ranging from 10 -14 s -1 to 10 -6 s -1' In regions of active crustal deformation strain rates generally range from about 10 -14 S -1 to 10 -15 s - (Pfiffner and Ramsay, 1982), and strain rates of 10 -2 s - to 10 - s - may be attained for very short times during periods of active faulting (Sibson, 1982). In volcanic environments, very high rates of deformation have been observed during dome eraplacement and during collapse accompanying eruptions (Dzurisin et al., 1983). In Figure 2, the brittle to plastic transition is at a depth of about 3 km and a temperature of about 360C (point A) when k = 0.38, the strain rate is 10 -14 s -, and optimally oriented open fractures have no cohesive strength. Increasing fluid pressure to greater than hydrostatic moves the brittle to plastic transition point along curve ABC to higher temperatures (greater depths) and lower stress differences at the transition point, as discussed above. Note that the calculated brittle to plastic transitions in Figure 2 would have been at slightly shallower depths if the rheologic properties of wet rhyolite had been assumed and at significantly greater depths if the rheologic properties of basalt had been assumed.

Figure 2 illustrates that a temporary increase in the strain rate to > 10 -14 s -1 in initially plastic rock will allow brittle behavior at relatively low stress differences at very high temperatures. For example, for initially plastic rock that has 200 bars cohesive strength and k = 1, temporarily increasing the strain rate to 10 -s s - moves the brittle-plastic transition to point G at a depth of about 4 km, where the temperature is about 870C. Thus, brittle behavior is moved into the temperature domain of melt or partial melt.

The calculated conditions for shear failure at high strain rates shown in Figure 2 are very approximate because there would be changes in the material constants at very high temperatures where partial melting occurs. However, the general conclusion that shear failure of very hot material can occur at high strain rates and small stress differences is valid. Some examples of shear failure that apparently occurred in partly molten magmatic bodies are given below. Examples of brittle behavior of magrrms

Fractures, veins, and dikes in many plutonic bodies appear to have formed and to have been subsequently offset before surrounding magmatic material completely solidified. This implies that partly molten material behaved in a brittle manner, allowing shear failure of crystal-rich magma to occur. Ex- amples of shear failure within a partly molten granitic body deep in the crust are given in a paper by R. B. Fournier (1968). That paper describes textural variations and structural features exhibited by mafic schlieren and by alaskitic, aplitic,


Stress difference (61-63), Bars 0 1000 2000 3000 4000 5000

= 0.38 (no cohesive strength) oo_ o

. Strain rte2.1ol: s 1. Brittle (point n) 30 400- .. ''' Plastic -- . km o' s'

'

l.' :i ..

600- '{ .' ' ' * -. 3.5 km .. I . . ..10- s - I

800- .' I i.. , T 10-s ' l: I . Brittle (oJnt G)

' ............. 4.0 lOOO

i< :" ) = 1.0 (with 200 bars cohesive strength)

* ' ' ' ' I .... I ' ' " ' I ' ' ' ' I

Fig. 2. Stress difference versus temperature, showing effect of strain rate, value of h, and cohesive strength on the brittle to plastic transition. See text for discussion.

and pegmatitic dikes enclosed in the Half Dome quartz monzonite of the Sierra Nevada batholith. All of the dike material appears to have been derived by segregation during differen- tial flow of the partly crystalline pluton at a depth of about 7 to 10 km. Many of the alaskitic dikes are associated with mafic schlieren, and these dikes usually have undulating walls that are not closely matching. The alaskitic dikes are interpreted by R. B. Fournier (1968) to have formed by movement of in- terstitial melt into shear zones (perhaps as a result of filter pressing) without the simultaneous evolution of a separate water-rich phase. This occurred while the surrounding partly molten quart. z monzonite was unsaturated with respect to water.

In contrast to the alaskitic dikes, the slightly later aplitic and pegmatitic dikes are interpreted to have formed with the simultaneous separation of a water-rich phase during injection by filter pressing into slightly underpressured shear zones. These dikes probably were not emplaced as a result of a simple hydraulic opening of parallel fractures perpendicular to the least principal stress, because some of the alaskitic and most of the aplitic and pegmatitic dikes exhibit a rectilinear pattern that implies shear failure at the time of their formation. In addition, the aplitic and pegmatitic dikes commonly crosscut and offset each other, again suggesting brittle failure at the time of formation. In some composite dikes of aplite and pegmatite, the upper contacts of the pegmatite zones

with the surrounding aplite are undulating, indicating that the aplite was still plastic at the time when the pegmatite segre- gated. R. B. Fournier (1968) concluded that the aplitic textures were the result of rapid crystallization of magma that had become supercooled at the time of separation and movement of water into slightly underpressured shear zones (Pf temporarily < Sv). The textures of the aplitic dikes are similar to groundmass textures exhibited by most porphyry Cu deposits.

Turning to the shallower porphyry Cu environment, the oldest "A" quartz veins of Gustafson and Hunt (1975) formed without parallel walls and generally before the rock was able to sustain continuous brittle fracture. They also show an illus- tration (their fig. 16) of an alkali feldspar seam cutting plagioclase phenocrysts but with no trace of the seam cutting inter- vening fine-grained, aplitic-textured quartz and alkali feldspar. Similar features are present in the porphyry Cu deposit at Ely, Nevada, where this texture was interpreted to show that fracturing and K feldspar replacement of plagioclase took place before the viscous, glassy groundmass had crystallized, at the time of a pressure quench of the porphyry that supercooled the residual magma by tens of degrees (Fournier, 1967). This type of brittle behavior probably occurred during very short-lived episodes of very high strain rate, as discussed above, with the system returning to a plastic condition almost immediately thereafter because differen- tial stress was diminished by the shear movement. However,


even though such fractures may be quickly closed, while they last they may be important avenues for rapid movement of magmatic fluids from deeper to shallower levels through normally plastic rock and then into the normally brittle domain.

Development and Significance of Plastic Domains around Igneous Intrusive Bodies at Shallow Depths

In the early stages of development of a volcanic complex, upward-intruding magmas encounter relatively cold, brittle country rocks in which hydrostatic Pf would probably prevail to depths of at least 6 to 8 km. In this situation, thermal gradients at the margins of newly intruded plutonic bodies that come to rest quickly within relatively cold crustal material would be very steep (Cathies, 1977; Norton, 1982). There- fore, the rinds or shells of plastic rock that surround the still molten portions of the plutonic bodies, separating them from brittle rock in which hydrostatic pressure prevails, would be correspondingly thin. In this situation, there would be little room for the accumulation of exsolved magmatic fluids at less than magmatic temperatures in plastic rock adjacent to the intrusion; the crystallizing and cooling magmatic body would have to serve as the receptacle for its own exsolved fluids, or lose them into the surrounding, hydrostatically pressured domain. A rapid loss of magmatic fluids across the relatively thin shell of plastic rock around the intrusion, initiated relatively easily by any of the mechanisms discussed later, would probably result in a decrease in Pf at the interface between that exsolved fluid and its water-saturated parent magma. A decrease in confining P would, in turn, result in an increase in volatile evolution from that water-saturated magma (Burn- ham, 1967, 1979).

In the later stages of development of a volcanic complex, upward-intruding magmas encounter rocks that have been heated to successively higher temperatures by previous igneous injections. Eventually, a considerable volume of rock becomes hot enough to behave plastically at depths as shallow as 1.5 kin. In addition, within a large volume of plastic rock the least principal stress will generally be the lithostatic load, Sv. When and where a3 __ Sv in plastic rock, that hot rock can act as a large receptacle for the temporary storage and cooling of fluids exsolved from newly intruded batches of crystallizing magmas (discussed in the next section). These fluids, which may accumulate over a long interval of time, become available to act as mineralizing agents in the formation of epo ithermal mineral deposits. It may be more than a coincidence that both the development of a sizable body of plastic rock at a relatively shallow depth beneath a volcanic complex and the initiation of epithermal ore deposition commonly take place only after a long interval of preceding barren volcanic and plutonic activity in a given region (e.g., deposits described in Silberman, 1983; Heald et al., 1987).

Accumulation of Magmatic Hydrothermal Fluid in, and Movement through, Plastic Rock

Accumulation in horizontal lenses when and where a s: Sv It is commonly assumed that the aqueous fluids that exsolve

from crystallizing magmas become concentrated beneath a carapace at the top of newly crystallized subsolidus rock, and that fracturing and brecciation occur as a result of increasing

fluid pressure, as in illustrations in Burnham (1979, 1985). In addition, Shinohara et al. (1995) present a plausible model for transporting volatile components from a large volume of magma to the top of the magma chamber, where metal-rich aqueous fluids accumulate. However, it seems unlikely that individual cavities or fractures beneath a carapace in plastic rock at the top of a crystallizing igneous intrusion will main- tain much vertical extent for the long periods of time required for significant quantities of fluid to exsolve by fractional crystallization. This is because deformation in response to buoy- ancy in large bodies of plastic rock results in the lithostatic load becoming the least principal stress (a3: Sv), and fluid- filled fractures or cavities are deformed by plastic flow into relatively thin, horizontal overlapping lenses in which P: Sv (Bailey, 1990).

Consider a cavity (or interconnected network of fractures) filled with aqueous fluid within a rock that behaves plastically, with a3: Sv. At very high temperatures (>600-700C) plastic flow, in response to a very small stress difference (Fig. 2), will force P at the bottom of the cavity (or fracture network) to approach the rock pressure, Sv, at that particular depth. This pressure is transmitted upward hydraulically through the fluid and P within the cavity at shallower depths becomes > Sv according to the following relationship:

Pf = Sv + gh(pr-pr), (2) where g is the gravitational constant, h is the effective vertical dimension of the hydraulically interconnected cavity, pr is the average density of the plastic rock surrounding the cavity, and pr is the density of the fluid in the cavity. Thus, cavities or interconnected networks of open fractures with considerable vertical extent in rock that behaves plastically are mechanically unstable. Near their tops they will tend to de- form by spreading out laterally, while the lower part of the cavity is squeezed shut. Because the stress differences required to initiate plastic flow decrease with increasing temperature while the rate of squeezing shut increases exponentially (eq. I and Figs. i and 2), fluid-filled cavities with considerable vertical extent are likely to persist for only very short periods of time in plastic rock immediately above a crystallizing magma, where temperatures exceed 600C. In this situation an exsolved fluid makes room for itself by lift- ing the overlying rock, and Pr is fixed at approximately lithostatic pressure independent of the degree of fractional crystallization of the magma. At 400 to 500C, hydraulically interconnected cavities with somewhat greater vertical extent may persist for longer periods of time than at 600C because greater stress differences are required to initiate plastic flow at lower temperatures.

The relatively flat-lying "B" veins in the E1 Salvador, Chile, porphyry Cu deposit, described by Gustafson and Hunt (1975), are examples of cavities that probably formed when exsolved magmatic fluid at lithostatic pressure began to accumulate in plastic rock. They are prominent structures that formed after most "A" veins (discussed previously). The "B" veins exhibit coarse-grained quartz crystals that grow inward from the walls, and fluid inclusion filling temperatures are as high as >600C. In addition, Fyfe et al. (1978) noted that horizontal lenslike veins or water sills are common in epigenetic hydrothermal deposits beneath shale or other impervious


materials and concluded that they form where lithostatic Pt is attained. There also is stable isotope evidence that magmatic fluids alecouple from their magma earapace or source and then reequilibrate in lower temperature plastic rock at lithostatic pressure before subsequent decompression events that produce epithermal orebodies (Rye, 1993; Deenet al., 1994).

When measuring and interpreting the attitudes of veins that are likely to have formed in rocks when temperatures were >400C, one should keep in mind that initially flat-lying veins may take on moderate to steep dips as a result of later deformation.

Rapid upward movement of fluid through plastic rock when rr s < Sv

In the above discussion emphasis was placed on the accumulation of fluids at lithostatic pressure and near magmatic temperatures in horizontal lenses near the tops of crystallizing plutons when rr 3 = Sv. In this situation, rapid upward movement would probably occur only during a short-lived episode of high-angle shear failure, when the rock behaved in a brittle manner. In contrast, when and where rr 3 < Sv in plastic rock, and that rock is cool enough to behave elastically and undergo hydraulically induced failure, an exsolving magmatic fluid will make room for itself by first creating and then en- larging hydraulic fractures approximately perpendicular to rr when Pt -- < Sv. As soon as the vertical dimension of such a fracture attains a particular length, depending on the fracture toughness of the surrounding rock, fluid in the hydraulic fracture starts to move upward by opening at its top leading edge while simultaneously closing at its lower edge (Secor and Pollard, 1975, Pollard, 1976). Thus, vertical lenslike bodies of fluid are transported upward relatively quickly and effi- ciently until they encounter a different stress field (perhaps where rr = Sv in the overlying plastic domain, or where open shear fractures exist in the brittle domain). An additional re- stfiction for upward flow by this mechanism is that the least principal stress gradient with depth must be greater than the fluid pressure gradient. Where the reverse is true, hydraulic fracturing may move fluid downward rather than upward (Pollard, 1976).

In natural systems, crystallization and rapid cooling inward from the walls of dikes or small plutons that have been intruded to shallow depths can result in the development of significant tensile stresses that act normal to those walls (rr acting at about fight angles to Sv). Thus, the condition rr 3 < Sv may develop at >400C in and adjacent to such structures and apparently persist long enough for high-angle hydraulic fracturing to occur. Two examples of ore deposits that are thought to have had relatively rapid upward movement of magmatic fluids through plastic rock in and adjacent to dikes at >400C and lithostatic Pt are the Big Gossan Cu-Au skam deposit, Ertsberg district, Irian Jaya (Meinert et al., 1997), and the Far Southeast porphyry Cu- Au deposit in the Philippines (Hedenquist et al., 1998).

Self-Sealing at the Brittle-Plastic Interface Given that magmatic fluids tend to accumulate in plastic

rock at lithostatic pressure where rr 3 = Sv, and that fluids at hydrostatic pressure circulate through brittle rock where permeability is maintained by recurrent seismic activity, the nature of the interface between these two different hydrologic regimes

is of considerable theoretical interest and practical importance in the development of models for ore deposition. In the brittle regime, when dilute solutions in contact with quartz at hydrostatic pressure are slowly heated at pressures ranging from about 34 to 900 bars (maintaining saturation with respect to the solubility of quartz), a point is reached at which there is a change from dissolution of quartz to precipitation of quartz with further heating (Kennedy, 1950). Figure 3 shows calculated solubilities of quartz at the vapor pressure of the solution at selected isobars. It also shows the region of retrograde solubility where precipitation occurs upon heating at constant Pt (stippled area). Where the circulation of a relatively dilute fluid at hydrostatic pressure extends downward to near the top of a cooling pluton, at a depth of about 3 to 4 km (300-400 bars in Fig. 3), the onset of quartz deposition with heating would occur at about 370 to 390C. Increasing salinity and increasing pressure (at greater depths of circulation) move the point of maximum quartz solubility to > 400C (Fig. 4a and b). However, at >400C quartz diorite starts to become quasiplastic (discussed previously) and this limits the time that fractures are likely to remain open for fluid flow.

Coming at the brittle-plastic transition zone from the other direction (fluid moving from the higher Pt region in plastic rock into the lower Pt region in brittle rock), there is a large potential for the precipitation of silica from dilute and highly saline fluids with decreasing pressure at >400C (Fig. 4a and b). Decompression is also likely to result in massive evapora- tive boiling of brine (discussed later) that causes additional supersaturation with respect to the solubility of quartz.

In the above discussion emphasis was placed on quartz deposition as a major factor in the formation of a self-sealed zone because quartz veins commonly occur in hydrothermal systems. Other commonly observed vein minerals, including carbonates, sulfates, sulfides, oxides, and other silicates, also might play major roles in the development and/or reestab- lishment of a self-sealed zone.

Mechanisms for Breaching the Self-Sealed Zone and Discharge of >400C Fluid into Cooler Rock

Under normal circumstances, without a major perturbing event, the leakage of fluids (probably mostly gas) from plastic rock into an overlying hydrostatically pressured hydrothermal system occurs very slowly. However, in a subvolcanic hydrothermal-magmatic system there is likely to be a precarious balance between processes that create permeability and those that diminish permeability. Episodically, minor to major breaches of the self-sealed zone occur in response to various processes. One such process is continued degassing of crystallizing magma and accumulation of the evolved fluid beneath a carapace until that carapace becomes stretched sufficiently to rupture by tensile failure, as postulated by many investigators (e.g., Norton and Cathles, 1973; Phillips, 1973; Burnham, 1979, 1985). A variation on this theme of "accumulation of fluid near the top of a still partly molten pluton leading to rupture of overlying rock" is exsolution of volatiles at depth and transport to the top of the pluton by convection of the magma (Shinohara et al., 1995).

Another, and probably important, mechanism for initiating breaches of a self-sealed zone is upward injection of a new pulse of magma from depth. This can increase the strain rate


5OOO

4000

$000

2000

IOOO

I I I I I I I !

Quartz solubility in water

/

*/ / i o/ / i el oOl l // '/oo./

/// i/ 1o' / / / /

/ / / / / I / / / / /

Fig. 3. Calculated solubilities of quartz in water as a function of temperature along isobars ranging from 200 to 1,000 bars. The stippled pattern shows a region of retrograde solubility in which the solubility of quartz decreases with increasing temperature at constant pressure. From Fournier (1985).

within the overlying rock to such a degree that the brittle to plastic transition is temporarily moved to a deeper and hotter condition, as discussed above. Because fluids in the initially plastic rock would be at lithostatic Pf ( = 1), the change from plastic to brittle behavior with increasing strain rate would

result in breaching of the self-sealed zone by shear failure in response to a small stress difference (just sufficient to overcome the cohesive strength of the rock). A new pulse of upward magma injection also would be a source of additional volatiles upon crystallization. In addition, it would

a I I [

40

300 Bars

$o

Temperature, C

b i l Quartz Solubility l

40 500 Bars zo

,_

,o 0 I00 200 :500 400 500 600

Temperature, C

Fig. 4. Comparison of calculated solubilities of quartz in water and in NaCI solutions at (a) 300 bars pressure, and (b) 500 bars pressure. From Fournier (1983).


heat previously existing brine trapped in horizontal lenses (probably inducing boiling), and the resulting rapid expansion of that fluid would cause an additional increase in the strain rate, possibly leading to failure of the overlying rock.

Large seismic events that affect permeability and rates of flow within an overlying hydrostatically pressured hydrothermal system may also lead to a breaching of the self-sealed zone if the increase in flow rate results in extraction of heat at the base of the circulating system at a rate more rapid than heat can be supplied by conduction through plastic rock from below. This would lead to a progressive downward decrease in temperature and simultaneous expansion of the permeable region as a result of volumetric contraction and tensile crack- ing (Lister 1974; Norton and Knapp, 1977; Norton and Knight, 1977; Carrigan, 1986). Thus, as rock temperatures decrease, the depth at which brittle failure can occur is low- ered into previously plastic rock where fluids had been trapped at lithostatic pressure. The high Pt in the previously plastic, but now brittle, rock facilitates shear failure of the self-sealed zone in response to a relatively small stress difference (Fig. 2).

Foundering of a large slab of roof rock (detached by horizontal hydraulic movement of fluid) in such a manner that it slides down one side of a magma-filled pipe, causing rapid upward flow and vesiculation of water-saturated magma on the other side of the pipe, might result in failure of the overlying rock. This is much like the turnover and foundering of solidified crust on a lava lake. The convection model of Shi- nohara et al. (1995) has elements in common with this mechanism, except here convective flow and rapid accumulation and expansion of exsolved fluid at the top of the system is induced by downward movement of a relatively dense body of solid rock.

Characteristics of Fluids in the Magmatic-Epithermal Environment

Salinity variations and fluid inclusions in plastic rock at >400C

Phase relations in the systems NaC1-HgO and NaC1-KC1- HgO over widely ranging temperatures and pressures provide a good first approximation of how salinity is likely to vary in hydrothermal fluids derived from crystallizing magmas (Sourirajan and Kennedy, 1962; Nash, 1976; Cunningham, 1978; Eastoe, 1978; Bloom, 1981; Cline and Bodnar, 1994). Figure 5a, modified from figure 55.4 in Fournier (1987), shows phase relations in the system NaC1-HgO when Pt is controlled by lithostatic pressure (assumed average specific gravity of overlying rock = 2.5 gm/cm3). The brittle-plastic boundary is shown at about 390C (Fournier, 1991) but could be at a higher or lower temperature for the various reasons described previously. The dot-dashed lines show boiling- point curves for aqueous solutions containing 30, 50, and 70 vet percent NaC1. The curve labeled "A" is the 3-phase (gas + liquid + halite) boundary for the simple system NaC1-HgO. It separates a field of gas + halite (G + S) from a field of gas + liquid (G + L). The curve labeled "B" shows the approximate position of the gas + liquid + solid salt phase boundary when both dissolved NaC1 and KC1 are present and the NaJK ratio is fixed by the presence of coexisting albite and K feldspar

(Orville, 1963; Hemley, 1967). Between curves A and B is a field of gas + liquid + solid salt (mainly halite).

The short dashed curve passing through point C in Figure 5a shows the boiling point curve for a 10 vet percent NaC1 solution at pressures less than about 520 bars, and the dew point (or condensation point) curve for a gas containing 10 vet percent dissolved NaC1 at pressures greater than about 520 bars. These two curves meet at the critical point for a 10 vet percent NaC1 solution (point C in Fig. 5a). To the left of the condensation point curve, water with 10 vet percent dissolved NaC1 commonly would be considered a supercritical fluid. To the right of the condensation point curve, water containing 10 vet percent dissolved NaC1 is not stable; it will dissociate into hypersaline brine and a less saline gas or vapor. In a natural system the positions of the boiling point and condensation point curves may be shifted upward or downward by addition of other dissolved constituents (e.g., downward by addition of CaClg and noncondensable gases, such as COg).

In the remainder of this presentation, HgO-rich vapor that is in equilibrium with brine at pressures > 221 bars (the critical pressure of water ) will be referred to as "steam." The quotation marks are used to emphasize that this "steam" has a relatively high density compared to steam at 100C and at- toospheric pressure, and that it is capable of carrying significant concentrations of dissolved constituents. Quotation marks will not be used for steam at pressures less than 221 bars.

Isochores of the solubility of NaC1 in "steam" are shown in Figure 5b. Note that at pressures greater than about 400 bars and temperatures less than about 800C, the solubility of NaC1 in "steam" decreases as it expands with increasing temperature and with decreasing pressure.

The five circles labeled "E" through "H" in Figure 5 indicate hypersaline brine (Fig. 5a) and coexisting "steam" (Fig. 5b) that have exsolved at about 750 to 800C from crystallizing magma at depths of 1, 2, 3, and 6 km, respectively. At a depth of 6 km (circle H) about 20 wt percent NaC1 can dissolve in the "steam," which is likely to be greater than the salinity of aqueous fluids exsolving during the crystallization of most magmas (Hedenquist et al. 1998). However, as pressure decreases the solubility of NaC1 in "steam" decreases (Fig. 5b). Accordingly, at a depth less than about 5 km an ex- soMng magmatic fluid will dissociate immediately to highly saline brine and "steam." At a depth of 3 km and 750 to 800C, the coexisting brine and "steam" will contain about 65 and 1.25 vet percent NaC1, respectively (circle G in Fig. 5a and b). At a depth of i km there is so much salt dissolved in the water that it is more realistic to consider the liquid to be an NaC1 melt that contains about 10 to 15 vet percent dissolved water (Fig. 5a, circle E). The coexisting "steam" or vapor contains less than about 0.05 vet percent NaC1 (Fig. 5b, circle E).

Noncondensable gases, such as COg, HgS, and SOg, prefer- entially partition into the "steam" phase that evolves with brine at magmatic temperatures. The SOg/HgS ratio in this "steam," and subsequent changes in this ratio upon cooling, are of considerable importance in regard to epithermal mineralization. With decreasing pressure the SOg/HgS ratio increases in the vapor phase that exsolves from a crystallizing magma of given composition (Carroll and Webster, 1994).


2OO

400

600

800

1000

1200

1400

1600

Temperature, C 200 400 600 800 1000 0

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,

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200

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600

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Temperature, C

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%%% %%

%

10 wt% NaCI ' in gas ("steam")

() 6 km- I , , , i , , , i , , , i , , ,

Fig. 5. Temperature-depth diagram showing phase relations in the system NaC1-HO at lithostatic Pc as an analog for aqueous, chloride-rich fluids exsolved from crystallizing magma. G = gas, L = liquid, and S = solid salt: (a). Dot-dashed lines are contours of constant wt percent NaC1 dissolved in brine; short dashed line shows the boiling point curve for a 10 wt percent NaC1 solution at pressures and temperatures below its critical point (C) and at pressures and temperatures above point C it shows the dew point curve (or condensation point curve) for "steam" containing 10 wt percent dissolved NaC1; curve A shows the three-phase boundary, G + L + S, for the system NaC1-HO; curve B shows the three-phase boundary, G + L + S, for the system NaC1-KC1-HO, with Na/K in solution fixed by equilibration with albite and K feldspar at the indicated temperatures; the triple dot-dashed line shows the approximate temperature of the brittle-plastic boundary when the strain rate is 1044 sec 4 (modified from fig. 55.4, Fournier, 1987)). (b). Dashed lines show dew point curves (or condensation point curves) for "steam" containing indicated values of wt percent dissolved NaC1; curve A shows the three-phase boundary, G + L + S, for the system NaC1-HO (modified from fig. 55.6, Fournier, 1987). See text for discussion.

Within that vapor phase, HS reacts with water at high temperatures and low pressures, producing SO and H (Gerlach, 1993; Rye, 1993). Giggenbach (1997) also shows a figure (fig. 15.7) in which HS is converted to SO at higher temperatures and lower pressures, with the temperature of conversion highly dependent on the oxidation state of the system. Thus, magmas with identical initial oxidation states and dissolved sulfur concentrations will yield relatively HS-rich aqueous fluids when crystallization takes place deeper in the system (e.g., circle H in Fig. 5a) and relatively SO-rich fluids when crystallization takes place at shallower depths (e.g., circle E, Fig. 5a). During the subsequent cooling of these magmatic fluids the HS/SO ratio is affected by the degree of fluid-rock interaction that occurs, staying about the same above 400C if there is no reaction with wall rocks, or becoming HS dominant if there is reaction with the wall rocks that buffer oxygen fugacity (Ohmoto and Rye, 1979; Rye, 1993). Below about 400C, SO. reacts with liquid water yielding HSO4 and H.S (Holland, 1965).

If the "steam" and hypersaline brine remain in contact during cooling in place (perhaps in isolated lenticular cavities), some or all of the "steam" will dissolve back into the brine, lowering the brine salinity as cooling progresses. At depths greater than about 2 to 2.5 km essentially all of the brine and "steam" (except the small quantities trapped in fluid inclusions) will reconstitute into a single fluid phase having nearly

the initial salinity of the magmatic fluid. However, the fluid inclusions in the rock would record a systematic decrease in salinity with decreasing temperature. Vapor-rich inclusions would probably be present in the rock, but these would have formed at the time of the initial dissociation of exsolving magmatic fluid to highly saline brine and "steam." After cooling to temperatures in the range 400 to 450C, the fluid would be either a moderately saline brine or supercritical gas, depending on the depth and the initial salinity of the exsolved magmatic fluid (e.g., circles K-N, Fig. 5a). The HS/SO ratio for the gas in the cooled fluid would be considerably higher than that for the gas in the initial fluid at near magmatic temperatures (Ohmoto and Rye, 1979). The total concentration of sulfur species could range from about the same as to less than that in the initially exsolved magmatic fluid (because there is no loss of "steam" containing noncondensable gases) as a result of precipitation of sulfides during cooling.

At depths less than about 1.5 km, cooling from about 800 to 450C would probably result in precipitation of salt. For example, at a depth of 1 km, cooling of the brine + "steam" mixture represented by circle E (Fig. 5a) to the condition shown by circle K will move that mixture into the field of gas ("steam") + liquid + solid salt in the system NaC1-KC1-HO at temperatures of about 700 to 450C. Additional cooling to 450 to 400C, with no loss of "steam," would move the fluid back into the field of gas + liquid (containing about 30-50 wt


% NaC1 equiv). Brine-filled fluid inclusions that formed during cooling would have widely ranging salinities and could contain halite and sylvite daughter crystals. In some inclusions these daughter minerals might disappear upon heating after the vapor bubble disappeared.

If the "steam" and brine physically separate before significant cooling occurs (perhaps by "steam" escaping upward from the system while the brine remaines behind), the salinities of brines cooled to about 450 to 400C at depths greater than about 1.5 km would be the same as those of the corresponding brines at 700 to 800C. In contrast, at depths less than about 1.5 km, if "steam" escaped early there would be very little brine present in the rock after cooling to 450 to 400C, and this brine would have a salinity in the range 30 to 50 wt percent NaC1 equiv. At all depths the residual brine would be greatly depleted in sulfur relative to the initially exsolved magmatic fluid because most H.S and SO. would be carried away with the "steam" phase.

Results similar to those described above for separation of "steam" upward would be obtained if the brine moved hori- zontally while the "steam" remained behind. In both of the above scenarios ("steam" escaping upward, or brine moving laterally away from "steam"), the fluid inclusions that formed during cooling would record the lowering of temperature but exhibit closely clustering salinities. However, early-formed vapor-rich inclusions would probably be found with early- formed brine inclusions where most of the brine cooled in place after separation and escape of steam. In contrast, vapor- rich fluid inclusions would not be found with highly saline brine fluid inclusions where there had been lateral movement of brine away from its initially coexisting "steam" phase. In the latter situation, the lack of vapor-rich inclusions coexisting with highly saline brine inclusions might lead to a false conclusion that the initial magmatic fluid had a much higher salinity than it had in reality.

The fluids in P-T space represented by circles K to N in Figure 5a may have cooled in plastic rock at constant pressure, or they may have followed a cooling path that involved upward movement through plastic rock. Normally, upward movement of brine (or moderately saline supercritical fluid rich in gaseous sulfur species) through plastic rock would be very slow. An exception, noted previously, would be where contraction within cooling dikes allows hydraulic fracturing to occur along vertical channels, such as appears to be the case at the Big Gossan Cu-Au skarn deposit, Ertsberg district, Irian Jaya (Meinert et al., 1997), and the Far Southeast porphyry Cu-Au deposit in the Philippines (Hedenquist et al., 1998). These deposits give permissive support to thepossibil- ity that rapid upward hydraulic movement of fluid within cooling dike systems may be important. Effects of decompression from lithostatic to hydrostatic pressure

The effect of a sudden decompression from lithostatic to hydrostatic pressure conditions that would occur when fluids are discharged rapidly from plastic rock into the brittle domain can be seen by comparing Figures 5a and 6. Note that in Figure 6 the depth-hydrostatic pressure relationships are shown for a water table at or very near the earth's surface. In reality, the water table (potentiometric surface) may lie above

0

100

200

300

400

500

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700

800

Temperature, C

200 400 600 800 1000

.:' /-

i',X

h \' 70 wt% i C'"\ \, \ 5km - 30 wt%

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0 wt . . , , ,

Fig. 6. Temperature-depth diagram showing phase relations at hydrostatic Pf in the system NaC1-H20. Notation is the same as for Figure 5a. See text for discussion.

the earth's surface (artesian springs) or, more likely in a volcanic environment, tens to hundreds of meters below the earth's surface. Also, in large vapor-dominated systems, such as The Geysers, California, there is little change in Pf in the vapor-dominated region from the top to the bottom of the reservoir over a depth range of i to 2 km (White et al., 1971; Walters et al., 1992). For this system, the effective water table for an underlying brine would be at least i to 2 km below the earth's surface, and the curves shown in Figure 6 would be deeper by about i to 2 km at corresponding temperatures.

At near-magmatic temperatures, decompression from lithostatic to hydrostatic Pfwill cause brines to become much more saline with the simultaneous evolution of much "steam." Flu- ids at or near magmatic temperatures may experience sudden decompressions in Pf because the distance from the top of a magmatic intrusion (where temperatures are 700-800C) to the contact with the brittle (hydrostatic) domain (where temperatures are about 370-400C) may be as short as a few meters. An example of the probable discharge of magmatic fluid across a very thin brittle-plastic interface into a gas + solid salt environment is the mineralization at Alunite Ridge, near Marysvale, Utah (Cunningham et al., 1984; Rye, 1993). At this locality banded veins of coarse-grained, nearly pure alunite were deposited as open-space fillings in extensional fractures as much as 20 m wide and100 m in vertical extent from a low density, SO.-rich magmatic fluid (Rye, 1993). The magmatic source for these fluids apparently was at a relatively shallow depth, such as circle E in Figures 5 and 6. Adiabatic cooling upon decompression brought the fluids into the field of gas (superheated steam) + solid salt, and this gas likely vented to atmospheric pressure as a superheated fumarole, such as are common in volcanic environments.

Decompression from lithostatic to near-hydrostatic pressure will have little effect on brine salinities in plastic rocks


where temperatures are in the range 400 to 450C, and at depths greater than about 4 km (e.g., circles M and N in Figs. 5a and 6) because boiling is not induced. In contrast, at 2 km and 400 to 450C a sudden decrease in Pf from lithostatic (circle L in Fig. 5a) to hydrostatic (circle L in Fig. 6) will probably bring fluid into the field of gas + brine + solid salt. If the initial salinity of the fluid at lithostatic pressure were moderate (perhaps 5-10 wt %), a very large volume of steam would form immediately upon decompression and a frothlike mixture of highly saline brine and "steam" might bubble up and out into the brittle domain.

In general, decompression processes that induce the massive formation of "steam" would probably be accompanied by mineral deposition in veins and cavities, and the formation of primary and secondary fluid inclusions. The fluid inclusions in single specimens would be expected to exhibit widely ranging salinities in association with abundant vapor-rich inclusions. Evidence of entering the field of gas + liquid + solid would be the presence of fluid inclusions that, upon heating, exhibit the disappearance of their vapor bubbles before the disappearance of salt daughter minerals. Fluid inclusion evidence for actually boiling dry would be difficult to find because many brine-bearing fluid inclusions probably would be present, formed both before the pressure quench, and during the subsequent process of boiling to eventual dryness. Within portions of a late-stage porphyry at Ely, differences in composition between K feldspar phenocrysts and K feldspar in the surrounding aplitic-textured groundmass indicate that that particular intrusive mass probably boiled dry during a pressure quench and stayed dry thereafter until the system had cooled to near present-day temperatures (Fournier, 1967).

The movement of brine into the field of gas (superheated "steam") + solid salt has important implications regarding the concentrations of HC1 that may be transported when and if "steam" escapes into overlying rock. Shinohara (1994) and Shinohara and Fujimoro (1994) showed experimentally that the concentration of H C1 in a vapor in contact with brine at 600C becomes greater as pressure is decreased from 2,000 to 400 bars, and it is well established that HC1 tends to partition into a vapor phase in contact with boiling brine (Hemley et al., 1992; Candela and Piccoli, 1995; Williams et al., 1995). However, at 600C in the system NaC1-H20, Fournier and Thompson (1993) measured an abrupt and a greater than predicted increase in the concentration of HC1 in steam when NaC1 began to precipitate at pressures below about 300 bars. This increase occurred because hydrolysis reactions that produced HC1 and NaOH by the reaction of NaC1 with H20 become important only at pressures sufficiently low for halite (and probably also NaOH) to precipitate. In addition, an order of magnitude more HC1 was obtained at comparable pressures and temperatures when quartz was present compared to when no quartz was present. Presumably this oc- cured because the quartz reacted with NaOH to form sodium silicate.

In natural systems the generation of HC1 by hydrolysis is likely to be even more efficient in the presence of mica and plagioclase, with albite, K feldspar, and epidote being prod- ucts of the reactions (Fournier and Thompson, 1993). In addition, Bischoff et al. (1996) investigated hydrolysis in the system CaCls-HsO, yielding HC1 and Ca(OH)s at 380 to 500C.

At 500C the concentration of HC1 in the vapor phase abruptly increases at pressures below about 520 bars and at 380C at pressures below about 230 bars. Also, in contrast to the system NaC1-H,20, the hydrolysis reactions proceed without precipitation of CaCls. The highly associated and nonre- active HC1 that enters the vapor phase becomes reactive only where cooling and condensation of steam occurs.

The separated "steam" phase The physical characteristics of the "steam" that separates

from boiling brine during decompression from lithostatic to hydrostatic Pf, and how these characteristics change with additional decompression, are important in regard to mineral transport and ore deposition. At temperatures >374C and P >220 bars this "steam" would be called a supercritical fluid when decoupled or moved away from its parent brine. The types of decompression paths that may be taken by this supercritical fluid are shown in Figure 7, which is a graph of enthalpy versus pressure for pure water, contoured with selected isotherms. In Figure 7, point B is the critical point for water. Increasing concentrations of dissolved salt will move the critical point to higher temperatures and higher pressures (Sourirajan and Kennedy, 1962). There is a field of liquid water + steam at pressures below the critical pressure (220 bars for pure water). The boiling point curve for water extends downward to lower pressures and temperatures to the left of the critical point, and the condensation point curve (or dew point curve) extends downward to the right of the critical point. A field of liquid water is to the left of the boiling point curve, and a field of superheated steam is to the right of the dew point curve. For the probable composition of a natural fluid, the 2-phase field would be somewhat larger than the field shown in Figure 7.

400[ I II ]] o Sup .... itical / fluid

200 2kin I. // 19 \'. Super leated]

[ Boiling pt / \\l \ J

0 , . . c . . .... ' 0kin 0 500 1000 1500 2000 2500 300 3500

Enthalpy, Joules gin- 1

Fig. 7. Pressure-enthalpy diagram or H20 showing selected temperature contours. Arrows indicate adiabatic (vertical) and partly conductive (decreasing enthalpy with decreasing depth) cooling paths. See text or discussion.


Adiabatic decompression of "steam" (or supercritical fluid) that has an initial enthalpy greater than about 2,800 Joules gm - will bring that fluid into the field of superheated steam or gas (for example, the path from point F toward G in Fig. 7). Fluid following this path would discharge at the earth's surface as a superheated steam vent. Along this path there is very great expansion of the fluid (resulting in underground deposition of the less volatile dissolved constituents) but no formation of a true liquid. In the absence of liquid water, any HC1 or SOs initially present in the "steam" remain unreactive. Therefore, there would be little acid alteration of the wall rocks associated with the gaseous discharge. This apparently was the type of path taken by the mineralizing fluid at Alunite Ridge, men- tioned above and discussed by Rye (1993).

Fluids having initial enthalpies between about 2,060 and 2,800 and Joules gm - will, upon adiabatic decompression, move into the field of superheated steam at a pressure below the critical pressure, and then intersect the dew point curve where liquid water is produced by condensation of previously superheated steam. An example would be point H moving toward point D in Figure 7. At point D the less volatile dissolved constituents would partition into the newly forming liquid phase (point E in Fig. 7) while the steam would retain most of the noncondensable gases. The newly condensed liquid water would be very acidic because HC1 carried in the superheated steam would immediately dissolve in that water and dissociate to reactive H* and CI-, as discussed above. In addition, SOs would react in the liquid water, producing HsS and HsSO4 (Holland, 1965). The net result of this decom- pressional condensation would be a very corrosive liquid, initially unsaturated with respect to many of the minerals in the wall rock. Continued decompression increases the mass of liquid water relative to steam, but the volume occupied by the remaining steam steadily increases. This expansion of steam keeps fluid moving rapidly upward through the system and may result in brecciation at shallow levels. A similar scenario would apply to fluids that initially have enthalpies greater than about 2,800 Joules gm - but cool partly by conduction during decompression so that the superheated steam inter- sects the dew point curve (e.g., path F-J, Fig. 7). This would be a likely situation at the outer margins of a shallow, high- temperature system, away from the main discharge zone.

Adiabatic decompression of"steam" initially at lithostatic Pf and about 425C in plastic rock at a depth of 2 km is shown by the arrows from point A to point B in Figure 7. The 2- phase field of liquid + steam is intersected near the critical point for pure water (point B) at about 220 bars and 374C. Additional decompression to about 165 bars would produce nearly equal masses of liquid water (point E) and steam (point D). Slight conductive cooling in addition to the de- compressional cooling will cause supercritical fluid at point A to move first into the liquid-only field and then intersect the 2-phase field on the boiling point curve, such as point E in Figure 7. Further decompression will follow along the well- known boiling point curve. This would be a favorable path for the formation of bonanza-type Au deposits (discussed below). With significant conductive cooling along the decompression path (a slower rate of flow) the initially supercritical fluid moves into and stays within the liquid only field with no boiling during ascent to the earth's surface (e.g., the path from

point A toward L in Fig. 7). This cooling path would favor water-rock chemical equilibration along the flow path and re- suit in the conversion of all SOs to HsS (Ohmoto and Rye, 1979; Rye, 1993). The ore deposition model of Hemley and Hunt (1992) is applicable for these conditions. Deposition of silica from decorepressing "steam" and formation of bonanza Au deposits

In Figure 8a and b the solubility of quartz in supercritical HsO is shown at 425C and 1,000, 750, 500, and 250 bars, plotted as a function of the enthalpy of the fluid. For comparison, the solubilities of quartz and amorphous silica in liquid water and quartz in steam are shown for conditions along the boiling point curve for water. Consider a situation in which 425C brine in plastic rock at a depth of about 2 km (500 bars Pt lithostatic pressure) experiences a sudden decompression and starts to boil. The decompression path and intersection of the 2-phase liquid-steam field was discussed previously (path A-B-E, Fig. 7). The solubility of quartz in the "steam" that begins to separate from the boiling brine is about 1,180 mg kg -1 (point A, Fig. 8a). For now, assume that the "steam" rapidly becomes saturated with respect to the solubility of quartz, and thereafter does not precipitate quartz in response to the rapid adiabatic decompression. The concentration (wt %) of silica in the decompressing supercritical fluid would remain constant until the boiling point curve is intersected, and then silica would follow the trend A to E in Figure 8a. Supersaturation with respect to the solubility of amorphous silica at the vapor pressure of boiling water would be attained at >350C.

In contrast to the above adiabatic cooling scenario, if the "steam" moved relatively slowly along the decompression path and cooled by conduction without precipitating quartz (an unlikely circumstance during slow cooling) it would become supersaturated with respect to amorphous silica only at temperatures below about 230C (path A-L-H, Fig. 8a). If there were precipitation of quartz along the cooling path, supersaturation with respect to amorphous silica probably would not occur.

A more likely scenario would result in (1) some precipitation of quartz during the initial rapid decompression and decrease in temperature from about 425 to 350C, (2) a small component of conductive cooling, such as shown by path A to E in Figures 7 and 8b, and (3) great supersaturation in respect to the solubility of quartz at the vapor pressure of the solution everywhere along the flow path. In Figures 7 and 8b boiling of liquid water starts at point E, and this boiling con- tinues along the boiling point curve with additional upward flow and decompression. The separation of steam results in an increase in the concentration of silica in the residual liquid along the path E-F-G that moves directly away from point S in Figure 8b. At point F (about 260C and 47 bars) the boiling liquid becomes saturated with respect to amorphous silica. Additional decompressive boiling results in high degrees of supersaturation with respect to amorphous silica (point F- G). Polymerization of dissolved silica occurs, which produces colloidal particles dispersed in the liquid. Some colloidal silica particles probably would become attached to the walls of fractures at or near their point of formation, while others would be swept along in the rapidly moving, boiling water and


2000

1500-

lOOO

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0

0

01000 bars 425C E O 750 bars

//' '--' 425C Amorph Silica / _ , V.E soln ' QtzV. E soin

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/ .- 250 bars

I i I I I i I I I I I , I I I I I : : I I I i I I ilk

500 1000 1500 2000 2500 3000 Enthalpy, J/g

2000

1500

lOOO

500

1000 bars 425oc

A750 bars '--, 0 425oc

Amorph / AS00 ?rs Silica / /, 425 C V. Esoln ' \O - '

. [ Qtz E 165 rs V.P. sol s0 C

''cal Pt. / / C 221 bars

o- - 250 bars

,r ,r, SO0 1000 1500 2000 2500 3000

Enthalpy, Jig Fig. 8. Dissolved silica-enthalpy diagram showing solubility of quartz (solid line) in liquid water and in steam at the vapor

pressure of the solution (from Fournier and Potter, 1982b); circles show the solubility of quartz at 425C and the indicated pressures (calculated using equation of Fournier and Potter, 1982a). Dot C is at the critical point for water. Dot S shows the approximate enthalpy of steam in the range 175C and 9 bars to 350C and 165 bars. The dashed line shows the solubility of amorphous silica in liquid water at the vapor pressure of the solution (Fournier and Marshall, 1983). Amorphous silica solubilities above 300C estimated taking account of variations in fluid densities and the solubility variations of quartz over the same temperature and pressure range. Temperatures along the abscissa are for boiling liquid water with corresponding enthalpy values. (a). Arrow from A toward E shows silica changes resulting from adiabatic cooling from 425C and 500 bars to 350C and the vapor pressure of the solution, assuming no precipitation of silica; horizontal arrows from A to L and H show path taken during conductive cooling to less than 250C at the vapor pressure of the solution, assuming no precipitation of silica. (b). Path A to E shows effect of partly adiabatic and partly conductive cooling from 425C and 500 bars pressure to 350 and 165 bars pressure with some precipitation of silica; path E-F-G shows effect of boiling at the vapor pressure of the solution below 350C without precipitation of silica. See text for discussion.

deposit elsewhere along the flow path, particularly at throt- tling points.

Repeated episodes of rupturing, healing, and rerupturing of the brittle-plastic boundary result in a buildup of banded layers of amorphous silica in the surrounding brittle rock. Later crystallization of these amorphous silica layers produces bands of interlocking anhedral quartz (Boydell, 1924; Lover- ing, 1972; Ramdor, 1980; Fournier, 1985). At the same time, boiling could lead to the solution becoming supersaturated with Au (Drummond and Ohmoto, 1985; Spycher and Reed, 1989), and this supersaturation also could result in the formation of dispersed colloidal particles of Au (Saunders, 1994, and references therein). Saunders (1990, 1994) presented textural evidence that the transport of colloidal Au along with colloidal silica was important in the formation of the Sleeper bonanza Au deposit in Nevada.

The accumulation of magmatic fluids in plastic rock at about 400 to 425C in the depth range 2 to 3 kin, and episodic discharge of that fluid into the brittle, hydrostatically pressured domain appears to be optimal for the formation of bonanza Au deposits for the following reasons: (1) in this temperature and pressure range most, if not all, of the SO2 that may have been present initially in the magmatic fluid will have been converted to H2S, enhancing the solubility of Au (Benning and Seward, 1996; Gammons and Williams-Jones, 1997); (2) the solution acidity would probably be in an appropriate range for deposition of K feldspar (commonly present in the silica bands that host the ore) because the Pt would be sufficiently high to inhibit the production of HC1 by hydrolysis of salt, and little SO would be available to produce

HSO4; (3) adiabatic decompression of lithostatically pressured fluid in this depth-temperature range will result in the relatively dense "steam" phase (supercritical HO) intersect- ing the field of liquid + steam on the boiling point limb rather than on the dew point limb, which is favorable for keeping dissolved constituents in solution down to lower pressures; and (4) the solubility of silica, both in brine and in coexisting "steam" that form as a result of decompression, is sufficiently elevated that a high degree of supersaturation with respect to the solubility of amorphous silica is likely to occur after cooling to below about 250 to 300C.

At depths less than about 1 to 1.5 km (250-375 bars lithostatic Pf) and at >425C, the decompression of"steam" would not be likely to lead to major silica deposition along the flow path. At 425C and 250 bars the solubility of quartz in "steam" is only about 100 mg kg 4 (Fig. 8a), and at 700C and 250 bars it is only about 316 mg kg 4 (Fournier and Potter, 1982a).

A General Model of Physical Processes in a Shallow Plutonic Environment Spanning

the Brittle-Plastic Transition

A general physical model for the transition from magmatic to epithermal conditions in a shallow, subvolcanic environment is shown schematically in Figure 9a. It is assumed that there have been multiple injections of magma to shallow levels, in accord with estimated depths to the tops of porphyry Cu systems (Sillitoe, 1973; Nielson, 1976; Dilles and Einaudi, 1992), so that a considerable volume of rock around the last


intrusion is at a temperature greater than 400C. In this figure, the tops of degassing intrusive bodies would probably be in a depth range of i to 3 km. A depth scale is not shown in Figure 9a because the general model is applicable for tops of intrusions less than i km deep and deeper than 3 km, depending on many factors, including the topographic relief, the elevation of the water table, the vertical extent of any vapor-dominated system, and the partial pressures of noncondensable gases that might be present. The main constraint on depth in Figure 9a is that Pt in the brittle rock immediately overlying the plastic region be less than about 300 bars so that brine will boil, producing a large volume of "steam," when there is a drop in pressure resulting from a breach in the self- sealed zone (Fig. 9b).

The lake that occupies the small summit crater, and the vapor-dominated region within the volcanic pile, are common features of natural systems and affect subsurface Pf. They are not necessary for the model. Where a vapor-dominated system is present, its vertical dimension may be a few tens of meters or it may exceed i to 2 km, extending down to the brittle-plastic transition, as may be the present situation at The Geysers geothermal field in California (Walters et al., 1992).

Relatively dilute water (dominantly meteoric in origin) circulates through brittle rock at hydrostatic pressure where temperatures are less than about 360 to 370C. Details of the chemical variations found within that convecting hot spring system are not shown in Figure 9a because they have been discussed extensively by others (e.g., Old and Hirano, 1970; Henley and McNabb, 1978; Henley and Ellis, 1983; Hedenquist and Lowenstern, 1994). While the self-sealed zone remains intact (Fig. 9a), ongoing acid alteration is chiefly confined to near the earth's surface where H2S is oxi- dized to H2SO 4. The transition from brittle to plastic behavior occurs across a narrow band of rock, generally at temperatures in the range 370 to 400C. Within plastic rock, hypersaline brine and "steam" accumulate in isolated, relatively thin, mechanically stable horizontal lenses or networks of fractures that have limited vertical interconnectivity, as discussed above.

An important point to keep in mind when looking at the distribution of "steam" in Figure 9 is that at moderately high temperatures (>374C) and high pressures (>220 bars) it generally would be considered a supercritical fluid that, because of its moderately high density, is capable of transporting relatively high concentrations of dissolved constituents, such as silica (discussed in detail above) and metals. The density of "steam" at 425C and lithostatic Pt at a depth of 2 km would be about 0.13 gm cm a. By comparison, steam as we generally know it at 100C and atmospheric pressure has a density of only 0.00062 gm cm a.

Heinrich et al. (1992) found appreciable Cu concentrations in vapor-rich fluid inclusions in reduced porphyry Sn-W deposits, possibly carried as a volatile Cu sulfide complex. Bod- nar (1996) also found fluid inclusion evidence for Cu partitioning into the vapor phase in the Bingham porphyry Cu deposit. It would appear that a major factor in the ability of moderately high density "steam" to carry dissolved Cu (and probably also Au) is the concentration in that "steam" of HS, which can act as a complexing agent.

Slow movement of fluid across the brittle-plastic boundary Dense hypersaline brine that slowly escapes across the brit-

tle-plastic boundary into the hydrostatically pressured part of the system will tend to move outward and downward by gravity flow along open fractures in brittle rock. Because of the relatively slow rate of movement, water-rock chemical reequi- libration along the flow path is likely to occur in response to declining temperatures and pressures. If the leakage occurs at a depth shallow enough for the brine to boil upon entering the hydrostatically pressured domain, most of the separated "steam" would probably condense by conductive loss of heat upon entering the cooler rock (path A-L, Fig. 7). The condensate would be a liquid of relatively low salinity, which may mix with local ground water.

Where slow leakage occurs 2 to 3 km deep, separated "steam" would probably contain relatively little HC1 and SO, so the condensate would not be very acidic. In contrast, where leakage occurs at less than about 1.5 to 2 km, the cor- rosiveness of the condensate would be greater because the "steam" would carry increased amounts of both HC1 and SO, as discussed previously.

Rapid movement of fluid from across the brittle-plastic boundary

Episodically, major breaches of the self-sealed brittle-plastic transition zone occur (Fig. 9b) as a result of one or more of the mechanisms discussed above. The pressure surge into the brittle region initiates faulting there, which increases permeability and facilitates the movement of the magmatic fluid upward and outward. The rapid escape of fluid from >400C rock results in a temporary drop in Pt, vaporization of brine in confined spaces, and expansion of "steam." This expansion results in brecciation, increases the rate of expulsion of fluid across the initial brittle-plastic interface, and increases the stress difference and the strain rate within the >400C material. This, in turn, allows a downward propagation of brittle fractures and brecciation into previously plastic rock. Rapid deposition of quartz, amorphous silica, and other minerals along channels of flow counteract the processes that increase permeability. Porphyry-type ore deposition may occur at this stage of activity if the drop in pressure extends down into in- completely crystallized magma that is rich in S, Cu, and/or Mo.

If the brittle-plastic transition is at a relatively shallow depth (


alteration

hot

Hydrostatic Pf Dilute to moderate salinity

(xm Acid-sulfate alteration

:. Hydrostatic Pf ' '. ' :1 E J a '-

'X \'.'- :' Pressure surge V' o . --

/:' . , ]'{ "Steam" + salt lX 1'

4 ' ':'[, k + k %.i:' - Dilute to

*' / Chied +% moderate s W

t+ , Hydrosc Pf

Fig. 9. Schematic model of the transition from magmatie to epithermal conditions in a subvolcanic environment where the tops of intruded plutons are at depths in the range 1 to 3 km. (a). The brittle to plastic transition occurs at about 370 to 400C and dilute, dominantly meteoric water circulates at hydrostatic pressure in brittle rock while highly saline, dominantly magmatic fluid at lithostatic pressure accumulates in plastic rock. (b). Episodic and temporary breaching of a normally self- sealed zone allows magmatic fluid to escape into the overlying hydrothermal system. See text for discussion.

1208 ROBERT 0. FOURNIER

may mix with brine droplets, producing a brine of intermedi- ate composition.

The rapid expulsion of "steam" and brine into the overlying, previously hydrostatically pressured part of the system results in displacement of the relatively dilute, meteoric-derived water upward and outward ahead of the advancing magmatic fluid (Fig, 9b). This is an ideal environment for initiating many types of brecciation (Sillitoe, 1985). In particular, local heating and expansion of ground waters in small fractures adjacent to the main channels of fluid flow result in hydraulic fracturing and brecciation. Phreatic explosions are likely to occur above the most permeable channels of upward flow. Such eruptions have been suggested by many investigators to have initiated ore deposition in Miocene Kuroko deposits (e.g., Horikoshi, 1969; Henley and ThornIcy, 1979). Where considerable condensation of highly expanded "steam" occurs by conductive loss of heat as it moves through cooler rock, there may be a decrease in the volume of upward-moving magmatic fluid, thereby decreasing the magnitude of the pressure surge.

Eventually igneous intrusive activity wanes, the plutons fin- ish crystallizing and begin to cool, and the brittle-plastic transition recedes downward. Low-temperature hydrothermal alteration is superimposed on the earlier magmatic alteration, and sulfides are remobilized.

A discussion of the transport of metals in the fluids that might evolve at the various depths and temperatures shown in Figures 5 and 6, or the types and place of epithermal ore deposition in the environments shown in Figure 9, has not been attempted because of the large number of variables regarding initial compositions and possible decompression paths, and a need for more experimental work on the degree to which metals partition among the phases that are present through- out the pressure and temperature range appropriate for the shallow environment of epithermal mineralization. A good start in this latter endeavor has been made by experiments at relatively high pressures looking at the partitioning of halo- gens and metals between magmas and coexisting brines (e.g., Candela and Holland, 1984; Candela and Piccoli, 1995; Williams et al., 1995; Gammons and Williams-Jones, 1997; Webster, 1997) and studies of metals in fluid and melt inclusions (Heinrich et al., 1992; Cline and Bodnar, 1994; Bodnar, 1995, 1996; Lowenstern, 1994, 1995). The experiments of Hemley et al. (1992) in which Fe-Cu-Zn-Pb sulfide solubility relationships were determined in rock-buffered systems at 300 to 600C and 0.5 to 2 kbars are appropriate for deeper epithermal deposits where boiling does not occur. Hemley and Hunt (1992) discussed in detail probable sequences of hydrothermal alteration and ore deposition from saline solutions based on the Hemley et al. (1992) experimental results. However, the pressure-temperature range of the Hemley et al. (1992) experiments restricts their application mainly to cooling paths where boiling is not important (leakage across the brittle-plastic transition at depths >3-5 km) and for relatively reduced conditions in which essentially all the sulfur is present as H2S. Thus, these experiments probably bear greatly on the formation of low-sulfidation deposits and less directly on the formation of high-sulfidation deposits where boiling and condensation of acidic superheated steam are important processes (Hedenquist, 1987; Hedenquist et al., 1996).

Summary and Conclusions

In subvolcanic environments, meteoric fluids circulate at hydrostatic pressure in brittle rock where temperatures are less than about 370C, whereas brine and "steam" exsolved from crystallizing magma tend to accumulate at >400C in thin, overlapping horizontal lenses in plastic rock at lithostatic pressure (where the least principal stress is the lithostatic load). Where magmatic bodies have been repeatedly intruded to relatively shallow depths, a large volume of 400 to 500C plastic rock develops, and the brittle-plastic interface occurs at depths as shallow as i to 2 km. A combination of precipitation of vein minerals as a result of heating to >350C at the low-pressure, low-temperature side, and plastic flow and de- compressional deposition of vein minerals at the high-pressure, high-temperature side creates a self-sealed zone at this interface. Temperature and fluid pressure gradients across the interface are very steep. Episodically, breaches of the self- sealed zone occur, and "steam" and/or brine that had ponded in plastic rock are discharged into the brittle, hydrostatically pressured domain. Faulting, brecciation, hydrothermal alteration, and vein mineralization ensue. Banded veins extending far into the brittle domain result from repeated cycles of breaching, healing, and rebreaching of the self-sealed zone.

A suggested mechanism for initiating the breaching of the self-sealed zone is upward injection of a new pulse of magma that increases the local strain rate to such a degree that the brittle to plastic transition is temporarily moved to a deeper and hotter condition. This, in turn, allows brittle failure within the previously plastic rock in response to very small stress differences where lithostatic Pf prevails. New pulses of magma injection also heat previously existing brine trapped in plastic material, causing rapid expansion of that fluid, which also leads to an increase in the strain rate.

Discharge of exsolved fluids directly from magmas into the brittle regime yields fluids with very high SO2/HS ratios at shallow depths, and somewhat lower SOs/HS ratios at greater depths. SOs/H2S ratios also decline when fluids pond in plastic rock at temperatures considerably below magmatic. Thus, the depth and reservoir temperature of fluids discharged from plastic rock into the hydrostatic domain has an influence on whether a high- or low-sulfidation epithermal deposit forms. Magmatic alunite vein deposits form without acid alteration of the adjacent wall rocks where the "steam" (supercritical fluid) that separates from boiling brine has an initial enthalpy greater than about 2,800 to 2,900 Joules gm -, and there is little loss of heat by conduction during decompression. In contrast, adiabatic decompression of "steam" having initial enthalpies in the range 2,100 to 2,800 Joules gm - will result in its movement into the field of superheated steam (at a pressure


field at pressures below 220 bars, followed by intersection of the boiling point curve. The initial "steam" would be rich in H2S and contain relatively little SO2 and HC1. This environment favors the development of low-sulfidation bonanza Au deposits.

The contraction within a dike that occurs as a result of crystallization and cooling can lead to a local tensile stress field in which the least principal stress in plastic rock acts normal to the wall of the dike and nearly at a right angle to the lithostatic load. In this situation hydraulic fractures may open and fluids may be propelled upward relatively quickly through rock at >400C. Thus, dikes can be effective conduits for the upward flow of fluids through plastic rocks that otherwise would impede upward movement of hydrothermal fluids. The fluid that moves upward through these dikes may become ponded in overlying plastic rock where the least principal stress is the lithostatic load, or it may move through the brittle-plastic boundary and discharge into the hydrostatically pressured region. Mineralization is likely to occur during this process as a consequence of declining pressure.

Acknowledgments George Bergantz and John L. Muntean provided helpful

formal reviews of this manuscript, and informal discussions with Jeffrey w. Hedenquist have been very appreciated. Jeff Hedenquist and John Muntean called my attention to several important references that I had overlooked. I am particularly indebted to John Muntean for providing a copy of his unpub- lished Stanford Ph.D. dissertation, "Magmatic-hydrothermal gold deposits of the Maricunga belt, Northern Chile." That dissertation contains some of the best descriptions of textures and mineralization within porphyry Cu-Au deposits and con- temporaneous epithermal vein deposits that I have come across. He independently came to many of the conclusions presented here and has related Cu-Au deposition to brittle- plastic phenomena in the magmatie-epithermal environment. July 20, 1998; May 12, 1999

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