Isotopic exchange in mineral-fluid systems: III. Rates and mechanisms of oxygen isotope exchange in the system granite-H2O ± NaCl ± KCl at hydrothermal conditions

Geochimica et Cosmochimica Copyright 0 1992 Pergamon

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Isotopic exchange in mineral-fluid systems: III. Rates and mechanisms of oxygen isotope exchange in the system granite-H20 f NaCl k KC1 at hydrothermal conditions DAVID R. COLE,’ HIROSHI OHMOTO,‘*~and GARY K. JACOBS~ ‘Chemistry Division, Oak Ridge National Laboratory, Oak Ridge, TN 37831,USA *Department of Geosciences, Pennsylvania State University, University Park, PA 16802, USA ‘Institute of Petrology, Mineralogy, and Economic Geology, Faculty of Science, Tohuku University, Sendai, Japan 4Environmental Science Division, Oak Ridge National Laboratory, Oak Ridge, TN 37831, USA (Received March 14, 199 1; accepted in revised form October 30, 199 1)

Abstract-Variations with time in the alteration mineralogy and the oxygen isotopic composition of solutions, minerals, and rocks have been experimentally investigated in the system granite-Hz0 +- NaCl + KC1 at T = 170” to 300°C, P = L/V to 0.3 kb, water/rock mass ratios (0.2 to 6), and grain sizes ( -0.1 mm to 2.5 mm) for periods up to 1006 h. Alteration assemblages formed in the experiments are dominated by chlorite (after biotite) , &cite-zeolite-albite (after K-feldspar and plagioclase ) , and hematite (after magnetite and pyrite). The abundance of these minerals continued to increase with increasing reaction time, temperature, NaCl (or KCl) concentration of solution, and surface area of solids. However, major element concentrations attained steady state in less than 330 h. Reaction of rock (6 “0, N +8%0 ) and aqueous solution (6 180, N - 10% ) resulted in depletion of “0 in the solid and enrichment of “0 in the fluid. The magnitudes of change increase as temperature, time, salt concentration, and surface area increase. The trends in isotopic shifts are directly related to changes in the style and intensity of mineralogic alterations in the granite. The degree of isotope exchange in the experimental systems between granite (minerals) and solutions was computed from the comparison with calculated equilibrium fractionation factors and yield values of less than 5 to 50% exchange. The 6180 changes of the rocks, minerals, and fluid were observed to follow closely with those expected from a firstorder rate law. At temperatures between 170 and 300°C rate constants for the oxygen isotope exchange between granite and pure water range from 10m9 to 10e8.’ (moles 0 mm2 see-’ ) whereas rates in the granite-O. 1 m NaCl system range from 1O-8.5to 10-7.6. Rate constants were also retrieved for biotite, K-feldspar, and plagioclase interacted with either pure water or 0. l-l m NaCl. Based on a simple closed-system model with (W/S) mass ratios between 0.5 and 5, we estimate the minimum time required for granite-fluid isotopic equilibration to be roughly 200 years or less for grains 1 cm in radius or smaller (porous-media analog) reacted at 3OO”C, and between lo4 and lo5 years for grains ranging from 0.1 to 1 m in radius (fractured-media analog) reacted at 300°C. between a mineral (rock) and fluid proceeds in two steps: ( 1) instantaneous isotopic effects (i.e., kinetic isotope effects) accompanying formation of a new mineral (i.e., alteration mineral including recrystallization of an old phase) and (2) isotope exchange between the new mineral and solution. The final products of the alteration process depend on the fluid composition, the chemical, mineralogical, and physical properties of the solid (e.g., porosity, permeability, grain size), the fluid/solid ratio, temperature, and pressure. Significant amounts of isotopic exchange involving these mechanisms may occur within a relatively short period of time. For example, oceanic tholeiitic basalts experimentally reacted with seawater between 300 and 500°C for durations up to 272 days exhibited pronounced depletion in I80 in response to the formation of product phases dominated by smectite, albite, zeolites, quartz, and anhydrite (COLE et al., 1987). Once chemical equilibrium has been attained between the fluid and the rock and no new minerals are formed, isotopic exchange will occur at much slower rates by a diffusion mechanism. Isotope exchange controlled by diffusion is typically several orders of magnitude slower than surface reaction-controlled exchange (COLE et al., 1983) and usually can be ignored for rock-water systems where significant mass transfer is occurring-e.g., in weathering zones, hydrothermal ore deposits, geothermal systems, and altered oceanic crust.

INTRODUCTION OXYGEN ISOTOPESTUDIESof natural systems have revealed the presence of mineral assemblages clearly out of isotopic equilibrium (e.g., TAYLOR and FORESTER, 1979; GREGORY and TAYLOR, 1981; CRISS and TAYLOR, 1983, GREGORYet al., 1989) or yielding isotopic temperatures inconsistent with those estimated from phase equilibrium or fluid inclusion evidence. Collectively, these studies indicate that ( 1) isotopic disequilibrium can occur at high as well as at low temperatures, (2) different minerals exhibit varying susceptibilities to retrograde exchange (e.g., feldspars versus quartz), and (3) the mechanisms of isotope exchange are varied and depend on the geochemical conditions (e.g., fluid chemistry, temperature) at the time of the interaction. Recent emphasis on the systematics of stable isotope exchange kinetics between minerals and fluids in natural systems has been made in the context of modal mineralogical variations and mass balance, both in open and closed systems (GREGORY et al., 1989). Critical to the interpretation of the extent of isotope equilibrium in natural systems is the knowledge of the rates and mechanisms of isotope exchange between rocks (minerals) and solutions. In the previous two papers of this series (COLE et al., 1983, 1987)) we have suggested that the rate of isotopic exchanp 445

D. R. Cole, H. Ohmoto, and G. K. Jacobs

446

For cases where both diffusion and surface reaction contribute significantly to isotope exchange, usually at temperatures above about 4OO”C, the two mechanisms can be treated simultaneously using finite difference methods ( BURCHet al., 1990). This paper presents the results of our experimental study on the oxygen isotope exchange between granite and aqueous solutions containing up to 1 m NaCl or 0.1 m KC1 at temperatures ranging from 170 to 300°C. The chemical and Table 1. Run Number

I.

Temp. ('C)

mineralogical changes attendant with the interaction ofgranite and water have been the focus of numerous experimental studies (e.g., CHARLESand BAYHURST, 1983; BENJAMINet al., 1983; SAVAGE, 1986, RAFALSKIYet al., 1987). None of these studies, however, examined the consequences of fluidrock interaction on the redistribution of the stable isotopes. The main objectives of this study were ( 1) to determine the effect of temperature, solution chemistry, mineralogy, fluid/ rock ratio, surface area, and time on oxygen isotope shifts

Summary of Experimental Conditions for the Granite-Water Runs. Rock Type (a)

Initial Fluid Comp. (b)

Initial Fluid 6 '0 (c)

Mass (gm) Rw. W. (d)

Surface

Duration (Hrs.)

Circulating or Rocking AutoclaveSystem

1c

170

wgn

w

-10.3

1277.2 640.7

76.9

444

1L 1M

250

bqm

W

-10.0

1525

520.2

91.9

122

200

-10.1

1497.7 511.9

2A 2B

200 170

bqm bqm

W

-11.6 -11.6

1319.4 636.6 1262.3 647

90.2 79.5

414

v-cm

O.lm NaCl O.lm NaCl

2T

300

v-cm

O.lm NaCl

-10.5

706.3

2s

200

bqm

O.lm NaCl

10

170

O.lm NaCl-

-9.0 -9.4

lHA3 R4Al

285

bqm bqm

0.5m NaCl

-10.3

883.6

0.0083

73

O.lm NaCl

-11.0

162.3

RlA2

300

l.Om NaCl

-11.1

149.8

665.3

0.0047 0.0083

772 667

RlA4

250

bqm bqm micr

149.8 85

RlA6 RlA7

290 270

micr micr

l.Om NaCl l.Om NaCl

-8.8 -8.8

10 10

l.Om NaCl

-8.8

10

256.8 258.7

II.

285

76.0

204

42.5

291

1339.5 607.1

84.3

472

1387.4 569.9

83.6

343

666

600

0.300

120

600 600

0.300 0.300

121 127

69.8

6.46 6.51

838

Static System 200

12

300

grgn grgn

W

-10.8 -11.1

13

250

grgn

W

-8.4

261.7

15 16

200 250

bqm bqm

W

-10.0

W

255.5 256.0

17 18 21

300 300 200

bqm bqm

W

-9.8 -10.6

22 23 24

300' 200 250

11

354

grgn grgn bqm grgn

W

O.lm O.lm O.lm O.lm O.lm

NaCl NaCl NaCl NaCl NaCl

O.lm NaCl

-10.2 -9.7 -11.7 -10.1 -11.0 -10.7

253.6

71.4 71.5 54.6 54.5 64.4 72.9

260.5 252.9 254.5 256.9 250.2

61.1 74.5 51.0 64.9

257.9 259.9

72.0 69.9

6.58 6.43 6.43 6.38 6.55 6.36 6.40 6.46 6.29 6.49

834 834 834 1006 840 840 835 809 836 841

838 250 bpm 25 O.lm KC1 -10.0 6.52 840 250 bqm 26 NOTES: (a) Rock Type - grgn is granite gneiss, bgm is biotite quartz monzonite, micr is microcline + quartz + amorphous silica (b) w&s pure distilled water (c) 6 0 is relative to SMOW (d) Total weights of solid and water (e) Total surface areas based on B.E.T. absorption method (see text for specific surface area values)

Oxygen-isotope exchange between granites and brine during experimental granite-water interaction, (2) to determine the rate constants and activation energies for the rockfluid isotope exchange and individual mineral-fluid exchange, and (3) to compare these rates with those estimated for simple mineral-fluid systems (COLE et al., 1983; COLEand OHMOTO,

447

,,

Line to Presrure Intensifier

1986) and the basalt-seawater system (COLE et al., 1987). The fourth paper in this series will address the application of these rate data to quantifying the duration of isotopic exchange in natural hydrothermal systems. EXPERIMENTAL METHODS A total of twenty-seven granite-fluid and microcline-fluid experiments were conducted at temperatures ranging from I70 to 300°C. Solution compositions were varied from pure distilled water to 1 m NaCl (or 0.1 m KCl), with the time of interaction ranging from 120 to 1006 h. The experiments were carried out in one of three types of hydrothermal apparatus-rocking autoclaves, circulating flow-through autoclaves, and static stacked-vessel autoclaves. These systems are explained in detail in COLE ( 1980), and only briefly summarized below. Table 1summarizes the experimental conditions used in this study.

Experimental Systems Rocking autoclaves and circulating flow-through hydrothermal systems

One seriesofhigh-temperature experiments was conducted in either Barnes-type rocking autoclaves (BARNES, I97 1) or in constant volume, circulating hydrothermal systems (Fig. 1) These two systems allowed for fluid sampling through time and could accommodate more frequent sampling with little or no perturbation of the system. Solids were examined after the termination of the experiments. A known mass of granite (approximately 85 to 1500 gm) was loaded into a I 100 cc reaction vessel, which was evacuated and filled with a known mass of solution. Two different grain sizes were used in these experimental systems. One set ofexperiments employed grains with an average diameter of approximately 1 mm. A second set of runs used discs of granite approximately 2.5 cm thick by 7.5 cm in diameter. The water/rock mass ratios ranged from 0.3 to 6. The disc experiments with water/rock mass ratios ranging from 2 to 6 were designed to simulate a fractured, rather than a granulated-porous system where alteration products could be concentrated over a smaller surface area. In addition to the granite experiments, isotope exchange experiments were coupled with experiments carried out by SOLOMON (1978) who measured the rates of Na-K ion exchange at 250, 270, and 290°C in the system microcline + quartz + amorphous silica + HZ0 + NaCI.

FIG. 2. Schematic diagram of the static stacked-vessel system (after BARNESet al., 1979). Each vessel is approximately 30 cm long.

The autoclaves were heated to, and maintained at, the desired run temperature (25°C) while solution samples were periodically withdrawn. Figure I shows a schematic of the circulating system, with arrows indicating the direction of hydraulic flow. Circulation flow rates in this system were approximately 40 cm3/h at all temperatures and pressures. To fix the temperature of the circulating solution after it leaves the reaction vessel, the heat exchanger, solution sampling system and constant-volume circulating pump are all immersed in a constant temperature (65°C) bath. Fluid pressure was kept constant during sampling by a piston-driven hydraulic pump connected to a fluid reservoir. In contrast, rocking autoclave experiments experienced pressures equal to those along the liquid/vapor curve for the reactant solutions. Static stacked-vessel autoclaves This experimental system consists of a series of five reaction vessels (TUTTLE,1949; LUTH and TUTTLE, 1963) connected end-to-end as shown in Fig. 2. Each reaction vessel is contained in a separate splittype furnace with an independent temperature controller to reduce thermal gradients throughout the chain of five vessels to less than ?5”C. A high-temperature valve was positioned at the bottom of the lowermost furnace and maintained at the temperature of the reaction vessel. The individual reaction vessels, which have volumes of roughly 25 cm3 each, were filled with 50-55 gms of 0.1-1.0 mm diameter granite chips which had been washed to remove any adhering dust particles. Typically, each system of five reaction vessels contained a total of 50-75 grams of solution and 250 grams of rock at temperature (water/rock mass ratio of 0.3). The system was pressurized with argon to about 750 bars at 25°C. As the reaction vessels were heated to temperature, solution was bled off to maintain the system in the liquid phase at pressures around 350 bars. Samples were collected at one-week intervals, at which time the solution-argon interface was lowered to the next vessel, enabling constant pressure to be maintained on the fluid. This sampling procedure served to isolate the reactant grains in each successive vessel from the reaction solution at one-week intervals. The rocks in the uppermost vessel were reacted for one week and those in the lowermost vessel for five weeks. This procedure permitted the evaluation of chemical and isotopic compositions of both rock and solution at specific time intervals. Solutions were analyzed by atomic absorption spectroscopy. Starting Materials

FIG. 1.Schematic diagram of the circulating system (after BARNES et al., 1979). This figure is not drawn to scale.

Two different granitic rocks and one type of pegmatite were used in the hydrothermal experiments, a granite gneiss (grgn) and a biotite quartz monzonite (bqm) and microcline from the Spruce Pine pegmatite (spm). The chemical compositions of these solids are given in Appendix A, along with compositions of individual minerals, The minerals, their abundances and isotopic compositions (relative to SMOW scale) for the two rock types are shown in Table 2.

448

D. R. Cole, H. Ohmoto, and G. K. Jacobs specificareaof 5.0 mlgrn-' , corrected for fracture and grain-boundary porosity as outlined by CATHLES( 1983) and NORTONand KNAPP (1977). The chemical composition of the microcline, sampled from the Spruce Pine Pegmatite, North Carolina, and used in the Na-K ion exchange experiments by SOLOMON( 1978), is given in Appendix A and its isotopic composition is included in Table 2. The specific surface area of the microcline was 3.02 X 10-’ m* gm-’ , whereas 9.12 X 10e2 m* gm-’ and 9.34 X lo2 m2 gm-’ were measured for the quartz and amorphous silica, respectively. All three surface areas were determined by the N2-BET absorption method (SOLOMON, 1978 ) Solutions used in the granite-fluid experiments were pure water, 0.1 m NaCl, 0.5 m NaCl, I m NaCl, and 0.1 m KC1 solutions. Because these hydrothermal experiments were conducted in large-volume systems (greater than 35 cc), it was necessary to use large quantities of solutions. Consequently, each starting solution had its own unique isotopic value (see Table 1). The rangein b"O values fortheexperimental waters was from about - 12%0to -8%0, relative to SMOW.

Several major differences exist between these rocks types that should be noted. The most obvious difference is that the amount of biotite present in the bqm is about twice that in the grgn. Secondly, plagioclase in the bqm is an oligoclase containing about 29% anorthite component. Plagioclase in the grgn contains less than 10 mol% anorthite and the alkali feldspar is almost pure microcline. The feldspars in both rock types showed minor (less than I %) deuteric alteration to sericite or clay. Finally, the grgn exhibits a weak to moderate foliation, with grain sizes averaging less than 1.5 mm in diameter. This texture is in contrast to the bqm which exhibits a porphyritic texture with large microcline grains ranging up to 10 mm in diameter, set in a matrix of fine-grained (~2 mm) quartz, biotite, and oligoclase. To compensate for this grain size difference and enhance reaction rates in the experiments, each rock was crushed into grains ranging from 3 mm to less than 0.1 mm in diameter. Two splits were then taken from these grains, one containing coarse grains ( l-2 mm) and a second containing chips from 0. I to1 mm in diameter. The specific surface areas of these two fractions determined by the Kr-BET absorption method were 6.02 X 10e2 mZ grn-’ and 1.26 X 10-l m2 grn-’ for the coarse and fine splits, respectively (SOLOMON,1978). The coarse fraction was used exclusively in the circulating system while the fine fraction was used only in the static system. Discs of bqm employed in the high water/rock experiments had an average

Tabel

2.

Oxygen for isotopic analysis of both starting and run waters was liberated by reaction of water, introduced by a glass capillary, with BrFS in nickel reaction vessels at 225 f 25°C after a method modified

Mineralogical and Isotopic Starting Materials.

Phase I.

Isotopic Analysis

Granitic

0

wt. % Gneiss

‘i

Composition 6l*O

SMOW (O/00)

(grgn) 25.2

0.255

8.0

43.9

0.447

7.8

22.7

0.248

8.9

Biotite

4.7

0.041

3.6

Magnetite

1.5

0.008

6.4

Sphene

1.0

nd

nd

Pyrite

1.0

na

na

K-feldspar Plagioclase

(Anlo)

Quartz

Whole Rock

II.

Biotite

7.9kO.2

na

Quartz Monzonite

(bqm)

26.5

0.266

8.6

37.0

0.38

8.4

Quartz

22.6

0.254

9.9

Biotite

10.6

0.095

4.5

Magnetite

1.3

0.008

4.60

Hornblende

1.0

nd

nd

Pyrite

1.0

na

na

K-feldspar Plagioclase

(An2g)

Whole Rock III.

Spruce

na

8.5eO.20

Pine Microcline

X? = atomic fraction (a) respect' to whole rock. nd = not determined:

of the

10.0*0.1 of oxygen

in ith phase

na = not applicable.

with

449

Oxygen-isotope exchange between granites and brine from O'NEILand EPSTEIN( 1966). The oxygen was converted to CO1 for isotopic analysis by reaction with carbon produced by heating disc-shaped graphite electrodes with two projector lamps. If the conversion of the water to CO2 failed to achieve a yield of 99.970, a

second sample was converted. Whole-rock starting material and run products ( -25 grams for static, - 150grams for circulating runs) were crushed with a stainless steel automatic crusher and sieved to less than 100 mesh. Mineral separates were ground to less than 100 mesh in agate mortars. Whole rocks and mineral separates averaging 10 mg were reacted with BrFS in nickel reaction vessels at 600°C for a minimum of 12 h, as described by CLAYTON and MAYEDA( 1963), to liberate oxygen. Measurements are reported in the usual “8” notation in parts per thousand (“per mil”) relative to Standard Mean Ocean Water (SMOW, as defined by CRAIG,1961) with a precision of &O.lO%. The per mil notation is defined as &‘80, = lo3 (RJR,, - 1) where R = 1sO/‘6O and the subscripts x and std represent the sample and standard, respectively. EXPERIMENTAL Mineralogical

RESULTS

and Chemical Changes During Reaction

Detailed examination of run products by both X-ray diffraction and thin-section techniques indicate that a complex series of mineralogic reactions has taken place during granitewater interaction. Table 3 summarizes the major mineralogic trends observed in the experiments. Quantitative estimates of relative abundances for the various phases are based primarily on thin-section observations. The replacement of biotite by chlorite (7 A variety) is the most ubiquitous reaction observed during granite-water interaction involving rock chips. Chlorite appears after 160 h of reaction for all temperatures and persists for the duration

Table

Experiment TYPe

3.

Summary

of Alteration

1 week

(Salt runs) Static

(chip) (Bqm: Grqn) _ (pure water)

Bio - Chl Py + Hem Mte + Hem Feld + Zeal ?Qtz, Cc, Ser, Ab

Bio + Chl Plag + Zeol Py, Mte + Hem +Qtz, Ser

Disc (Bqm)

(pure water)

(Salt runs)

Mineralogy

Time 3 week

Bio + Chl Ksp + Ser Py, Mte - Hem Plag -t Zeal *Qtz, Cc, Clay

in Granite-Water

Experiments. +

5 week

Chl Bio Ksp Mte Feld

Hem + Ab -t Chl + Ser + Hem + Zeal

Temp =

most abundant t

Bio + Chl Feld + Zeal Py, Mte f Hem rQtz, Ser, Cc

?Otz: ED. Cc Bio + Chl + Hem Feld + Anal Py, Mte + Hem ?Qtz, Clay

Bio -t Chl, Verm Ksp * Illite Ksp 4 Mont, Kaol Illite + Mont Chl, Py, Mte + Hem *Qtz, Ser

Abbreviations: Bio = Biotite, Chl = Chloride, Hem = Hematite, Zeal = Zeolite, Feld = Alkali feldspars, Plag = Plagioclase, Anal = Analcime, Mont = Montmorillonite, Xonot = Xonotolite,

+

- Most Altered 2o0°c 25O'C 3o0°

Ah

Chl Ser Zeol Ab, Hem

t

Zeal Chl Hem

Chl Zeal Hem

Chl Ser Zeol

t

Ser

Ser Ab

Hem Ab

t

KSD + Ab

Bio + Chl Ksp + Ser Ksp + Ab Feld + Zeol *Hem. Otz Bio + Chl Py, Mte - Hem Feld + Ser, Mont ?Qtz, Cc, Xonot

(Salt runs) Circulating (chip) (Bqni Grgn)

of the experiments. In the case of the static system, increased chlorite abundances have been observed through time with almost total chloritization of biotite occurring by 800 h at 300°C in 0.1 m NaCl runs. Vermiculite replacement of biotite is observed in the bqm disc-salt runs in the first several hours. However, vermiculite alteration with accompanying illite-montmorillonite replacement of feldspars persists for only 160 or 330 h and is subsequently replaced by a chlorite + sericite assemblage. In addition to early chloritization of biotite, most granitewater experiments exhibit early formation of hematite and zeolites. Hematitization is observed in most runs, regardless of temperature, with hematite replacing either pyrite and/or magnetite grains usually associated with either biotite or hornblende. Hematite is also observed rimming chlorite grains, generally by 670 or 840 h of interaction. Zeolites are formed at all temperatures, usually on feldspars as replacements along twin or cleavage planes or as fine, radiating crystals on grain surfaces. The zeolites observed in the salt solution runs are tentatively classified as the Na-K variety which includes phases such as phillipsite and mordenite. Na- and Carich zeolites such as analcime and epistilbite, respectively, are more frequently encountered in pure water experiments. Although zeolite formation is common between 160 and 500 h in the experiments, sericitization and albitization prevail during the remaining weeks of interaction. However, formation of sericite on K-feldspar and plagioclase was observed in the first 500 h of interaction. Albite replacement of K-feldspar was not detected until roughly 500 h of rock-

Chl Zeal Hem Ser

t t

Chl Hem Zeol Ser

t least

most abundant

t

least Verm, Bio + Chl Feld + Ser Illite + Ser Chl - Hem + Ab ?Qtz

Py = Pyrite, Ksp = K-feldspar, Ser = Sericite Qtz = Quartz, Cc = Calcite, Ab = Albite, Verm = Vermiculite, Kaol = Kaolinite

450

D. R. Cole, H. Ohmoto, and G. K. Jacobs

fluid interaction and only in experiments where NaCl(aq) was present. Unlike zeolites, sericite forms irregular, patchy replacements embaying into and frequently cutting across feldspar grains. This style of replacement is most commonly observed where microcline checker-board twinning is being engulfed by sericite envelopes. Feldspar grains exhibiting disrupted twinning as a result of natural granulation typically display the most severe sericitization. Albite formed as the result of the Na exchange with K in K-feldspar exhibits a distinctive twinning referred to as “cheeseboard” twinning (MOOREHOUSE, 1959). The estimates of relative abundances of the alteration minerals summarized in Table 3 indicate that the greatest amounts of alteration occur in NaCl runs at 300°C followed by the 250 and 200°C experiments, respectively. The greatest amount of alteration is observed in static systems runs which employed the finest grain size. Chlorite formation dominates the alteration assemblages for both pure and salt solution experiments, followed by zeolites, sericite, hematite, and finally albite, in descending order of abundance. It is not uncommon for hematite abundances to exceed sericite in pure water runs, particularly at temperatures below 250°C. Sericite, however, becomes more abundant than both zeolites and hematite at 300°C for rocks reacted with 0.1 m NaCl solutions. In conjunction with mineralogical variations described above, the changes in the solution chemistry for a variety of experimental solutions reacted with granitic rock in the static system at 250°C are shown in Fig. 3. Differences in the absolute concentrations of K, Na, and Ca in the four representative examples shown are a function of rock type, temperature, surface area, and initial solution composition. Typically, solutions initially containing 0.1 m NaCl or KC1 exhibit greater shifts in chemical as well as isotopic compositions upon reaction with granite than pure water. For example, K leached from grgn at 250°C in the presence of a 0.1 m NaCl solution is approximately one order of magnitude higher in concentration than in the equivalent pure water run (mk+ = 4 X 10m3m vs. 4 X 10m4m at 840 h). A comparison of the K and Ca data from the 0.1 m NaClbqm (Fig. 3a) and -grgn (Fig. 3b) runs demonstrates the influence that rock type has on solution chemistry. Potassium is slightly more enriched in the solution that was interacted with bqm than the K concentration observed in the grgn experiment. Conversely, Ca is about two times greater in the solution reacted with the grgn than the Ca concentration in the bqm run. The presence of greater quantities of K-bearing phases in the bqm (K-feldspar plus biotite) and Ca-bearing phases (calcic-plagioclase plus carbonate) in the grgn probably best explains such contrasts in the solution chemistry. The most significant observation to be made from this figure (and data given in COLE, 1980) is that regardless of the absolute abundances of cations in solution, most of the concentration-time curves are similar in shape. They generally exhibit a rapid flux of elements into solution producing the initially steep curve between zero time and 160 h, followed by the attainment of near-steady state conditions after 160 h. A near-steady state behavior for K+ was also documented by SOLOMON( 1978 ) for microcline- 1 m NaCl experiments,

but over a shorter time interval ( - 120 h). The major exception to this is the 0.1 m KCl-bqm run (R26, Fig. 3c) which exhibits continuously decreasing concentrations of Ca and K after 160 h. The rapid increase in cation concentrations followed by steady state is in sharp contrast to the less rapid, continuous increase in the 6 I80 values of the solutions. Variations in the S’*O of Whole Rocks, Individual Minerals, and Solutions

Changes with time in the 6 I80 values of solutions and solids (rocks, minerals) as a result of water-rock interaction are summarized in Figs. 4 (whole rock) and 5 (mineral). A summary of the data is given in Appendix B. In these figures, isotopic data are represented as either depletions (-) or enrichments (+) in the I80 content from the starting material as a function of time with A, = 6’80; - s’*@,

A, = 6’80; - 6’80;

(1)

where w and s refer to solution and solid (rock, mineral), respectively, and i and f refer to initial (time = 0) and final (time = t), respectively. The depletions and enrichments represent mean values since solid and solution samples were usually analyzed at least twice. We considered only those experiments where the solution and rock samples collected at a given time gave 6 I80 values that satisfied the following mass balance equation b’80system= 6 180’ wxi w +

,3’80i~f

=

6’8OLXL

+

6’80fXf

(2)

where X is the atomic fraction of oxygen in the particular phase. The satisfaction of the mass balance equation indicates that the changes in the 6 I80 values of solutions and rocks were due to the isotope exchange between them and not due to loss of solution or due to formation of separate oxygenbearing phases (e.g., iron hydroxides on bomb walls) by reaction between solution and the vessel. Eflects of temperature and run duration From the data presented in Figs. 4 and 5, several similar isotope trends become obvious for both types of granitic rock, their minerals, and corresponding solutions. A continuous increase in the 6 “0 values as a function of time of the solution is accompanied by a continuous decrease in the 6 I80 values of rocks (and minerals). The maximum change in 6 “0 after 800 h was approximately 4%‘~for solutions, 2%0 for rocks, and over 10% for biotite. The magnitude of enrichments of I80 in the solutions usually exceeds the magnitude of depletions of I80 in rocks because these experiments are “rock dominated,” where the atomic oxygen content of the rocks is equal to or exceed by a factor of two that in solution. In general, the magnitude of the 6 “0 shifts observed in rocks, minerals and solutions increase with an increase in temperature and time, regardless of the experimental system used. It is important to point out that the effects of temperature on the magnitude of isotopic shifts become more distinguishable after approximately 300 h of reaction. With the exception of the 3OO”C-NaCl experiments, most isotopic data for solutions and rocks at roughly 160 h plot virtually on top of one another (see 160 h, Figs. 4 and 5). After this time,

Oxygen-isotope

exchange

between granites

451

and brine

6

2 200 400 600 8001000

0 12

12 10

5

10

5

._s

8

4

8

4

9 'b

6

3

6

3

’ -3

4

2

2

1

4 2

'9 1 'b

0

200 400 600 8001000

200 400 600 8001000

0

TIME(hrs)

TIME(hrs)

FIG. 3. Concentration of Na+, K+, Ca*+, Al’+, Si02, and ‘*O in solutions plotted against time of reaction for selected 250°C static system experiments (a) Run 25; (b) Run 24; (c) Run 26; (d) Run 13. Solution concentrations are in molality. The symbol X is used for 6’*0 change in solution, 6’*0: - 6’80L.

better separation in the data occurs such that obvious trends in the data may be observed.

Effects of solution composition and rock (mineral) type The solution the fluid-rock

composition isotopic

also has a pronounced

exchange.

Reaction

with

effect on a NaCl

so-

1n-relosparl 0

-G $

-1

cc -2 4 ioz

_2 :

1Plagioclasel

3

L '3 2 ii? 70

1 0

F 0

200

400

600

800

1Biotitel

1000 -14

FIG. 4. Changes in the oxygen isotope composition of granite gneiss (a) and biotite quartz monzonite (b) and solutions as a function of reaction time, temperature, and solution type. The i and f refer to initial and final (time = t), respectively, and r and w refer to rock and water, respectively. Symbols designate the following: square = grgn/static; circle = bqm/static; diamond = grgn/circulating; triangle = bqm/circulating; inverted triangle = bqm/disc. Open = pure water; filled = 0. I m NaCI; partially filled = 0.1 m KCl.

I

0

TIME(hrs)

I

200

I

I

400

I

I

600

I

I

800

I,

1000

TIME(hrs) FIG. 5. Changes in the oxygen isotopic composition of K-feldspar, plagioclase, and biotite as a function of reaction time, temperature, and solution type. See caption for Fig. 3 for definition of nomenclature and symbol usage. Circle containing X refers to microcline -1 m NaCl experiments. Also, open inverted triangle = bqm/disc/l m

NaCl.

452

D. R. Cole, H. Ohmoto, and G. K. Jacobs

lution results in a greater 6 ‘*O shift in the run product. For example, runs at 200°C in 0.1 m NaCl generally exhibit more isotopic shift than runs at 300°C in pure water. In addition, the type of salt solution has a marked effect on the magnitude of isotopic depletion in minerals. Plagioclase reacted with 0.1 m KC1 at 250°C (Fig. 5), for example, exhibits somewhat greater depletion in “0 (2%0) than plagioclase reacted with 0.1 m NaCl at 300°C ( 1.4%0) for roughly 840 h, suggesting that a similarity between the cation (Na+) in solution and the dominant cation in a particular mineral phase (Na-plagioclase) results in less enhanced rates of oxygen isotopic exchange. In general, the biotite and K-feldspar exhibit much greater 6”O depletions when reacted with NaCl compared to KC1 solutions (Fig. 5). These results are consistent with those obtained by O’NEIL and TAYLOR ( 1967). The type of the granitic rock used in the experiments also influences the isotopic trends. Isotopic shifts measured for bqm are usually slightly greater than those measured for grgn. This is particularly true for pure water runs. For example, bqm reacted with pure water at 300°C in the static system is depleted in “0 by 0.7%0at about 840 h whereas grgn reacted with pure water exhibits a depletion of only 0.3%0. Inspection of the magnitude of “0 depletions for various minerals reacted for more than 330 h indicates that biotite is most depleted, followed by K-feldspar, and finally plagioclase. Isotopic data show that neither quartz nor magnetite underwent measurable isotopic exchange in the 840-h period for either pure water or salt runs.

Eflects of grain size Large differences in grain size can produce different magnitudes of isotopic exchange. For example, the “0 depletion of a whole bqm disc reacted with I .O m NaCl for 4 weeks at 300°C (Fig. 4) is 0.5%0 (for outer - 1 mm), while depletion in finer grained bqm reacted with 0.1 m NaCl at 300°C is 1.7%0for a comparable time period. These data indicate that order-of-magnitude grain size differences between similar rock types should result in measurable differences in isotopic shifts in both fluids and rocks. Conversely, smaller differences in starting grain size (i.e., circulating vs. static systems) were apparently not enough to produce noticeably significant differences in the magnitudes of the isotope shifts. Grain size (surface area) effects also contribute to the magnitude of isotopic shifts observed in minerals. Coarsegrained microclines reacted with 1.O m NaCl solutions have I80 depletions comparable in magnitude to that observed in finer grained K-feldspars from granitic rocks reacted with only 0.1 m NaCl at similar temperatures and time. Surface minerals analyzed from two long-term disc runs, RlA2 and R4A 1, are generally less depleted in ‘*O than similar experiments employing higher surface area material reacted with less concentrated salt solutions (0.1 m NaCl). Isotopic analyses of the total-bulk disc, as well as the outer rind ( -0.5 to 1 mm depths into the surface using micro-drilling) of minerals indicate that virtually all of the isotopic exchange is localized in the surface layer ( 1 mm or less), with the interior of the disc generally unexchanged. Relation with degree of alteration

2.4

2.4

2.0 1.6 1.2

.I=

0

0.8

70

0.4

co

0

II

6

12

18

24

30

0

2.4

12

2.0

10

1.6

8

1.2

6

0.8

4

0.4

2

0

0

__‘-

%

ALTERATION

% ALTERATION

FIG. 6. The 6 “0 depletion ( -A) observed in granite and minerals plotted against volume percent of alteration in granite and minerals, determined from thin section point-count estimates, from selected experiments. The i and f refer to the initial and final isotope comp&ion of thejth phase: (a) whole rock, (b) K-feldspar, (c) plagio&se, (d) biotite. See Fig. 3 caption for definition of symbol usage. Data are for 300°C experiments unless noted otherwise.

A comparison of the alteration assemblages (Table 3) with the oxygen isotopic data suggests that whole rock-fluid isotope exchange observed in the 160 h of reaction was controlled by biotite transformation to chlorite t hematite and, to a lesser extent, zeolitization of feldspars. Continued formation of chlorite, zeolites, and hematite coupled with increase in sericite and albite formation after 500 h resulted in the continued increase with time in isotopic exchange. The magnitude of the “0 shifts observed in solutions and rocks appears to be a function of the type and proportion of new minerals formed, which in turn, are controlled by temperature, solution composition, time, and surface area of the solids. The similarity in type and paragenesis of alteration minerals formed in the finer grained experiments in the temperature range of 200 to 300°C suggests that variations in mineral abundances have had the most influence on the degree of oxygen isotopic exchange. In order to test this hypothesis of increased alteration producing increased isotopic exchange, point-count estimates were made of the extent each starting mineral was altered from a select group of pure and salt solution experiments. These estimates, given as volume percent, were determined for K-feldspar, plagioclase, and biotite, and were weighted to obtain an overall percentage for whole rock alteration. These values are compared with the 6 “0 shifts observed in the various minerals and whole rocks in Fig. 6. Our estimates indicate that after 800 h. K-feldspar, plagioclase, and biotite from bgm have been altered approximately 36, 13, and 94%, respectively, during interaction with

Oxygen-isotope exchange between granites and brine 0.1 m NaCl at 300°C. Similar phases from grgn reacted with 0.1 m NaCl at 300°C for over 800 h were altered roughly 30, 12, and 90%, respectively. At this condition, the feldspars and biotite exhibit 6 “0 depletions in excess of 2%0 and 8%0, respectively. At roughly 500 h of reaction, we observed percentages of alteration and magnitudes of “0 depletion slightly more than half of those determined for over 800 h of reaction. From mass balance calculations, the estimated magnitude of

whole rock alteration is as high as 25% for bgm at 3OO”C0.1 m NaCl-840 h, with an accompanying S’*O depletion of about 2%. Although the scatter is large and the patterns are not linear, data in Fig. 6 leave little doubt that 6 ‘*O shifts are produced by the degree (and type) of alteration of the starting phases.

453

c

ALBITE P

150

200

250

300

150

200

250

300

DISCUSSION Equilibrium Rock (Mineral)-Fluid

Isotope Fractionation

In order to make reasonable quantitative estimates of the rates of isotope exchange for either altering rock-fluid (r-w) exchange or altering mineral-fluid (m-w) exchange, we need to know the equilibrium fractionation factors, lO%(u, and 1031na,_,, respectively. In this context, we need to consider the isotopic effect accompanying the transformation of rock (mineral) to a new assemblage, not simply the exchange between rock and water. The magnitudes of these values depend on the temperature and the final proportions and types of new and residual minerals which, in turn, are controlled by the bulk chemistry of the rock (its mineralogy) and the fluid chemistry, including cation and anion concentrations and pH, and water/solid ratio. Because the final equilibrium mineral assemblage cannot be attained via the experiments described in this study (i.e., a final equilibrium state has not been reached), we modeled the equilibrium mass transfer between the granites and fluids used in our experiments to define the equilibrium state. The computer programs (EQ3NR and EQ6) used in this study are those developed by WOLERY (1978, 1979, 1983). EQ3 computes ( 1) distribution-of-species in the aqueous solution, (2) concentrations and activities of ions and complexes, and (3) degree of saturation of the mineral phases with respect to the solution. EQ6 is then used to calculate mass transfer and chemical equilibrium in aqueous solutionmineral systems. A closed-system model is directly applicable to our hydrothermal granite-fluid experiments including the circulating experiments. Reactant solid phases are titrated at specified relative rates into solution, and the chemical composition of the solution is progressively modified by dissolution of the reactant phases, as well as by precipitation and/or dissolution of all the various product phases, which remain in the system and continue to react. Reaction progress is computed until equilibrium is completely attained among all the solid phases and the evolving solution. Input for the closed-system model consists of the temperature, pressure, fluid compositions, initial pH, rock/mineral compositions and abundances, and fluid/solid mass ratios. An equilibrium mineral assemblage was calculated for each of the twenty-seven experimental runs described in Table 1. The thermodynamic data (DATA FILE

T(OC)

FIG. 7. Mole percent oxygen contribution to granite gneiss (a) or biotite quartz monzonite (b) based on “average” equilibrium mineral assemblages predicted by water-rock mass transfer modeling (EQ3 / 6) plotted against temperature of interaction. Run conditions summarized in Table 1 provided the input to 27 individual water-rock interaction computer simulations. The results at each temperature (i.e., 170,200,250,285, and 300°C) have been averaged to generate this plot. The mole percent oxygen contribution to each starting rock by unaltered minerals is shown on the right of each figure.

= DATAQ.3245 R46) for both the aqueous species and minerals are described in detail by WOLERY( 1983 ) . From the equilibrium assemblage, the contribution (in mol% oxygen) of each phase to the rock is computed. The results of these calculations are summarized in Fig. 7a,b which shows the mol% oxygen contribution of each phase to the rock plotted against temperature. The calculations reveal that the final mineral assemblage (type and abundance) is similar for a given rock type at a given temperature; thus, these plots represent the “average” for grgn and bqm. The presence of an electrolyte (NaCl or KCl) has a lesser influence on the final assemblage compared to temperature or rock type; and primarily controls the type and intensity of zeolitization. The similarity between predicted and observed mineral assemblages is quite good, with the exception of smectites which were predicted rather than chlorite. The calculated equilibrium mineralogy for the grgn remains nearly constant as a function of temperature with albite, zeolites (heulandite), quartz, K-feldspar, biotite, smectite (saponite), prehnite, and tremolite in decreasing order of abundance. The calculated phase behavior for altering bqm is more complex. It is clear from Fig. 7b that between 200 and 25O”C, zeolites (heulandite + clinoptilolite) and smectites (saponite, nontronite) dominate along with albite, K-feldspar, biotite, and quartz. At temperatures above and below this

454

D. R. Cole, H. Ohmoto, and G. K. Jacobs

range, epidote, muscovite, and smectites (beidellite, nontronite, and saponite) are stable phases along with albite, Kfeldspar, biotite, and quartz. The equilibrium oxygen isotope fractionation factors for altering rock-fluid (Fig. 8) were computed by combining the appropriate mineral-water equilibrium fractionation factors for each phase with their proportions recast in terms of the atomic oxygen fraction:

n

0

+ aqtz-wXqtz+

* * *.

O.lm NaCl

200

400

600

800

1000

(3) 0.6,

The following sources were used for the equilibrium mineralwater isotope fractionation factors: quartz, albite, and Kfeldspar (MATSUHISA et al., 1979); tremolite (MATTHEWS et al., 1983); muscovite (O’NEIL and TAYLOR, 1969); magnetite (BECKER and CLAYTON, 1976); zeolites (analcimewater as analog, KARLSSONand CLAYTON, 1990); smectite (KULLA, 1979); biotite (BOTTINGAand JAVOY, 1973, 1975); and epidote (zoistite as analog, MATTHEWSet al., 1983 ). In Fig. 8, the calculated oxygen isotope fractionation curve for altered grgn-fluid lies slightly above the curve for bqmfluid, and is nearly coincident with the albite-water curve of O’NEIL and TAYLOR ( 1967). Lower equilibrium fractionation factors are computed for bqm-fluid because of its higher mica content (i.e., biotite and minor muscovite). The bqmfluid curve is roughly coincident with the An3,,-water derived

,

,

,

,

,

,

,

,

,

,,

G

“._

0

200

400

600

800

1000

TIME(hrs) FIG. 9. Fraction of oxygen isotope exchange (F) plotted as -In( 1 - F) vs. time (h) for exchange between granite gneiss (a) or biotite quartz monzonite (b) and fluid (pure water; 0.1m NaCI) from static system experiments at 200,250, and 300°C. Lines on a first-order rate model (see text).

in Fig.

9 are based

V°C) 300

250

200

170

from MATSUHISAet al. ( 1979). The An30-water curve has been proposed by TAYLOR ( 197 1, 1974) as being a good analogue for equilibrium rock-water isotope exchange for silicit-rich compositions (e.g., granite, rhyolite), and our results confirm that this proposition is valid. The fractionation equations calculated for grgn-fluid and bqm-fluid are

11.0

10.0 ‘;; 8 3

3 = I n E v) -0

9.0

8.0

1000 In cum,_, = 2.68( 106/T2) - 3.57

(4)

1000 In abqrn+ = 2.67( 106/T2) - 4.02,

(5)

and 7.0

E ii? -ga

6*o

8 0 F

5.0

respectively. The error approximately +0.3%0.

about

these

two

curves

is

Degree of Isotopic Exchange and Reaction Order

The extent of isotopic equilibrium in these experimental granite-fluid systems can be quantified by the following

4.0

3.0 2.5

I

I

3.0

I

I

3.5

I

I1

4.0

I

4.5

I

I,

5.0

I

5.5

106iT2( K) FIG. 8. Equilibrium oxygen isotope fractionation factors between “altered” granite gneiss (Grgn-dashed line) or biotite quartz monzonite (Bqm-solid line) and solution plotted against 106/ T2 (K). Note that the dashed line for Grgn lies in nearly the same position as albite-water determined by O'NEIL and TAYLOR ( 1967). Mole percent oxygen contribution by each phase to the rock (has been determined from equilibrium mineral assemblages calculated by EQ3/ 6 mass transfer computer modeling. Square = grgn; circle = bqm; filled = 0. I m NaCl; open = pure water. The curves for quartz-water and Anjo-water are from MATSUHISA et al. ( 1979).

F==

_

a’ -

f

creq

where F is the fraction of isotopic exchange, cu’sare the isotopic fractionation factors between solid (rock or mineral) and fluid at time t(af), at t = O(oc’), and at t = ax(aeq). Using Eqn. (6) and the equilibrium fractionation factors described in Eqns. (4) and ( 5 ), we have computed the F values for whole rock-fluid runs. Examples of these data from the static experimental system are given in Fig. 9. The whole-rock data plotted in Fig. 9a (grgn), b (bqm) exhibit a systematic behavior of increasing F with an increase

4.55

Oxygen-isotope exchange between granites and brine in time. For these examples, where grain size and fluid/solid ratio are similar, the factors controlling the rate of change in F with time are temperature, fluid composition, and rock type. In general, at any one time, the F values are greatest at 3OO”C,followed by 250°C and then 200°C. Some ambiguity occurs for the shortest duration data (less than 175 h reaction time) where F values for all three temperatures are similar, between 0.1 and 0.03. The F values determined for the bqmfluid runs are somewhat greater compared to grgn-fluid experiments. Whole rocks reacted with 0.1 m NaCl exchanged far more than whole rocks reacted with pure water. At approximately 840 h, F values estimated for the pure water system range from 0.05 to 0.15 (200 to 300°C) compared to between 0.2 and 0.43 for the 0.1 m NaCl systems. It is noteworthy that the separation in data between 300 and 250°C is greater than between 250 and 200°C for these salt runs. The whole rock-pure water F values exhibit similar trends of increasing F with increasing temperature and time; however, the separation between lines for different temperatures is much less than observed in the salt runs. Finally, partial exchange data (not plotted) from mineral separates (K-feldspar, plagioclase, biotite) also show the same kind of trends as observed for the whole rocks. Before the rates of oxygen isotope exchange between solid and solution in our experiments can be estimated, the order of reaction must be determined. Because we have a mix of different types of data on whole rock-fluid and mineral-fluid isotope exchange (i.e., some as a function of time, some at one time only) we have used ORGLS, a general least-squares regression program (BUSING and LEW, 1962; BUSING, 1970) which utilizes all of the data to statistically evaluate rate models. We tested the following three general rate-order models: zero-order:

F g rot

first-order: ln( 1 - F) g -r,t second-order:

Fj( 1 - F) g r2t

(7) (8)

(9)

where r, the overall rate constant, is given as /&-E.IRr

(10)

with A0 and E, representing the pre-exponential term and the activation energy, respectively. Initial guesses are made of A0 and E, and the program computes new values for A0 and E, (from data on F, T, and time) and the agreement factor (Z), which is given as

(11) For unweighted data, the smaller the value for G, the better the fit is between the data and the model. Through the course of testing models, it was observed that the separate data sets representing each rock type or minerals from each rock type could be combined (i.e., the similarities in exchange behavior outweighed the differences). Thus, we regressed five groups of data: granite-O. 1 m NaCl; granite-pure water ; K-feldspar0.1 m NaCl; plagioclase-0.1 m NaCl; and biotite-0.1 m NaCl. Fits to the three general rate models utilizing all of the data (static + circulating + rocking autoclave) confirm that the first-order model best describes the bulk of these isotope

exchange data. With the exception of the granite-pure water data set, agreement factors for the first-order model range from 4 to 8 times smaller than 2 values predicted for either zero- or second-order models. For example, Z values of 0.139, 0.033, and 0.136 are estimated for zero-, first-, and secondorder for the granite-O. 1 m NaCl system compared to 0.022, 0.0 15, and 0.02 1 for these same reaction orders in the granitepure water system, respectively. For the mineral-O. 1 m NaCl systems, the i values are comparable to or less than the whole rock system-i.e., K-feldspar-O.1 m NaCl modeling yields 2 values of 0.139 (zero), 0.021 (first), and 0.133 (secondorder); plagioclase: 0.1 m NaCl yields 2 values of 0.074 (zero), 0.009 (first), and 0.073 (second), and biotite-0.1 m NaCl G values are 0.272 (zero), 0.077 (first) and 0.27 1 (second). The linear relationships of the granite-fluid data in -ln( 1 - F) vs. time space appear to be consistent, in part, with other hydrothermal mineral-fluid isotope exchange results. CRISSet al. ( 1987) note that for many mineral-fluid isotope exchange reactions (e.g., albite, quartz, and calcite-water), there appears to be a multiplicity of exchange processes influencing each experiment, and a single linear relationship between -ln( 1 - F) and time is not commonly observed. They conclude that different mechanisms of isotope exchange contribute to short and long times, but that each time frame conforms to simple kinetic theory-i.e., straight line segments defined by -ln( 1 - F). In the cases they cite, an inflection in slope demarcating rapid (steep) from slower (flatter) exchange in -ln( 1 - F) vs. time curves generally occurs at F values in excess of 0.4 to 0.5. The results from the granitefluid systems (whole rock, minerals) seem to fall only in the “short” time category, where F values typically do not exceed 0.4-0.5 (except biotite transformation to chlorite + hematite) after roughly 840 h of reaction. We believe that the steep “short” time trajectories are consistent with a mechanism of rapid grain regrowth or transformation reactions of the types observed in the granite-fluid experiments. Rate Constants and Activation Parameters for Oxygen Isotope Exchange

COLE et al. (1983) modified a simple isotope exchange rate model derived by NORTHROPand CLAYTON( 1966) to account for reactions occurring at surfaces. It is clear that isotopic exchange accompanying granite-fluid interaction is directly controlled by alteration of the rock. If we assume the rate-limiting step is addition and removal of oxygen atoms from the surfaces of minerals, then the forward rate constant rf can be expressed in Eqn. ( 12) with inclusion of a factor A representing the total constant surface area ( m2) of the solid (rock or mineral): rf= -ln(l

- F)(WS)/(W+S)(A)(t)

(12)

where ( W) and (S) represent the total moles of oxygen in the fluid and solid at Pand T, respectively, t is time in seconds, and rf is in units of moles of 0 me2 see-‘. The F values (measured or predicted) together with information on the total surface area, the mass of water and solid, and duration of each run can be used to compute the rate constants. As a first approximation, the surface areas for individual phases contained within either bqm or grgn were estimated by mul-

D. R. Cole, H. Ohmoto, and G. K. Jacobs

456

tiplying the total whole-rock surface area by the weight fraction of that phase. Examples of predicted F values from the first-order model are compared with appropriate measured F’s in Tables 4-8 for specific run conditions involving either bqm- or grgnfluid experiments. Included in these tables are values for WS/ ( W + S)A (moles 0 mm2), time (set) and r y, the forward rate constant based on measured F’s. These r? values represent either ( 1) average rates calculated from individual F and WS/( W + S)A values for each static run condition or (2) rates based on final measured F’s for experiments where the solid isotope composition was determined only at the end of a circulating or rocking-autoclave experiment. The modeling of first-order behavior described above indicates that exchange data from the three types of experiment apparatus are compatible and can be treated as a common group. The rate constants (moles 0 m-* set -‘) based on measured F’s (examples given in column 4, Tables 4-8) are presented in Tables 4-8 and are also shown on Arrhenius plots in Figs.

Table

4.

Run No. a 18 22 2T RlA2* 1RA3* R4Al

Summary of rate isotope exchange O.lm KC1 solution. Rock Type b

300 300 300 300 290 285

log r’; = a( 103/T) - b

F,

FP

C

d

rate and

constants O.lm NaCl

for oxygen solution or log L"f

("+?(A) e

$06) f

4

0.43 0.43 0.40 0.40 0.05 0.08 0.0017** (<0.0001)** (
0.08044 0.08383 0.31787 466.31 463.3 410.10

3.024 2.912 1.048 2.412 0.263 2.603

-7.82 -7.84 -7.59 -6.48 -7.05 -7.83

%P

250 250 250

0.25 0.29 0.21

0.25 0.24 -

0.07751 0.08022 0.07916

3.028 3.017 3.024

-8.13 -8.04 -8.21

21 23

Grgn BQm

200 200

0.23 0.23

0.13 0.12

0.07366 0.80440

3.006 3.010

-8.16

10

Bgm

170

0.02

0.02

0.21190

1.235

-8.46

24 25 26*

Grgn BP’

-8.19

aRefer to Table 1 for run conditions. bBcpn= biotite quartz monzonite; Grgn = granite gneiss. 'Measured fraction of oxygen isotope exchange, averaged from both solid and water data. dPredicted fraction order model (see

of text,

(13)

where a is the slope and b is the intercept. From the slopes oftheselines, a = (-EJR(2.3 X 103)), theactivationenergies (kcal mol-’ ) for oxygen isotope exchange controlled by alteration reactions are calculated. Table 9 summarizes the various reactions, equation for the regression lines, the regression coefficients, and the activation energies (E,) for oxygen isotope exchange. Inspection of rate constants plotted on Fig. 10 for granite0.1 m NaCl solution reactions suggests the data may be treated as two separate groups. Above 250°C the rate constants exhibit good linearity in 1/T space with an E, valueof 10.4 kcal mol -’ and a R 2 of 0.8 1. The values of log r? range from approximately -8.13 at 250°C to -7.75 at 300°C. The trend of these data is parallel, but offset from that of rates plotted at 200” (-8.18) and 170°C (-8.46). These lower temperature data yield an E, valueof approximately 9.1 kcal mol -’ .

parameters and between granite

T('C)

Grgn Grgn

10 and 11. Straight lines have been regressed through the data for each group (e.g., granite-O. 1 m NaCl) in the form

oxygen isotope eqn. 12).

exchange

from

a pseudo-first

eThe ratio of the product of the number of moles of oxygen in water (W) and solid (R for rock: M for mineral) to the sum of the moles of oxygen in the system (W+R) times the total surface area (A, m*). 'Run duration in seconds. gLog rate constant (moles of oxygen rn-'set-') estimated from F,. *Data not used in the non-linear least-squares regression. Run RlA2 is a lm NaCl experiment, lHA3 is at 0.5m NaCl, R26 is at O.lm KCl. **F, value adjusted to account for total disc: analysis made on outer surface equaling less than 0.5 percent of total rock volume.

Oxygen-isotope exchange between granites andbrine Table 5.

Summary of rate parameters and rate constants for oxygen isotope exchange between granite and pure water.

Run No.

Rock Type b

T("C)

12 17 1A 3c

Grgn B?P Grgn Bgm

13 16 1L

FL?

FP

log r", (W+J)S(A)

(x:06) f

C

d

e

300 300 300 300

0.07 0.14 0.05 0.06

0.14 0.15 0.03 0.05

0.08075 0.07528 0.20050 0.17290

2.437 3.024 1.307 1.674

-8.61 -8.42 -8.11 -8.20

Grgn Bgm Bqm

250 250 250

0.09 0.15 0.02

0.11 0.15 0.01

0.08015 0.06673 0.1893

3.002 3.622 0.439

-8.60 -8.52 -8.10

11 15 1M

Grgn Bqm Bqm

200 200 200

0.04 0.09 0.03

0.05 0.07 0.01

0.07993 0.06436 0.1895

3.019 3.024 1.274

-8.97 -8.70 -8.35

1c

Grgn

170

0.01

0.02

0.2395

1.598

-8.82

a

?3ee

g

footnotes for Table 4.

The apparent small differences in rate between 200” and 250°C were first noted in the -ln( 1 - F) vs. time plots (Fig. 9) where the separation between 200 and 250°C trends was less than between 250 and 300°C. Two possible explanations may be invoked to describe the apparent dichotomy in the rates. ( 1) We know that the uncertainties for the log r p values become larger at lower temperature, in large part because F values become small and the associated uncertainties for the F values become large.

Other errors can be attributable to estimation of total surface areas, oxygen isotope fractionation factors, and sampling, extraction, and isotope analysis. Errors in log r 7 of 50.2 to 0.3 at low temperature (~250°C) and between 0.1 and 0.2 at high temperature (>25O”C) are not unreasonable, given the magnitude of F values used to calculate the rates (Tables 4-8), as well as the other factors described above. The 170 and 200°C data, when the uncertainities are taken into account, may fall close to the extension of the line derived from

Table 6. Summary of rate parameters and rate constants for oxygen isotope exchange accompanying the alteration of K-feldspar by NaCl(agueous solutions). Run No. a

Rock Type b

18 22 2T RlA2*

Grgn Grgn Bqm

BW

300 300 300 300

R4Al

Bg-m

285

24 25

Grgn Bqm

250 250

0.20 0.34

0.27 0.33

0.1524 0.1494

1.8036 3.0168

-7.72 -7.69

21 23

Grgn Bqm

200 200

0.12 0.40

0.11 0.16

0.1473 0.1552

1.8288 3.0090

-7.99 -7.58

2B 10

Grgn =P

170 170

0.02 0.02

0.01 0.02

0.3772 0.3542

0.7344 1.2348

-7.96 -8.27

Micro Micro Micro

250 270 290

0.005 0.011 0.018

0.9436 0.9436 0.9436

0.4320 0.4572 0.4356

-7.95 -7.60 -7.41

RlA4 RlA7 RlA6 ?See

T("C)

E, C

0.55 0.50 0.11 0.01

log r:

F&J d 0.52 0.53 0.09

(W+;(A) e 0.1556 0.1454 0.4170 138.21

(0.0005)** 0.0002 481.1

$06) f

g

3.0240 2.9124 1.0480 2.4012

-7.38 -7.46 -7.32 -6.24

2.6028

-7.03

footnotes for Table 4.

?OOO In (I.~~. feld_W = 3.14 (106/Tz)- 4.9 used in the estimation of F,; based on predicted (EQ3/EQ6) model abundances of alteration phases. The calculated fractionation factors are very close to those given by Matsuhisa et al. (1979) for albite-water.

458

D. R. Cole, H. Ohmoto, and G. K. Jacobs Table 7. Summary of rate parameters and rate constants for oxygen isotope exchange accompanying the alteration of plagioclase by NaCl or KCl(aqueous solutions) Run No.

Rock Type b

T('C)

BQm Grgn Grgn

300 300 300

Bqm

285

24 26*

Grgn Bqm

250 250

0.19 0.53

0.20

25 21 23

BQm Grgn Bqm

200 200 200

0.06 0.08 0.11

2B 10

Grgn BQm

170 170

0.01 0.01

a 18 22 2T R4Al

"%ee %ee

K, C

0.16 0.26 0.05

FP

___ (W+;(A) e

log r", $06) f

g

1.8144 3.0240 1.0476

-7.84 -7.70

2.6028

-7.44

0.1214 0.1307

3.0276 3.0240

-8.08 -7.49

0.04 0.08 0.09

0.3282 0.1175 0.1325

1.5264 1.8288 1.7928

-7.88 -8.27 -8.07

0.01 0.01

0.3450 0.3225

0.7344 1.2348

-8.21 -8.35

d 0.17 0.25 0.04

0.1339 0.1428 0.3827

(0.0002)** 0.0001 467.6

-7.89

footnotes for Table 4. footnote for Table 6.

the high-temperature rates. (2) Alternatively, the apparent parallel offset may be due to the fact that the style (zeolite + hematite + chlorite), intensity, and kinetics of alteration differ less between 200 and 250°C than between 250°C and higher temperatures. Above 250°C we observed alteration

assemblages dominated more by muscovite, albite, and chlorite and less by zeolites and hematite, which are more common at 250°C and below. Errors propogated during data gathering and manipulation, and real differences in mineralogy both contribute to the

Table 8. Summary of rate parameters and rate constants for oxygen isotope exchange accompanying the alteration of biotite and NaCl or KCl(aqueous solutions). Run No.

T('C)

a

Rock Type b

18 22 2T R4A2*

Bqm Grgn Grgn Bqm

300 300 300 300

R4Al

Bqm

285

24 26*

Grgn Bqm

250 250

0.70 0.60

0.62

21 23 2A 25

Grgn Bqm Grgn Bqm

200 200 200 200

0.75 0.56 0.25 0.39

2B

Grgn

170

0.07

F, C

0.86 0.96 0.31 0.019

FP

(W+:(A) e

log P; $06) f

g

3.024 2.9124 1.0476 2.6028

-6.67 -6.80 -5.41

2.6028

-6.88

0.19884 0.18074

3.0276 3.0240

-7.09 -7.10

0.53 0.55 0.16 0.17

0.19629 0.18458 0.43181 0.40370

3.0060 3.0100 1.4940 1.5264

-7.04 -7.30 -7.08 -6.88

0.05

0.37090

0.7344

-7.44

d 0.98 0.99 0.30

(0.0007)**0.01

0.18496

0.19756 0.4417 485.05 478.14

-6.90

=-'Seefootnotes for Table 4. hlOOO In salt. bio_w = 1.35 (106/Tz)- 4.76 used in the estimation of F,; based on observed modal abundances of chlorite, hematite and unaltered biotite. Chlorite-water fractionation factors from Cole (1985) and hematite-water from Becker and Clayton (1976; magnetite as analog).

459

Oxygen-isotope exchange between granites and brine

-6.5,

38 5

V°C)

300 I

-7.0

-

-7.5

-

-8.0

-

-8.5

-

-9.0

-

250 I

170 I

200 I

I

b z

c E; _o

-9.5L

I 1.7

’

’ 1.8

’

’ 1.9

’

’

’

’

2.0

2.1

’

’ 2.2

’

’ 2.3

1

103iT(K) FIG. 10. Arrhenius plot of experimentally determined (based on measured F’s) oxygen isotope exchange rate constants, r 7, in units of moles of 0 m-* of solid surface area set-’ , for granite-fluid exchange. The symbols are the same as those described for Fig. 3. The solid lines are the least-squares fits to three separate data groupings (i.e., 250-300°C with 0.1 m NaCl; 170-200°C with 0.1 m NaCl; 170-300°C with pure water). See the text for an explanation of the

groupings.

magnitudes of rates plotted in Figs. 10 and 11. In the case of granite-O. 1 m NaCl, a line regressed through all the data yields an E, value of only 5.7 kcal mol-‘, which we consider unreasonably low for the kinds of reactions observed in these rocks. We favor an E, value of approximately 10.4 kcal mol-’ for the granite-O. 1 m NaCl system. The rate constants estimated for the granite-pure water system can be treated as one group. The best fit to these data yields an E, value of 7.4 kcal mol-’ for the temperature range of 200 to 3OO”C, with an R2 of 0.82 (Table 9). A rate constant estimated for 170°C (-8.82) is similar to the average calculated for 200°C (-8.84). Assuming a reasonable error of kO.2 log units could shift this value to coincide with the extension of 200-300°C rate data. Although the data are limited for mineral-fluid reactions, isotopic rate constants retrieved from the partial exchange results (Tables 6-8) also suggest the presence of two trends as observed for the whole rocks (Table 9; Fig. 11). This is expected because the whole-rock behavior is assumed to be an approximate average of the rates controlling individual mineral reactions. Regression of feldspar data from 250 to 300°C yields E, values of 8.2 and 7.4 kcal mol-’ for K-feldspar and plagioclase, respectively. An E, estimate of 11.0 kcal mol-’ is determined for biotite-0.1 m NaCl reaction (biotite altered to chlorite plus hematite). R2 values for these mineral -0.1 m NaCl reactions typically range from 0.75 to 0.92. Finally, an 18.3 kcal mol-’ value for E, is estimated for microcline-1 m NaCl, with an R2 of 0.98. Mechanisms of Oxygen Isotope Exchange and the Salt Effect

Two important mechanisms of oxygen (and hydrogen) isotope exchange between minerals and fluids, as summarized

by COLEet al. ( 1983, 1987), and COLEand OHMOTO( 1986), are diffusion of oxygen-bearing compounds between fluid and mineral, which may or may not involve a change in the mineralogy, and surface chemical reactions (e.g., dissolution/reprecipitation) that result in the formation of new minerals of the same or different chemical composition. Because of the relatively coarse-grained nature of the starting granites and the moderate to low temperatures of exchange, we can rule out the possibility that some of the measured oxygen isotope exchange is attributable to diffusion. To illustrate this point, let us consider diffusion of oxygen from a well-mixed solution of limited volume into microcline at 300°C. Using a diffusion coefficient of 2.44 X lo-” cm2 set-’ ( YUND and ANDERSON,1974), a grain radius of 0.05 cm (similar to the granite), and a corrected volume ratio (solution to solid) of 0.38, we estimate a time of 325 years to reach 10% isotope exchange (F = 0.1) assuming a spherical grain geometry. This difision exchange rate is over three orders of magnitude slower than what is observed in the granite experiments. Note that the diffusion coefficients for oxygen in other granite phases such as quartz or biotite (use phlogopite as analog) are several orders of magnitude less than those for microcline. Thus, all the discussion and interpretation presented earlier in this paper-i.e., calculations of F, aeq, r F values, are all justifiably based on the alteration-control model, rather than a diffusion model (COLE and OHMOTO, 1986 ). The influence of electrolyte solutions on the rates of isotope exchange observed in our experiments is confirmatory evidence that reactions involving chemical exchange at mineral surfaces (exposed or along grain boundaries) are promoting isotopic exchange. CHAI ( 1975) observed a similar circumstance where the rate of oxygen isotope exchange between

T(OC) 300 I

-6.0

250

200

170 I

-6.5 z. "E 5

-7.0

8

-7.5

E -

-8.0

EL -0 m

-8.5

f-9.01

’ 1.7

’

’ 1.8

’

’ 1.9

’

’ 2.0

’

’

2.1

’

’ 2.2

’

’

1

2.3

103/T(K) FIG. 11. Arrhenius plot of experimentally determined (based on measured F’s) oxygen isotope exchange rate constants, r t , in units of moles of 0 m-* of solid surface area see-’ , for mineral-fluid exchange. The meaning of the symbols is the same as described for Fig. 3; the only difference is that filled = K-feldspar; untilled = plagicclase; unfilled with cross = biotite; unfilled with X = microcline. Initial fluid compositions are either 0.1 m NaCl or 1.Om NaCl (microcline only). Least-squares regression lines have been fit to various data sets for specific temperature intervals. Table 9 gives the equations for these fits.

460

D. R. Cole, H. Ohmoto, and G. K. Jacobs Table

9.

Temperature for various

functions of the oxygen isotope exchange rate constant granite-fluid and mineral-fluid reactions.*

System

T"C Interval

Log r",=

R2

(a) Granite-O.lm Granite-O.lm Granite-H,0

NaCl NaCl

K-feldspar-O.lm

NaCl

Plagioclase-O.lm Plagioclase-O.lm Biotite-O.lm Biotite-O.lm

NaCl NaCl

NaCl NaCl

Microcline-O.lm

"Rate constant

NaCl

*Only

pure

250-300 170-200 200-300

-2.728(103/T) - 3.711 -1.979(103/T) - 3.993 -1.615(103/T) - 5.439

0.81 0.99 0.82

10.4 9.1 7.4

250-300

-1.908(103/T)

- 4.057

0.92

8.7

250-300 170-200

-1.619(103/T) - 4.986 -2.118(103/T) - 3.499

0.74 0.77

7.4 9.6

250-300 170-300

-2.392(103/T) -1.482(103/T)

- 2.532 - 4.167

0.70 0.83

11.0 6.8

250-290

-3.989(103/T)

- 0.302

0.98

18.3

based on F, in moles of 0 m-2sec-1.

bLinear least-squares

‘Activation

(b)

E, Kcal mol.' (c)

regression

coefficient.

energy. water

or

O.lm

(NaCL or KCl)

calcite and 2 m NaCl at 585”C, 2 kb was roughly 5 times

faster than the rate between calcite and pure water. Rates of dissolution of certain minerals are also known to increase with addition of NaCl or other salts to solution, as in the case of quartz ( JDOVEand CRERAR, 1990). In Fig. 12, we show a plot of log r vs. NaCl (aq ) concentration to illustrate how the magnitudes of the rates at 300°C are influenced by the solution composition. Although the data are limited, we observe that the log r values increase significantly with an increase in NaCl concentration. For example, in the granite-fluid system we observe an increase in log r from about -8.25 at NaCl = 0 to -6.75 at NaCl = 1.0 m at 300°C. We also observe that the individual mineral-fluid reactions (i.e., biotite to chlorite + hematite and K-feldspar to muscovite + albite * zeolites) exhibit parallel trajectories in log r vs. NaCl(aq) space, but different magnitudes in log r compared to the granite-fluid system. We have combined our result for microcline-1 m NaCl isotope exchange at 300°C with data taken from O’NEIL and TAYLOR ( 1967) for sanidine-NaCl ( HzO) at 300°C to generatea curve in Fig.12 for “microcline”-fluid isotope exchange. The calculated rates for “microline”- 1 and 3 m NaCl solution are significantly lower than those calculated for the granite-fluid system curves. The whole-rock, K-feldspar, and biotite curves were determined from complex systems involving a number of minerals reacting simultaneously to produce new alteration minerals. The microcline and perthite curves were determined from mineral-salt solution systems, where only cation exchange occurred. Above approximately 1.Om in NaCl (aq ) , the increase in rate accompanying cation exchange appears to be less pronounced. The activation energy for our microcline- 1 m NaCl ion exchange experiments,

are used in the regression

analysis.

18.3, is in good accord with that estimated by COLE et al. ( 1983) for sanidine-3 m NaCl, 17.5 kcal mol-‘. Additionally, the range of activation energies, 1O-20 kcal mol -’ , isappropriate for systems where reactions at surfaces are rate controlling ( BERNER, 1978 ). Rates of Isotope Exchange in Granite-Fluid

Systems

The rate constants, r I, obtainedfrom our experimental systemscan be used to compute the approximatetime requiredtoachieve agivenfraction ofexchange( A general expression relatingthe(W/S)moleratio,densityoftherock (p),averagenumber of moles of oxygen in the solid(X,), thegrainradius(u), thedegree of isotopeexchange(F),and time (t) was derived by COLEet al. ( 1983) for fluid interaction with a spherical grain in a closed system where t = -W 1 - ~)W/~)~s(~H~) 3[1 + (W/S)Jr;1( 10e4)

(14)

( 10m4isa factor used to convert cm2 to m*). The results of calculations that used this expression are given in Fig. 13 where the time to attain 90% isotopic exchange is plotted as a function ofgrain radii for 200 and 300°C and water/granite mass ratios of 0.05, 0.5, and 5. Figure 13 indicates, for example, that a granite system with a = 1 cm and (W/S),, = 0.5 requires approximately 6800 years to reach 90% oxygen isotope exchange at 200°C, but only about 1730 years at 300°C for interaction with pure water. A decrease in grain radius by one order of magnitude decreases the time needed for 90% equilibration by one order of magnitude (i.e., 680 and 173 years for 200 and 300°C granite-pure water, respectively).

Oxygen-isotope exchange between granites and brine

e

-5.0

-

-6.0

-

E

-7*o

z

E +

-8.0

-

-9.0

-

that allow flow through the rock. Reaction halos will form around fractures that transmit flow. These halos will grow with time and eventually all of the rock may be isotopically altered. At any one time the fraction of rock interacting chemically with the fluid will be reduced by a factor equal to the ratio of the thickness of the reaction zone to half the average separation between flow fractures ( CATHLES,1983 ) . Limits on the thickness of the reaction zone near a fracture (also known as skin depth, I’) can be evaluated from the following relationship taken from CATHLES( 1983)

%

‘0

461

g

Yn -0 I 0.0

I

I

I

I

I

0.5

I,

I

I

I

1.0

I

lib

I

I

2.0

I

I

I

I

(15)

I I1

3.0

NaCl(m) FIG. 12. Relationship between the oxygen isotope exchange rate constant (log r ‘;, moles of 0 me2 of solid surface area set-‘) for granite, K-feldspar, biotite, perthite, and microcline (sanidine) and the NaCl(aq) content of the starting fluid (O-3 m). Data from O’NEIL

and TAYLOR( 1967) have been used to estimate the log ry values for perthite and sanidine (3 m). With the exception of the 585°C perthite experiments all other data are for runs at 300°C.

where DE is the effective diffusional porosity of the matrix block; p. and p are the base density of water ( 1 g cm-3), and the density of water at temperature, respectively; 4 is the total rock porosity; A is the surface area per cm3 fluid-filled pore space; and k is the isotopic exchange rate constant, recast in units of cm see-’ . It is assumed for simplicity that pa/p = 1 and that

4A=6U

-4)

(16)

d The time to attain isotope equilibration is significantly decreased by the interaction of NaCl( aq) with granite. At 3OO”C, a= 1 cmand(W/S),,= 0.5, the time to reach 90% equilibration is approximately 530 years for granite-O. 1 m NaCl compared to 1730 years for granite-pure water. Increasing the NaCl(aq) concentration to 1 m, holding all other conditions the same, yields a calculated time for 90% exchange of only about 27 years (nearly a 20-fold increase over the granite-O. 1 m NaCl time). Note that a decrease from 5 to 0.05 in the water/granite mass ratio produces a one order of magnitude decrease in the time required to achieve 90% exchange. For example, at 300°C and a = 1 cm, the granite-O. 1 m NaCl system attains 90% exchange at about 1000 years for (W/S),, = 5, but only 100 years for ( W/S),,, = 0.05. The faster equilibration rate for low water/solid mass ratios is observed because the solid doesn’t have to shift as far in composition. In a simplistic way, a grain radius in excess of roughly 10 cm is analogous to considering granite in the form of a blocky matrix dissected by fractures. A system with large fracture apertures or numerous fractures would have a higher (WI S)mssratio compared to one with narrow or infrequent fractures. Conversely, a system with a grain radius less than l10 cm is analogous to a porous media (e.g., sediment, finegrained volcanic tuff, etc. ) . Therefore, one can view the results given in Fig. 13 as an illustration of the duration of isotope equilibration expected for granite grading from a fine-grained porous media (lower left portion) to a fractured matrix (upper right portion). It is important to point out that the results in Fig. 13 represent the minimum times required to attain 90% isotope exchange because we considered fluid contact with granite to be 100%. We have not considered the consequences of diffusional porosity in the solid away from a flow fracture. If the isotope reactions are rapid, the alteration reactions will not penetrate great distances into the rock from the fractures

where d is the grain size of the minerals in the rock (cm). For a grain size of 0.01 cm and a rock porosity of 0.02, we get a 4A value of 588 cm-‘. DE is typically 4% of the total porosity (0.02 in this case) times the diffusion constant of the reactant species in water, which we have assumed is roughly 2 X 10m5 cm2 set-’ . This yields a DE value of

5

4

3

2

1

,’

,/’

(W/S) = mass ratio

0 -2

2 l-0:

GrainORacd(cm)

FIG. 13. Time, in log years, to attain 90% oxygen isotope exchange vs. the grain radius of granite for granite-fluid interaction. The lines were calculated from Eqn. ( 14) described in the text for a closed system. The rate constants used in Eqn. ( 14) were taken from equations given in Table 9. The various lines represent the minimum times for the conditions under consideration because we have assumed 100% fluid contact with the grains.

462

D. R. Cole, H. Ohmoto, and G. K. Jacobs

approximately 1.6 X 10 --*cm*set -’ .Estimates ofk( cm set-’ ) can be obtained from relationships given in Fig. 13 by computing the time necessary to react a 1 cm radius grain to 90% isotopic equilibration for a specific temperature and (W/S),, ratio. This approach gives the following k values for (W/ S)“ass = 0.05: 4.5 X lo-l2 (200°C H,O); 9.5 x lo-” (300°C H,O); 3.2 X lo-” (3009C, 0.1 m NaCl); and 6.3 X 10m9 (300°C 1 m NaCI). Using these values along with the DE and &4 values described above yields skin depths of roughly 2.46, 0.53, 0.29, and 0.066 cm, respectively. An increase in (W/S),, by one order of magnitude produces an increase in these skin depths by 2.25. Additionally, an increase in the effective diffusional porosity of one order of magnitude causes an increase in the skin depth by a factor of approximately 3.15. The sharpness of oxygen isotope alteration halos will depend strongly on temperature, being quite sharply defined at temperatures above about 1OO’C (CATHLES, 1983). Rocks with fracture spacings of approximately 1 m will result in only a few percent of the total rock isotopically exchanging at any time. In such cases, the reaction zone will migrate toward the interior of the matrix blocks and the rate of isotopic exchange will decrease, because the reactants must diffuse through an altered halo with which they are in equilibrium to reach the zone of reaction (e.g., MURPHY et al., 1989; BLATTNERand LASSEY, 1989). This type of shrinking core reaction control will be addressed in the next communication which will couple fluid flow and isotope reaction kinetics. SUMMARY The systematics of oxygen isotope exchange between granite and water have been investigated experimentally. Isotopic redist~bution of oxygen during granite-fluid interaction is controlled by alteration reactions similar to those observed in natural systems (i.e., formation of chlorite, muscovite, smectite, albite, zeolites, and hematite). The abundance of these phases generally increased with an increase in temperature, time, and the salinity of the fluid. In reactions between granites of 6 I80 z + 8% and aqueous solutions of 6 I80 z -lO%o at T = 175 to 300°C depletion of ‘*O in the solid and enrichment of I80 in the fluid was observed at all conditions investigated, The magnitudes of isotopic shifts in the granite also increased with increasing temperature, time, and salinity and can be closely correlated with an increase in the intensity of alteration as well as style of alteration. In all cases, the final measured b”O of the rock and fluid do not appear to represent equilibrium values. This conclusion is based on a comparison of the final measured solidfluid isotope fractionation factors and fractionation factors estimated from equilibrium modal mineral abundances calculated by mass transfer modeling. This modeling generated bulk unite-water equilib~um fra~ionation factors that closely follow the Anso-water equilibrium curve. The fractions of oxygen isotope exchange(F) between granites and aqueous solutions range as high as 0.5 for the longest duration ( -840 h), highest temperature (300°C) 0.1 m NaCl-granite experiments. F values retrieved from mineralfluid subsystems range as high as 0.85 for isotope exchange accompanying biotite transformation to chlorite t hematite, and 0.55 for isotope exchange associated with K-feldspar al-

tered to muscovite + zeolite, both reacted for 840 h at 300°C with 0.1 m NaCl. The oxygen isotope rate data were observed to follow dosely with those predicted from a first-order rate model. Rate constants for oxygen isotope exchange between altered granite and pure water range from about 10 -9 to 10 +.’ mol 0 rn-’ sect’ for temperatures of 170 to 300°C. Rates for the granite-O. 1 m NaCl system range from 10 -8.5to 10 -’ 6 for this same temperature interval ( 170-300°C). Oxygen isotope exchange accompanying the alteration of K-feldspar and plagioclase exhibit similar rates for the temperature interval of 170 to 300°C namely lo-E.3 to 10-7.3, Oxygen isotope exchange associated with biotite alteration to chlorite yielded the fastest rates, 1O-7.4 to 10-6.5 between 170 and 300°C. Activation energies ranged from about 7 to 11 kcal mol-’ for the granite-~uid system including the individual mineral-fluid subsystems, and 18.3 for the microcline-1 m NaCl system. Using a simple closed-system model, we estimated the minimum time required for fluid-granite isotopic equilibration. For grain radii less than 1 cm (analogous to a porous media), we calculate a minimum time of 200 years or less for isotopic equilibration between granite and fluid (O-l m NaCl) at 300°C for fluid/rock mass ratios between 0.5 and 5. Coarser grain radii (analogous to matrix blocks) tend to equilibrate at 300°C in times ranging from 200 years to as much as lo4 to 10 5 years for block sizes between 0.1 and 1 m in radii at fluid Jrock mass ratios of 0.5 to 5, respectively. Acknon~le&nent.r-We wish to thank Drs. Dave Wesolowski and Jim Blencoe for comments and suggestions that greatly improved the manuscript. Thanks are extended to Cleve Solomon, W. F. Down< ,..___. J. D. Rimstidt, R. E. Erickson, and T. Giordano for maintaining the hydrothermal equipment and sampling. A special note of thanks is extended to Drs. H. L. Barnes, C. W. Burnham, and II. HolIand for their interest and advice on matters related to the experimental aspects of the study, and to Dr. J. G. Blencoe for assistance in generating the computer graphics. Reviews by Drs. Kurt Kyser and Allan Matthews greatly improved the manuscript and are gratefully acknowledged. Completion of this paper could not have been realized without the typing talents of Betty Benton and Regina Violet. This study was supported by grants from The National Science Foundation, EAR 7603724, EAR 80-07839, and EAR 8508379 to Ohmoto, and the Division of Enginee~ng and Geosciences, Office of Basic Energy Sciences (Cole), US Department of Energy, under contract DE-ACOS840R2 1400 with Martin Marietta Energy Systems, Inc. This manuscript has been authored by a contractor of the US Government under contract No. DE-AC05840R2 1400. Accordingly, the US Government retains a nonexclusive, royalty-free license to publish or reproduce the published form of this contribution, or allow others to do so, for US Government purposes. Editorial handiing: G. Faure

REFERENCES BARNESH. L. ( 1971) Investigations in hydrothermal sulfide systems. In Research Techn~ques~orHigh Presswe and High Tenl~eratures fed. G. C. ULMER), pp. 317-335. Sponger-Verlag. BARNESH. L., BURNHAMC. W., and Dowws W. F. ( 1979) Experimental evolution of chemical conditions in geothermal systems. Final Report on National Science Foundation Grant NumberAER 74-08473 by the Ore Deposit Res. Sect., The Pennsylvania State University. BECKERR. II. and CLAYTONR. N. ( 1976) Oxygen isotope study of a Precambrian banded iron fo~ation. Hamerskey Range, Western Australia. Geochim. Cosmochim. Acfa 40, 1153-l 166. BENJAMINT., CHARLESR., and VIDALER. ( 1983) Thermodynamic

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464

D. R. Cole, H. Ohmoto, and G. K. Jacobs

Appendix

A.

Chemical

composition

of granite

starting

materials

in weight percent

Grgn

wRa

Oxide

oxides.

Microcline

Kspb

Plagb

Biob

WRa

Kspb

Plagb

Biob

SiO2

69.5

69.5

68.7

67.1

35.21

62.5

64.2

60.3

37.2

65.19

A1203 TiO2

14.7

14.8

17.90

19.9

22.9

14.8

18.9

25.2

16.68

18.78

0.44

0.41

-

3.75

2.98

-

1.39

0.72

-

0.56

0.55

-

CaO

1.66

1.62

-

MnO

0.056

0.051

Na20

4.91

4.99

0.70

10.8

3.4

3.55

12.70

0.2

0.17

0.06

-

0.13

-

-

Fe203 Fe0

WC

K2O P2'5 H20-

(23.16)

0.86

2.12

5.82

(19.18)

2.90

2.0

11.96

1.92

-

2.76

0.07 11.76 0.05

5.9

0.089

8.50

3.50 3.65

0.09 16.0

0.03

0.51

0.08

8.2

0.11

0.37

0.4

9.33

15.79

0.25 0.24 3.52

H20+ (a) WR = Whole,

J. Blunchi,

(b) Electron microprobe analysis, G. Solomon (1978) Ksp = K-feldspar, Plag = Plagioclase, Bio = Biotite

analyst

Appendix B. Summary of average isotope compositions from granite and biotite quartz monzonite (bqm)-fluid experiments.

Sample Number

0.05

Time (hrs)

Mass Solution

Mass Solid

tsms)

(sms)

gneiss

(grgn

6: (O/00)

512-l 512-z 512-z s12-2 $12-5

171 342 508 677 a45

13.7 14.7 15.5 13.1 14.5

53.3 49.7 52.9 51.9 50.4

-10.6 -10.7 -10.6 -10.5 -10.4

7.7 7.6 7.6 7.6 7.5

s13-1 S13-2 513-3 513-4 $13-5

163 332 49% 665 834

13.1 14.6 14.8 15.6 13.4

51.3 52.4 52.6 52.4 52.9

-8.1 -8.0 -8.0 -8.0 -7.8

7.8 7.7 7.7 7.6 7.6

Sll-1 511-z Sll-3 Sll-4 511-5

166 333 504 667 839

13.6 14.9 13.5 14.4 13.3

50.9 50.6 51.2 52.1 52.0

-10.6 -10.4 -10.4 -10.5 -10.5

7.8 7.7 7.7 7.7 7.6

s22-1 s22-2 S22-3 S22-4 S22-5*

138 306 474 640 809

16.7 15.7 15.3 14.9 12.0

49.8 51.1 51.1 51.8 50.6

-11.5 -10.7 -10.0 -9.1 -8.2

7.5 7.2 6.8 6.4 6.0

7.4

7.5

2.1

8.9

6.3

7.3

-0.5

8.8

5.7

6.7

-9.0

8.8

524-l 524-2 S24-3 524-4 S24-5*

168 334 501 671 843.

12.4 12.5 13.6 13.2 13.3

49.1 50.7 50.0 50.1 50.3

-10.6 -10.3 -10.0 -9.3 -8.9

7.6 7.4 7.2 7.1 6.7

7.5

-1.0

9.0

s21-1* SZl-3* S21-3* 521-4 S21-5

166 332 508 667 835

13.1 12.7 12.2 12.1 11.1

53.0 50.8 49.6 50.1 49.4

-9.5 -9.1 -8.8 -8.4 -7.9

7.7 7.7 7.5 7.4 7.0

8.9

8.8

7.8

7.0

7.4

-3.9

8.9

6.7

7.2

-5.1

8.8

7.8

7.6

3.1

8.9

7.6

7.6

0.5

8.9

7.2

7.5

-3.1

8.7

Oxygen-isotope exchangebetweengranitesandbrine Appendix

B. (Continued) f

Sample Number

Time (hrs)

Mass Solution (gms)

C2T-13

291

666.0

706.3

-10.2

7.4

7.0

ClA-10

363

495.7

1366.6

-10.0

7.6

7.8

636.6

1319.4

-11.0

7.2

7.7

579.2

1360.0

-10.0

7.6

7.9

647.0

1262.3

-11.2

7.4

7.8

1277.2

-10.0

7.6

C2A-12* ClB-13 C2B-14*

415 559 204

465

Mass Solid (gms)

&)

7.3

-4.3

9.0 8.9

0.4

8.9

8.8 1.4

a.9

ClC-16

444

640.7

517-l S17-2 517-3 517-4 517-5

165 333 501 671 840

14.1 13.0 12.6 12.7 12.0

51.5 50.7 50.7 49.7 51.1

-10.2 -9.9 -9.6 -9.4 -9.1

8.3 8.3 8.1 8.0 7.9

S16-1 516-2 S16-3 S16-4 S16-5

165 504 672 843 1006

11.1 10.4 10.0 11.3 11.7

51.7 51.0 51.7 51.4 49.8

-9.5 -9.1 -8.9 -8.7 -8.6

8.4 a.3 8.2 8.1 8.0

515-l 515-2 s15-3 s15-4 s15-5

162 330 498 670 834

10.0 10.7 10.7 10.2 10.0

49.4 49.1 54.3 50.3 51.8

-9.8 -9.6 -9.5 -9.4 -9.3

8.4 8.3 8.3 8.2 8.2

S18-1 S18-2 S18-3 S18-4 S18-5

168 336 504 672 840

14.4 13.7 14.1 14.0 16.5

51.2 50.9 50.6 52.4 54.2

-9.5 -8.2 -7.5 -7.0 -6.4

8.0 7.5 7.1 6.9 6.5

S25-1 S25-2 525-3 S25-4 S25-5

168 334 501 671 841

14.7 15.0 14.6 14.1 13.6

52.6 52.2 51.6 51.2 50.3

-10.0 -9.6 -8.9 -8.5 -7.7

8.1 8.0 7.6 7.4 7.0

7.8

9.9

7.5

9.8

6.8

9.8

S23-1 S23-2 S23-3* S23-4 S23-5

162 330 498 666 836

13.6 14.1 14.2 14.6 15.5

51.2 50.8 52.3 51.2 51.5

-9.9 -9.5 -9.1 -8.7 -8.5

8.0 a.2 8.0 7.8 7.6

8.0

8.2

3.4

9.8

7.7

8.1

1.6

9.9

7.1

7.8

-1.1

9.9

S26-1 S26-3 S26-5

168 504 840

14.0 14.4 14.1

49.6 52.2 53.6

-9.8 -9.2 -8.2

8.4 8.1 7.7

7.9 6.9 6.4

-1.1 -2.6

9.9 9.8 9.9

c3c-14

465

435.1

1485.3

-10.4

8.1

ClL-10

122

520.2

1525.0

-9.9

8.4

C25-15

424

607.1

1399.5

-8.2

7.8

8.0

8.1

0.4

9.9

ClM-9

354

511.9

1497.7

-9.9

8.4

RlAZ-6

667

667.3

149.8

-11.0

8.2

7.4

8.0

1.3

9.8

R4Al-4

723

162.3

85.0

-10.9

a.4

8.0

8.1

2.9

9.8

73

883.6

149.8

-10.1

8.2

lHA3

9.9

8.5

9.9 9.8

8.4

7.5

a.2

2.9

9.8

7.0

7.9

-0.2

9.8

6.1

7.7

-4.5

9.8

D. R.Cole,H.Ohmoto,and G. K.Jacobs

466

Appendix B. (Continued) Sample Time Number (hrs)

RlA6 RlA7 RlA4

121 127 120

Mass Solution (gms) 600.0 600.0 600.0

Mass Solid (gms) 10.0 10.0 10.0

f ($0)

-8.8 -8.8 -8.8

f ($0)

I ($0)

9.7 9.8 9.8

(*biotite value has been corrected for the presence of minor quartz)