Isotopic exchange in mineral-fluid systems. II. Oxygen and hydrogen isotopic investigation of the experimental basalt-seawater system

Gtwchhrm~ca PI Cosmorhrmrca Acln Vol. 51. pp. 1523-1538 C Pergamon Journals Ltd. 1987. Pnnted in U.S.A.

0016.7037/87/S3.00

+ .@I

Isotopic exchange in mineral-fluid systems. II. Oxygen and hydrogen isotopic investigation of the experimental basalt-seawater system* DAVID R.COLE',MICHAELJ.MO-~~L'~~~

HIROSHI OHMOTO~

‘Geosciences Group, Chemistry Division, Oak Ridge National Laboratory, Oak Ridge, TN 3783 I. U.S.A. 2Hawaii Institute of Geophysics, University of Hawaii, Honolulu, HI 96822, U.S.A. ‘Department of Geosciences, Pennsylvania State University, University Park, PA 16802, U.S.A. (Received March 12, 1986; accepted in revisedform

March 5, 1987)

Abstract-Oxygen and hydrogen isotopic exchange reactions between basalt and seawater at T = 300” to 500°C were investigated, using oceanic tholeiitic basalt (&I80 = -5.70/w; dD = --70%), natural Seawater (8’*0 = 1.2’%0;dD = +50/w)and artificial seawater (B”O = -5.3%) as starting materials. The starting basalts varied in the crystallinity (from holocrystalline to glass) but were ground to approximately the same grain size (-100 mesh). The water/rock mass ratios ranged from I to 3, and the duration of the experiments ranged from 167 to 576 days. In general, depletion of “0 and enrichment of D in basalts occur at all temperatures, with the magnitudes of change being greater as temperature, time, and to a lesser degree, glass content increases. The trends in isotopic shifts are directly related to changes in the style and intensity of mineralogic alterations in the basalt (e.g., smectite at 300°C. talc-actinolite at 400”-500°C). The changes in the b’*O values of basalts and seawater in the experimental systems were observed to follow closely with those expected from a first-order rate law. Rate constants for the oxygen isotopic exchange between rock and water range from lO-95 to lO-8.omoles oxygen/m* of solid surface/set for temperatures of 300” to 500°C. The activation energy for the isotopic exchange reaction was calculated to be I I .5 Kcal/ mol. An application of our experimental rate data to natural systems suggests that the oxygen isotope equilibrium between basalt and seawater in the mid-oceanic ridge may take place within approximately 1000 years at 350°C.

Our experimental data also suggest the equilibrium oxygen isotopic fractionation factors between the altered basalt and seawater to be 3.5 + 0.5L at 3OO”C, 2.0 + 0.4%~ at 4OO”C, and 0.5 + 0.25% at 500°C. The observed hydrogen isotopic fractionation factors between the altered basalts and seawater in our experimental systems were about -74700 at 300°C. about -62%~ at 4OO”C, and about -48% at 500°C. These fractionation factors are probably attributable to the equilibrium fractionation between Fe-rich secondary phases and seawater. INTRODUCTION AND HYDROGEN isotopic compositions of rocks and fluids from natural hydrothermal systems have been used to quantify the history of fluid-rock interaction, such as the attempts to determine the origin of the water and the water/rock ratios of the system (e.g., TAYLOR, 1974; TAYLOR, 1979; FORESTER and TAYLOR, 1980). Typically, calculations of water/rock ratios and the origin of water have been made with an equilibrium model which is based on two critical assumptions. The first is that rock and fluid readily attain isotopic equilibrium at some specific temperature below which no further isotopic exchange can occur. The second is that the equilibrium isotopic fractionation factors between rock and fluid can be closely approximated by those for particular mineral-water systems, such as plagioclase-water for the oxygen system and chlorite-water for the hydrogen system. As useful as the equilibrium approach has been, it is not without its limitations. Rock and fluid in many geothermal and ore deposit systems do not appear to be in isotopic equilibrium even at temperatures in exOXYGEN

* Research sponsored by the Division of Engineering and Geosciences, Office of Basic Energy Sciences, U.S. Department of Energy, under contract DE-AC05-840R2 1400 with Martin Marietta Energy Systems, Inc.

cess of 300°C. whereas other similar natural systems are apparently in isotopic equilibrium even at temperatures as low as about 100°C (e.g., CASADEVALL and OHMOTO. 1977; COLE, 1980, 1983; GREEN et al., 1983). If the equilibrium model does not hold, the calculated water/rock ratios and the interpretation of the origin of fluids may differ significantly from reality. The selection of an equilibrium model in most previous studies was made because of the lack in our knowledge of the mechanisms, the rates, and the equilibrium fractionation factors for isotopic exchange between rock and fluid. When all of these are determined for a variety of fluid-rock systems and under a variety of geologic conditions, the isotopic composition of rock and fluid can be used as a quantitative tool in deciphering the geochemical history of natural hydrothermal systems. This paper is the second in a series of papers directed at increasing our understanding of the mechanisms, the rates, and the equilibrium fractionation factors for isotopic exchange reactions between rock and fluid under a variety of conditions. In our first paper (COLE et al., 1983), we examined published data from oxygen isotope exchange experiments between minerals and aqueous solutions, and computed the rate constants and activation energies for the isotopic exchange reactions for two important isotope exchange mechanisms. The first mechanism involves the formation of new minerals by surface re-

1523

1.524

D. R. Cole, M. J, Mottl and H. Ohmoto

actions (e.g., dissolution-reprecipitation. cation exchange, hydrolysis, etc.) between residual (reactant) minerals and solutions. The new minerals may have the same chemical and mineralogical compositions as the residual minerals (e.g., the overgrowth of new quartz on quartz crystals) or they may be entirely different (e.g., formation of K-feldspar from Na-feldspar, muscovite from paragonite). The second mechanism is the diffusional transfer of oxygen- (and/or hydrogen-) bearing species between mineral and fluid. We have demonstrated that the rates of isotopic exchange through surface reactions are several orders of magnitude greater than those through diffusion at temperatures below about 500°C. An important suggestion made in our first paper was that the rate of isotopic exchange between a mineral and fluid depends strongly on the ease of formation of new minerals in the system, which in turn depends on the fluid composition, the chemical, mineralogical and physical properties of the solid (e.g., porosity, permeability, grain size), the fluid/rock ratio, temperature. and pressure. If a fluid and rock are grossly out of chemical equilibrium before interaction, significant amounts of isotopic exchange may take place within a relatively short period of time. However, once chemical equilibration has been attained between the fluid and rock and no new minerals are formed, isotopic exchange will occur at much slower rates. This paper presents the results of our experimental study on the oxygen and hydrogen isotope exchange between basalt and seawater reacted at temperatures ranging from 300” to 500°C. The main objectives of this study were: (1) to determine the equilibrium isotopic fractionation factors between basalt and seawater. (2) to determine the rate constants and activation energy for the isotope exchange reactions between them, and (3) to compare these values with those for simple mineral-water systems. The third paper in this series (COLE and OHMOTO, in preparation) will address chemical and isotopic exchange in the experimental system granite-HzO-NaCl-KCl. EXPERIMENTAL METHODS Isotopic exchange was investigated in a total of 3 I basaltseawater experiments conducted by MOTTL (I 976). Details regardingthe experimental design and procedures used in these experiments are given by Morr~ (1976) and Morr~ and HOLLAND(1978). Only a brief summary is provided here. Table 1summarizes the experimental conditions used in this study. Experimental conditionsand startingmaterrals Oceanic tholeiitic basalts were reacted with either natural (i.e.. Sargasso seawater, Atlantic seawater) or artificial (i.e., no Mg, SO:-, HCOS) seawater for various durations at temperatures and pressures ranging from 300” to 500°C and 0.6 to I .Okbar, respectively. The experiments of longest duration lasted 576 days and the shortest for only I67 days. Experiments were conducted in standard large-volume split-furnace style hydrothermal apparatus. Each experimental charge consisted of weighed amounts of rock powder and solution, with argon as an inert filler gas, inside a 10 cm long gold capsule which

averaged 15ccin volume. The water to rock mass rams rangetr from one to three. The starting basalts were comprised of the foliuwmy four types: (I) holocrystalline basalt with no glass. dredged from the Blanc0 Trough, near the intersection of the Blanco Fraaure Zone with the Juan de Fuca Ridge, consisting of plagioclase. pyroxene, quartz, opaques, and a few percent deuteric alteration product; (2) inner parts of basalt fragments dredged from the median valley of the Juan de Fuca Ridge, containing about 80 percent crystals (plagiociase microlites. a few 1 ‘4mm plagioclase phenocrysts) and about 20 percent glass: (31the outer part of the basalt fragments with 40 percent glass: and (4) the outermost part consisting of nearly 100 percent. glass. AI1 samples were ground to a grain size that averaged about I Xi pm. Nitrogen and Krypton B.E.T. adsorption studies of these ground materials indicate that small but measurable differences in surface areas exist among the different starting materials The overall range in specific surface area (A,) is from 0.268 to 0.41 m2 g-‘. The 6”O values of the starting basalts are similar, i.e.. 5.6 rt 0.2%. 6D values range from --66 to 77% (see footnote. Table I). The natural seawaters (i.e.. Sargasso, Atlantic) used m most experiments are similar in their major, minor, and trace element chemistry, but differ slightly in the 6’*0 values. 1.2wwr~~ 0.8%0. An artificial Mg-SOi--HCOj free seawater was used in some experiments by MOTTLand HOLLAND(1978) in an attempt to reverse the direction of flux of certain elements. This artificial seawater had a 6’*0 value of --5.3%~ (see footnote, Table I ). Isotopicanalysrs Oxygen for isotopic analysis of startmg waters was hberated by reaction with BrF, in nickel reaction vessels at 231; _+-25°C after a method modified from O’NEILand EPSTUN (1966). Rock samples (reactants. products) were reacted with BrF: in nickel reaction vessels at 600°C to liberate oxygen as described by CLAYTONand MAYEDA(1963). Conversion of O? to CO: was carried out quantitatively by reaction with hot, disc-shaped graphite electrodes, at temperatures low enough to preclude the formation of carbon monoxide (<6OO’C). and CO, was analyzed for oxygen isotopes. Water was liberated from the basalts by heatmg them in platinum buckets under vacuum to 1100”-1200°C in an induction furnace and stored in breakseals. Any hydrogen formed by reaction between Hz0 and Pt was reconverted to HI0 by reaction with hot CuO (450’C). The total Hz0 was then converted to H2 by reaction with uranium at I’ ; 600°C in a manner similar to that described by FRIEDMAW and SMI-TH ( 1958) and GODFREY(1962). The uranium pump technique of FRIEDMANand HARDCASTLE(1970) was used for transferring the hydrogen gas from the preparation line lo the mass spectrometer. Isotopic analyses were performed on either a 3-inch radius (D/H) or 6-inch radius (‘*0/‘60) dual inlet isotoDe ratio mass spectrometer. Measurements are reported in the usual “6” notation in per mil.

where R is either the ‘8O/‘6O or D/H ratio, and the subscripts .Yand std refer to the sample and standard, respectively. The analyses are reported relative to the Standard Mean Ocean Water (SMOW; CRAIG,1961). Overall precision of analyses was about +O. 1 per mil for 6% and + I per mit for 6D values.

RESULTS AND INTERPRETATION The measured 6”O and 6D values of hydrothermally-reacted basalts are presented in Table 1. The final isotopic values of run solutions were not measured

it

A,,

576

1

9.0‘

9.0‘

(12.1,

0.502

0.2‘9

4.9

3.3

2A

400

x-t

SS”

272

1

7.50

7.50

5.9

0.116

0.207

0.833

om25

4.5

1.3

28

‘W

9,

SS”

272

I

7.50

7.M

5.6

0.‘

16

0.207

0.833

0.0525

‘.2

‘.t

2c

‘00

eg

SS”

272

t

7.50

7.50

7.7

0.‘

16

0.207

0.855

0.002%

1.4

‘.5

tfl

xw

ssw

212

2

8.25

‘.I3

4.8

0.458

O.IM

0.916

(2,

I.6

2.3

-63

(1,

6

55”

272

3

8.5’

2.8‘

7.2

0.172

0.078

0.9‘%

0.ooo911 1.‘

3.9

,I,

1.‘

2.5

-52

,I,

6

ssv

272

1

7.%0

7.50

4.8

O.‘M

0.207

0.833

o.w25

‘.2

0,

I.8

2.‘

-56

(I)

8

1.7

2.9

P6300

20

‘W

2E

‘00

2F

‘00

xhl

O.OOI‘

1.‘

3.9

II,

0.9

4.‘

a>

1.1

2s‘

(3,

s.7

2.1

-54

tt)

(I

t.5

3.0

-,2

tt,

9

‘00

xt xt

Sk

$7‘

I

7.50

7.50

0.116

0.207

50

‘00

XT

h-t

II‘

I

7.50

7.50

0.416

0.201

11.0

0.8

(3)

x:

‘00

(a

1,‘

!

,.%(I

7.50

0.416

0.201

I.3

1.7

(2,

50

‘cm

iu

SSW .4-t

17‘

1

0.1116

0.201

10.1)

2.8

(2,

-3.9

6.7

-3.e

6.2

5A

7.1101.50

‘.5

4.5

‘.L

(2)

-2.9 1.6

-53

(2,

8

3.7 3.1

5s

‘00

9’

kt

1,‘

t

1s

1.50

0.116

0.20,

to.7

2.‘

(2,

,A

500

XT

ssw

260

,

6.39

6.19

1.0

O.S55

0.17.S

0.709

O.WJ9

‘.5

3.0

(2)

2.5

0.5

-‘3

(2)

*

50

500

91

ss4J

268

I

6.39

6.39

6.7

0.355

0.176

0.709

0.0039

4.2

2.9

(3,

2.‘

0.5

-10

(I,

8

3c

500

x+a

SSW

268

I

6.39

6.39

1.2

0.355

0.176

4.4

3.2

(2,

2.‘

0.1)

directly, but calculated from mass balance equations recast in the following form:

where X refers to the mole fraction of either oxygen or hydrogen, i andfrefer to initial and final, respectively, and W and R signify water and rock, respectively. The final atomic fractions of oxygen and hydrogen for rock and water have been calculated taking into consideration the increase in water content of the basalts due to the formation of hydrous minerals (Column 9, Table 1). Because ofthe method used (chloride balance), these water contents represent an upper limit. This increase in water content of the basalts produces

a 10 to 20 percent increase in the atomic oxygen fraction and up to a IO-fold increase in the atomic hydrogen fraction of the basalt. In cases where the water contents of the final basalt could not be estimated (runs 5A-5E), the final atomic fractions are assumed to be equal to the initial values. The final isotopic compositions calculated for the water-added and water-absent case using the same data set (e.g., experiments 1,4,2, 3, 6) differ by approximateIy 1.2 to 1.8 per mil for hydrogen and generally less than 0.4 per mil for oxygen. These differences are not great enough to exclude the use of data from experiments 5A-5E in estimating equilibrium fractionation factors or rates of exchange. Oxygen isotopic exchange in the experimental basalt-seawater system Apparent fractionation factors. The measured fractionation factors between basalt and seawater, Ak.w

D. R. Cole. M. J. Mottl and H. Ohmoto

1526

and &w, are shown in Fig. 1, as a function of temperature, run duration and the crystallinity of the starting basalt. These data exhibit the following general trends:

(I) Decreases in & from L& are observed in all runs. The magnitude of these decreases ranges from values as lo\s as 0. I %I(e.g.. runs 4A, 4C, IA) to values over 10% (e.g., runs hB, 6C). Note that the initial oxygen isotopic compositions of the fluids were lighter than the starting basalt. (7) There is a pronounced temperature effect, where the magnitudes of (AL-W- L&,.~)increase with increasing temperature when the Ak.w, duration and crystallinity are the same. For example, holocrystalline basalt reacted with artificial seawater for 172 days at 300°C (run 4D) exhibits a 0.7% change in 6’*0. and over 4% fractionation for nearly the same conditions at 400°C (run 5D). (3) The differences between Ak.w and &V increase slightly m magnitude with an increase in the run duration, when the &.w, temperature and crystallinity are the same. This trend is observed, however, only for experiments conducted at 300” and 400°C. At XXX, similar A$+ values are observed regardless of the run duration for similar experimental starting conditions (runs 3A VS.6A). (4) The effects of crystallinity on the observed lsotoplc fractionations can be significant depending on the experimental conditions. In general, there is a tendency for the mag nitude of change in isotopic fractionation to be less for experiments where the basalt is holocrystalline. This trend 1s best demonstrated by data given in Fig. I for holocrystalline basalts reacted with isotopically lighter artificial seawater (runs 4D, 5D, 6D). Commonly, samples with 20 to 40 percent glass mixed with crystalline basalt exhibit isotopic shifts comparable to and sometimes greater than the 100 percent glass samples (e.g., compare runs 4B and 4E, 5B and 5E, 2D and 2B).

(74 d --f--T

4OO’C

I-

s6

& a

5

2364

-

272 d

4

FIG. 1.Changes in the oxygen isotopic fractionation factors between basalt and seawater as a function of temperature. run duration, and crystallinity of the starting basalt. -: starting LW value: triangles: -170 days: circles: 236-272 days; squares: 576 days; open symbols: holocrystalline: A, 0 and q: crystalline; 0: crystals plus glass; solid symbols: glass. Aw is the equilibrium isotope fractionation factor

(5) There ISan effect due to the magmtude rn :he tsorop~ difference between unreacted basalt and seawater. Larger changes in &.w - &&are observed in runs where isotopicallg light artificial seawater was used. I e’. larger -lk.% values. Clearly. these experiments stan out farther from isotopt~ equilibrium than those with smaller IL.,., value\ Essentially all previous investigators have interpreted their experimental data on isotopic exchange reactions, either between solids and fluids (heterogeneous hysterns) or between aqueous species (homogeneous systems), in terms of the NORTHROP-CLAYIW ( 19661 method (e.g., O’NEIL and TAYLOR, 1967. 1963; O’Nelr PI ul.. 1969: CLAYTON et al.. 1972). The relationship among Ak.w, A&_w, Azw (equilibrium isotope fractionation between rock and water). and F (the fraction of isotopic exchange) is given b5

which can be rearranged

as:

We have adopted the same approach and represented the data given in Table 1 on Northrop-Clayton typr diagrams (Figs. 2a-c). All of the basalt-seawater data plot in the upper lefi quadrant of the Northrop-Clayton diagrams because the starting fluid in each experiment is isotopically lighter in “0 than the basalt. The data should fall be-, tween the ordinate (0% exchange) and a lint with a slope of - 1 (100% exchange). If our experimental basalt-seawater systems have a unique Azw value at :S given temperature, all the trends through the data points should converge to a point on the ),-axis, which defines the Azw value. The lines in Figs. 2a-c are used IO show the range of possible F values for GUI-data-I ti, they are given as reference lines with which to estimate the F values for our data. These lines should not be interpreted as regressed trends through the data. Figures ?a-c demonstrate that the basalt-seawater data can be interpreted as sharing approximately the same fractionation factor, Azw, at a specific temperature Additionally, we can constrain the Azw values by as.. suming that Azw is less than or equal to the minimum J&-W measured at a particular temperature and greater than the minimum &,.w measured at the next highest temperature (see Fig. 1). The reader should note that we have not corrected the fractionations for salt-effects because our intent is to use. these data exclusively for understanding natural basalt-seawater systems. ‘The work of TRLJESDELL (1974) and GRAHAM and SHEPPARD ( 1980) can be used to make estimates of the basalt-pure water isotope fractionation for oxygen and hydrogen, respectively. Data given in Figs. 1 and 2a suggest that the equiiibrium fractionation factor, AZ,, at 500°C is approximately 0.5 rt 0.25%. Because we have 111)data al

Basalt-seawater isotopic exchange

FIG. 2. Estimation of the basalt-seawater oxygen isotope fractionation factors by partial exchange at 500°C (a), 400°C (b) and 300°C (c). Y-axis = Ak.w (or 1000 In Ak.w). A& is the equilibrium isotope fractionation factor. The symbols are the same as those described for Fig. 1.

higher temperatures to bracket these results, the lower limit may be as low as -0.4%0 if runs 6A and 6B ( 167 days, xt) had exactly the same F value. At 400°C (Fig. 2b), we estimate a value of 2.0 + 0.47~ for Azw. A line connecting the holocrystalline data (runs 5C and 5D, 174 days) gives a Azw value of about 1.6%0, a possible lower limit at 400°C. An upper limit of approximately 2.7%0 is given from data on crystalline basalts (runs SA and 5B, 174 days). This value is probably high because it exceeds the minimum ~$a_~value of 2.4%0. The 300°C data are more difficult to interpret because of the limited amount of exchange that took place. We have used the following criteria to estimate the equilibrium fractionation factor: (a) A& should be less than 4. I%o, the minimum &w value; (b) AZ, should be greater than 2.3%0, the minimum &w value at 400°C; (c) the F value for run 4D (holocrystalline at 300°C) should be less than the F value of 5D (holocrystalline at 4OO’C); and (d) the F value of runs P4 and P6 (576 days, gl and xf, respectively) should be greater than the F value for run 4D ( I72 days). Based on these constraints, we estimate that Azw at 300°C is approximately 3.5 + 0.5%0. Degree ofisotopic exchange. Using Eqn. (4) and the “average” equilibrium fractionation factors described above, we have computed the F values for each run. These data are given in Table 2 and plotted in Figs. 3a-c. The observed ranges in F values at 300”, 400” and 500°C are 0.1 to 0.36, 0.47 to 0.88, and 0.56 to over 0.95, respectively. The variations in F values at a given temperature can be attributed to, in part, the crystallinity of the starting basalt. At 300°C (Fig. 3c) and run durations of 236 days, F varies from about 0.1 for holocrystalline basalts to 0.3 for samples initially containing glass (20 to 100 percent). The 400°C (Fig.

3b) data indicate that for durations of 272 days, F values range from about 0.6 for holocrystalline basalts to over 0.85 for basalts with glass. Holocrystalline basalts reacted at 500°C for 268 days exhibit F values of approximately 0.8, whereas glass-bearing basalts reacted under similar conditions yield F values exceeding 0.95. It appears that at 500°C equilibrium was established after 268 days for most samples with glass, but at 3OO”C, only about 35 percent isotopic exchange was observed even after 576 days for basalts containing glass. Before rates of oxygen isotope exchange can be estimated, the order of the reaction must be determined. Because we have so few data at any given P-T condition as a function of time, the conventional concentration (or fraction of exchange) versus time plots are not satisfactory for estimating rates. Instead, we have used a non-linear least-squares regression method (BUSING and LEVY, 1962; BUSING, 1970) which utilizes all the data. We tested the following three rate-order models: zero-order approximated by F = k,$, ln( 1 - F) = -k,t for first-order, and F/( 1 - F) = k2t for second-order, where k, the overall rate, was approximated by AOe-EJRT.The program computed values for A0 and E, from data on F, T and time, and the agreement factor (or variance) which is derived from the square root of the sum of the observed minus the predicted F (see Table 2) values divided by the degrees of freedom. The smaller the value for the agreement factor, the better the fit. These values range from 0.1844 for zero-order, 0.0784 for first order and 0.0897 for secondorder. Larger standard errors for the coefficients were caiculated from the second-order model-i.e., A0 equals 0.078 -I 0.113 for second-order compared to 7.33 X IO-’ + 6.95 X IO-’ for first-order, an improvement

D. R. Cole. M. J. Mottl and H. Ohmoto

1528

r -= -In(l ---F)(WR) ~(W + R)f.,t

of over 30 percent. The errors calculated for & are considerably less, where E, equals 1.78 x 10’ t 1.83 X IO3for second-order versus I .36 x IO’ I I .3 1 x 10’ for first-order. Additionally, plots of observed minus calculated F’s (or residuals) against calculated F’s for both models (not shown) reveal distinct clusters for the second-order case. Clustering or other trends such as linear groupings of data, with significant deviation from zero by the residuals, is indicative of a less adequate model. Taken together. the lines of evidence described above suggest that the preferred rate model is first-order. Rate constants and activation parameters.hr isotopic~ exchange. The F values (measured and predicted. Table 2) together with the information on the total surface area (A in m*), the mass of water and rock. duration of the experiment (t in set), and the following rate equation for the oxygen isotopic exchange reaction between solid and fluid (COLE ef a/.. 1983) allow us to compute the rate constants, r:

Table

2.

Summary isotopic

log r = -2.52( lO’!r) .- 4.80

and rate conatante basalt and seawater.

._._______.-._-.

for

oxygen

.-___

-

fsec)

Log

__

. ..- .

WR Run

No.

F

T -

T -

300°C

400°C

5OO’C

c er to >f kle.wured %redlcted text).

cd)

A

Gf+R)

o.io

IA ID 1E

L

P

_~_____c...__C_.__e___

__B_.___

T -

F

?(

:J . 3 :I 0.22

0.21 0.21 0.17

f-69 3.56 2.01

0.176 0.176 0.127

f

-__----A_..

2.039 2.039 2.039

x x x

“H

.

107 107 107

-9.b: -9.o3 -9.1: -9.06 --9.47

IF 4A

o.22 0.10

o.15 0.16

1.17 3.69

0.084 0.175

2.039 !.486

x x

107 107

48

0.29

0.17

I.h9

0.167

1.486

x

10’

-8.98

hi 41)

0.13 o. :‘i

0. I! 0.12

2.44 1.44

0.172 0.169

I.486 1.486

x x

107 107

--9. -9. 30

4E

0.24

0.15

3.31

0.172

1.486

x

107

-9.03

P4 P6

0.36 ___~-_-~ 0.36

0.40 0.43

3.37 3.69

0.184

4.977 4.977

x x

107 107

-9.30 -9.35 ---.-

24 28

0.76 0.82

0.81 0.79

j.Ob 2.80

0.145 0.145

2.35 2.35

x x

107 107

-a.>4 -8.42

ZC*

0.58

2.96

0.147

2.35

x

10’

-8.73

20 ZE 2F

0.a8 0.79 o.a1*

0.74 0.71 o.82

.b3 I.12 .i)6

1 j

0.097 0.075 0.144

2.35 2.35 2.35

x x x

10’ 107 10’

-a. 27 -8.35 -a .40

5A

0.64

0.68

3.Ob

0.138

1.503

x

107

-8.50

iB 5C

0.81 0.52

0.69 0.53

3.06 2.01

0.138 0.138

1.503 1.503

x x

107 107

-8.30 -a.67

5D

0.47

0.52

7.01

0.138

1.503

x

107

-8.53

SE

0.52

0.65

2.80

0.138

1.503

x

107

-8.61

--.

3A 38 3C 30 3P* 6~ 6B 6C 60’

0.99 0.99 0.92 0.95 0.77 0.99 0.89 0.90 0.50

0.99 0.99 0.95 0.96 0.99 0.94 0.97 -

2.61 2.38 2.52 0.93 1.71 2.61 2.bl 2.38 1.71

0.118 0.124 0.118 0.087 0.118 0.124 0.121 0.123 0.118

2.316 2.316 2.316 2.316 2.316 1.443 1.443 1.443 1.443

x x x x x x x x x

10’ 107 107 107 10’ 107 10’ 10’ 107

-a.bs -7.99 -8.29 -7.92 -8.33 -7.82 -8.15 -8.10 -8.40

-.-

Table fraction fraction

---.

30

1 for

run conditions. of oxygen iaotope of oxygen isotope

exchange. exchange

from

first-order

model

P oral surface aree of solid run reactant. eThe ratio of the product of the number of frock (R) to the sum of the moles of oxygen

moles of oxygen in the system.

in

Run duration in seconds. BLog rate cwi(lt(Lnt (males *Data not used in non-linear

estimted regression.

messured

of

oxygen/m’/sec) least-squnres

I0 /

where 7‘ is in “K. This equation grves smoothed log I values for 300”. 400” and 500°C’ of --9.20. x.54 and -8.06, respectively. The smoothed log r \siues calculated from the predicted F’s (Column 3, Table 2) differ very little from these--i P., --9.25. 5.47 and

of rate parameters exchange between

_-___.~-_--_..

--.l-_l_

where W and R are the number of moles ot axkgen In water and rock, respectively. The r values based on measured F”s (Column 2. Table 2) are presented In Table 2 and are also shown on an Arrhenius plot in Fig. 4. The rate constants (moles oxygen/m’ of solid surface/set) calculated for 300”, 400” and 500°C c‘xhibit the following ranges (log vahres), respcctivelv: .9.6 to -8.98. -8.7 to -8.3, and -8.4 to 7 3 Ltxsrsquares regression of these data yields the thilowmg temperature function:

from

water

(see

(W)

F.

and

Basalt-seawater isotopic exchange TIME (dovsl

0.6 0.6 09 0.95

0.99

0.6

-4

iT

06 F

& -2 0.9

c -3

0.95

-4

(b)

o-.,

a

0.99

HX

*-

GL

0.6

-4 -

3OOT 06

-2

-

-3

F

-4

L

09 0.95

(cl

I 100

I 200

I

I

I

I

300

400

500

600

0.99

TIME (days)

FIG. 3. Fraction of oxygen isotopic exchange (F) plotted as In (1 - F) versus time (days) for oxygen isotopic exchange between basalt and seawater at 500°C (a), 400°C (b) and 300°C (c). The symbols are the same as those described for Fig. 1. HX refers to holocrystalline, and GL refers to glass.

-8.0 for 300”, 400’ and 5OO”C, respectively. Because of the large variations in r observed at each temperature, as well as the large uncertainty in the equilibrium fractionation factors, it is not unreasonable to assign an error of +0.5 log units to these data. An activation energy of 11.5 Kcal/mol is calculated from the slope of this regression line, compared to a value of 13.6 kcal/mol estimated from r values based on predicted F’s (non-linear least-squares method). Inspection of Fig. 4 indicates that the oxygen isotopic rate constants for the basalt-seawater system are similar in magnitude to those estimated by COLE et al. ( 1983) for several mineral-water isotopic exchange reactions. These include the conversion of paragonite to muscovite in the presence of KC1 solutions, the conversion of kaolinite to pyrophyllite and diaspore in the presence of NaCl solutions, and the reaction of albite with either Hz0 or NaCl-bearing solutions. Activation energies estimated for these isotopic exchange reactions range from 10 to 14.8 to 2 1 Kcal/mol for the kaolinite, paragonite and albite-solution systems, respectively. This range of activation energies, lo-20 Kcal/mol, is appropriate for systems where reactions at the surface are rate-controlling. Mechanisms of oxygen isotopic exchange reactions. Two important mechanisms of oxygen isotope exchange between minerals and fluids, as summarized

1529

by COLE et al. ( 1983) and COLE and OHMOTO ( 1986) are: diffusion of oxygen-bearing compounds between fluid and mineral, which may br 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 fine-grained nature of the starting basalts and the high temperatures of exchange, we cannot rule out the possibility that some of the measured oxygen isotope exchange is attributable to diffusion, particularly for experiments where new mineral growth is sluggish. COLE et al. ( 1983) and COLE and OHMOTO( 1986) have presented the diffusion coefficients and rate constants for oxygen isotopic exchange reactions for various mineral-fluid systems. Figure 5 compares the F values computed from a diffusion model with the observed F values. The diffusion coefficients are based on those for the anorthite-water system of GILETTI et al. ( 1978). and the equations necessary for the computations are described by COLE et al. (1983). We modeled diffusion at the exact conditions used in the basalt-seawater experiments to estimate the fractional approach to equilibrium (F). Data in Fig. 5 demonstrate that the fractions of isotopic exchange observed in our experiments are much larger than those accounted for from the diffusion model. Note that the diffusion coefficients for other possible analog candidates, such as pyroxene (GILETTI and PARMENTIER, 1978) or phlogopite (GILETTI and ANDERSON, 1975), are several orders of magnitude less than those for anorthite-water. The observed changes in the oxygen isotopes between basalt and seawater in our experiments can. however. be accounted for by the surface reaction model. Several lines of evidence support this contention:

( I ) There was extensive formation of alteration minerals in the run products (MOTTLand HOLLAND, 1978).The alteration assemblage consisted OEsmectite, albite, analcime or wairakite, quartz and anhydrite at 300°C; smectite, albite or oligoclase, anhydrite, tremolite-actinolite, and quartz at 400°C; and smectite, oligoclase or andesine, tremolite-actinolite. talc and minor quartz at 500°C. Swelling chlorite similar to that forming at Reykjanes, Iceland, was present in a few experiments, (e.g., IA, 3A) but always in smaller amounts than smectite. Because of the tine-grained nature of the run products. MOTTL and HOLLAND (1978) could not quantify the abundances of alteration phases. However, the extent of alteration can be estimated from the water content of the rocks after reaction and the identity of the hydrous phases that contribute the water. The intensity of alteration ranges from approximately 5 to 25 percent at 3OO”C, 30 to 60 percent at 4OO”C,and 60 to over 90 percent at 500°C. Differences in the crystallinity of the starting basalts have produced most of the variations in the intensity of alteration at a given temperature. For example, at 3OO”C, glass-rich starting basalts are generally more altered than the crystalrich basalts, and the dominant phase formed at the expense of the glass is smectite. Tremolite-actinolite formed much more abundantly at 500°C than at 400°C and, although present at all crystallinities, it is more plentiful as an alteration product of glass. Talc, which formed only at 5OO”C,could not be detected by x-ray in the glassy whole-rock powders. In the crys-

D. R. Cole. M. J, Mottl and H, Ohmolo

1530

r i t

1 I

I /

f

t

I__.__“_

!

1.3

i.2

1 (5

(4

I

I

1.6

t7

--i *8 --i----r’..

i&TlK)

FE. 4. Arrhenius plot of experimentally determined (based on measured I;‘s) oxygen isotope rate constan!.~. r, in units of moles of oxygen/m2 of solid surface/set. The symbols are the same as those described for Fig 1.The solid curve is the least-squares fit to the basalt-seawater data. Curves given by COLE ei al. ( 1983) fo: paragonite-KCI, albite-Hz0 and kaoIinite-NaCl isotope exchange experiments are shown for comparison. taline and hoi~~~line powders. however. it is generally more abundant than tremolite-actinolite. Clearly, the magnitudes of the oxygen (and hydrogen) isotopic exchange between basalt and seawater in our experiments

07

-I

P

0 03

'A A

i _+

3OO’C

Q2

01

0

~ , 0.1

03

0.2

CALCULATED F DIFFUSION

MODEI.

FIG. 5. Relationship between the measured fractions(F) of oxygen isotope exchange in the basalt-seawater system and the F values calculated from a diffusion model. This model used diffusion coefficient data from an analog system, anorthite-water (GILETII efal.,1978), and the same run conditions described in Table 1. The symbols are the same as those described for Fig. 1.

are controlled by the style and intensity of secondary mmerat formation. Greater “0 shifts in less altered, crystalline basalts compared to more altered, glassy basalts can occur only when secondary phases forming at the expense of crystalline material preferentially concentrate 160. Chlorite, tremulite-actino!ite and talc all tend to concentrate I60 compared to the smectites. (7) As noted previously, there is a close similarity in the rate constants and the activation energy for oxygen isotope exchange in the experimental basalt-seawater systems and isotope exchange accompanying hydrolysis or cation exchange reactions in mineral-fluid systems (see Fig. 4). The activation energy of I 1.5Kc&/mol determined for the basalt-seawater system is simiiar to those given for chemical reactions involving fluids and mineral surfaces {termed “surface controlled” bt BERNER, 1978). (3) There is also a very close sim~I~ty in the equll~br~~m rock-water isotope fractionation factors obtained in our experimental study with those calculated for “theoreticai” altered basalt-seawater systems. MO~L (1983) and REED (1983) computed the proportions of alteration minerals for systems where basalt and seawater react at different water/rock ratios, and when the basalt is completely altered (see Table 3). The theoretical Szw values given in Table 3 were computed b) combining the appropriate mineral-water equilibrium fra+ tionation factors for each alteration phase (see Appendix A) with their proportions, recast in terms of the atomic oxygen (or hydrogen) fraction. There are some differences between the alteration mineralogy observed in the experimental systems and the phase assemblages given in Table 3. The most noticeable discrepancy mvolves the formation of smectites {~minandy t~~h~mi~ in the experimental system versus chlorite and epidote in natural systems and the assemblages described in Table 3. However. because the AWvalues for smectite-Hz0 are stmilar to those for epidote-HzO, and chlorite comprises less than 20 weight percent ofthe theoretical alteration assemblage at waterj rock ratios less than 10, the Azw values given for water/rock mass ratios of I to 5 in Table 3 may be taken as the “‘theoretical” Azw values for our experimental systems (W/R .T I 103). These theoretical values are compared with our measured $W values in Fig. 6. Considering the fact there e&ts some

1531

Basalt-seawater isotopic exchange uncertainty (kO.2 to 0.5%~) in the individual mineral-water

in the hydrogen content of the rock took place from the starting basalt (Hz0 - 0.5 wt.%) to the reaction products (Hz0 - 1 to over 10 wt.%) due to the formation of secondary hydrous phases. In order to calculate F, r and AZ, values, the solid phase must have maintained approximately the same amount of H (or 0). This prerequisite is fairly well maintained for the Hydrogen isotopic “exchange” in the experimental oxygen system, but not for the hydrogen system. basalt-seawater system It is clear from our data that the changes in 6D of the basalt are directly related to the addition of water The measured hydrogen differences between basalt to the rock as the result of the formation of alteration and seawater in our experiments are shown in Fig. 7. minerals. However, we are not certain whether the dD From an initial AR.w value of about -750/w, the A&-w values of the new minerals were controlled by kinetic values become less negative with increasing temperaisotope effects, or by the equilibrium isotope fiactionture, i.e., approximately -69 to -61%0 at 400°C and ation factors between seawater and the new minerals. -51 to -357~ at 500°C. We observed only a few %O Typically, isotope fractionation due to a kinetic isotope increase in A4.w at 300°C. effect during unidirectional chemical reactions shows The hydrogen isotope data are insufficient to detera preferential enrichment of the lighter isotope in the mine the rates of isotopic exchange. More importantly, reaction products. Therefore, in the basalt-seawater the process of hydrogen isotope transfer between basalt system such a process would tend to deplete the rock and seawater in our experimental systems is not “iso- in D (produce more negative 6D values) which is the tope exchange”. This is because a significant change opposite of what is observed in these experiments.

fractionation factors, we may conclude that the agreement between the observed AZ, values and the theoretical A& values is excellent. There is also modest agreement between our experimental fractionations and those generated by BOWERS and TAYLOR (1985) from computer modeling of mass transfer reactions between basalt and seawater.

Table 3. Estimated mineral abundances (in weight percent) for altered basalt and calculated bulk rock-water equilibrium isotopic fractionation factors.

(W/R) mass

I

B

5

5

10

10

100

II

b

B

b

13.0 38.0 9.0

68.0

36.0

24.8

23.0

3.5 3.7

13.0

il

b

19.7 21.1 20.9 29.3

30.0 13.0

26.1 21.7 16.3 23.3

4.9

15.0 6.0 0.6

a.5

97

PhaseC Chlorite (Chl) Eoidote (ED) Aibite (abi Actinolite (Act) Tremolite (Trem) Quartz (Qts) magnetite We) gematite (Hem) Llmenite (Ilm) Sphene (Sph) Paragonite (Parag) Phlogopite (Phlog) Kaolinite (l&01) Anhydrite (Anhy) ninnesotaite Wnn)

A”P-ii qd

moot 3OO~C 4o0°c 5OO’C

4-w= 300°C 200°C 400 *c 500°C

16.8 20.4 23.1 36.2

3.5

4.1

1.2 23.0 a.0

4.1 9.0 7.0

6.0 11.0 16.0

20.0

7.62 3.64 1.50 0.40 -41.5 -39.8 -36.4 -30.4

7.49 3.54 1.49 0.32 -43.0 -40.7 -37.0 -30.5

1.17 3.55 1.50 0.34 -20.1 -32.8 -31.0 -28.1

%i”eral abundance estimated by Mottl (1983) using a calculation (see text for details). S(i”ere.1 abundances estimated by Reed (1983) using B %quatio”s for the equlibrium isotopic frsctionation dees CB” be found in Appendix A. Calculated equilibrium oxygen isotopic fractionation altered baealt and seawater. ‘Calculated equilibrium hydrogen isotope fractionation altered basalt and seawater.

7.14 3.28 I. 28 0.14 -45.1 -41.8 -37.1 -30.4

7.30 3.11 1.15 0.11 -39.4 -41.2 -31.4 -30.7

CIPU-noraative

4.87 1.80 0.08 -0.91 -52.0 -44.5 -39.3 -30.0

6.32 2.92 0.99 -0.17 -47.8 -39.3 -33.7 -25.0

type

paes transfer nodel. factors for these phafactors factors

between between

D. R. Cole. M. J. Mottl and H. Ohmoto

1532

Hydrogen isotope equilibration between the secondary hydrous phases and seawater is perhaps a more plausible explanation for the behavior depicted in Fig. 7. Because hydrogen isotope exchange rates are relatively rapid (GRAHAM, 198 l), we believe that isotopic equilibration of the new hydrous minerals probably did occur, particularly for the longer duration, higher temperature experiments. To test this hypothesis, we have calculated the hydrogen isotope fractionation between basalt and seawater as different amounts of new hydrous minerals were added to the rock in a closed system. In these calculations, we assumed an initial 6D’ of -72k for the basalt, and an equilibrium value of 5’30 for the water. The results are presented in Fig. 8 where ARmwis plotted against weight percent of alteration. Note that the mineral-water fractionation factors used in these calculations were not adjusted for the salt-effect-thus, the differences between the calculated curves and our measured fractionation data represent minimums (see GRAHAM and SHEPPARD, 1980, for details regarding the salt effect and hydrogen isotope exchange). The results indicate that there is generally poor agreement between the calculated and measured iso-

-iO

r

-20

-30

-40 + & -50 D E :: -60 0 -70

-80

-90

100 L

___L-_____2

__

. _ .___:

5 106/T'iKJ

500 9

400

T(‘CI 300

I

I

200

I

I

\

/

8

FIG,. 7. Measured hydrogen isotope fractlonauon factors between basalt and seawater in our experiment& systems. compared against mineral-water and rock-water curves drscribed in the text. The symbols are the same as those described m Fig. I. The curve is described by the equation: IO00 InBsti -17,38(10’/T) - 15.37(106/r), where Tis in K

7

6 i

2

I

0 1 -4

I/

I

I

2

3

4

1

j06/T2fK)

FIG. 6. Equilibrium oxygen isotope fractionation factors between basalt and seawater in our experimental systems (dots with error bars), compared against mineral-water and theoretical rock-water curves. See Appendix A and Table 3 for the source of these data, respectively.

tope fmctionations. We observe only modest agrc%ment in the case of actinolite at 500°C. In the case ofzoisite. the calculated rock-water fractionation trends pass through or lie very close to the measured fractionation data. However, unlike smectite, actinolite. talc or chlorite, zoisite (or other epidote-group minerals) was not identified in any run products. In order for the equilibrium hypothesis to have ment. the measured fractionations must be controlled by a phase or phases whose isotopic fractionation mimics that of zoisite. Hydrous minerals that have equilibrium hydrogen isotope fractionation factors that meet this requirement include the Fe-rich micas and amphiboles (e.g., Suzuo~l and EPSTEIN, 1976; GRAHAM cf al.. 1984). To illustrate this effect of Fe on fractionation, we have used hydrogen isotope fractionation data given by GRAHAM ef al. (1984) for arfvedsonite (- 35 wt.? total Fe0 compared to - 10 wt.% total Fe0 for actinolite) to calculate an additional curve for 400°C (Fig. 8). At 4OO”C, the equilibrium hydrogen isotope fractionation factor for arfvedsonite-water is approximately -50’5~ (GRAHAM et al., 1984). A three-fold addition of Fe (as total FeO) has resulted in a lowering of the “1R_U. by nearly 15%0at 50% alteration. A similar trend

1533

Basalt-seawaterisotopic exchange might be expected for the addition of other Fe-rich phases to the basalt for which we presently lack niineral-water equilibrium fractionation factors, such as Fe-rich smectite and chlorite. This contention is sup ported somewhat by our results. Substantial amounts of Fe are found in secondary phases in the basalt-seawater experiments (M. J. Morr~, unpublished data), which could explain why the observed fractionation factors lie somewhat closer to the annite-water curve than curves given for more Mg- or Al-rich hydrous phases (Fig. 7). This effect of Fe on isotope fractionation also probably accounts for the poor agreement between the measured and calculated rock-water fractionation factors (Table 3, Fig. 7). Clearly, more work on the effect of Fe substitution into hydrous minerals on isotope fractionation is needed before these experimental data can be fully understood. APPLICATIONS

TO NATURAL

SYSTEMS

Equilibrium fractionation factors bet ween basalt and seawater In estimating the Al80 and 6D values of fluid from the isotopic composition of bulk basalts at known temperatures, or in modeling basalt-seawater interaction, various investigators have approximated the equilibrium fractionation factors between basalt and fluids with various mineral-water fractionation factors (e.g., muscovite-water by SFOONER et al., 1977; plagioclase-water by TAYLOR, 1979; calculated basaltseawater by BOWERSand TAYLOR, 1985). The results of our experimental study suggest that the best method

for determining equilibrium oxygen isotope fmctionation factors in the basalt-seawater system is to use the average of alteration mineral-water fractionation factors, weighted according to the abundance of the alteration minerals. The equilibrium mineralogy and mineral abupdances can be estimated either from mass transfer calculations simulating water-rock interaction (BOWERSand TAYLOR, 1985) or by detailed modal analysis for natural systems. When the alteration mineralogy is not available, the fractionation curve of 1000ln~Ew=

-3.90(103/T)

+ 3.42( 106/T2)

fit to the experimental data given in Fig. 6 may be used for oxygen isotope exchange. This curve falls virtually on top of both the theoretical altered basalt-seawater curve (W/R = I ) calculated in Table 3 and the plagioclase (An = 0.5)-water curve (not shown), and within about OSL of the muscovite-water fmctionation curve (see Fig: 6). The experimental basalt-seawater curve is applicable only to systems where the water/rock ratio does not exceed 5. At higher water/rock ratios (> 10) this curve may be off by as much as about l?‘mat 500°C and 3% at 2OOOCfrom the true values (see Table 3). As a test, we used the experimental curve (Eqn. 7) to estimate the temperatures of oxygen isotope fractionation between altered basalts and seawater in the Reykjanes, Iceland geothermal system (drill hole 8). Oxygen isotope analyses for the fluids average - 1.2%0 and for the altered rocks (e.g., palagonite breccias, tuffaceous sediments and basalt) we obtained the following data: 6.0%0 at 438 m depth (T measured = 150°C);

-10

-20 -30 -40

-50 -60 5 -70 z? + ak -20 -30 -40 -50 -60 -70 0

20

40

60

80

0

20

40

60

80

!OO

WEIGHT PERCENT ALTERATION

FIG. 8. Calculated (solid or dashed curves) and measured (boxes) hydrogen isotope fractionation factors between altered basalt and seawater, AR_,,, , plotted against weight percent alteration. The 6D of the unaltered basalt was set at -72%, and an equilibrium value of 5% was used for the seawater. Fractionation factors for smectite. zoisite, actinolite, talc, and chlorite are taken from equations given in the Appendix. The 4OO”Cdashedcurve is based on arfvedsonite-water hydrogen isotope fractionation given by GRAHAM etal. (1984).

(7)

1534

D. K. Cole. M. J. Mott! and H. Ohmoto

(2) boiling--i.e.. approximately I 7% steam irtss occurred in drillhole 8, and (3) water-rock interaction. on the magnitude of “0 shifts in the fluid. The fi’Q shifts in the rock range from l.2 to over 4%, whereas the fluid may have shifted as much as !.h% in 6°C) (corrected for boiling and mixing). The fluid/rock maSS ratios range from 0.3 to about 1.6 with an average oi 1. For water-rock interaction at 27O’C (temperature at a depth of 1500 m) with grain radii on the order of 0.01to 0.1 cm, the time required to achieve 30% isotopic exchange at Reykjanes ranges from about 375 to 3750 years, respectively (Fig. 9). These magnitudes of time are similar to those estimated by ARNASON( 1976) for the durations of fluid flow from recharge zones into geothermal aquifers in other areas in Iceland with comparable hydrology and lithology to KeykJanes. However, these times do not compare favorably with calculations based on sulfate-sulfide sulfur isotope ex change given by SAKAI et al. ( 1980). which Indicate that the fluids at Reykjanes have a residence time of less than 50 years. At present, the reason(s) for this discrepancy in the time estimates is not understood. Rates of oxygen isotope e.rchange between The difference in the relative rates of sulfur and oxygen basalt and seawater isotope exchange-i.e., SO4 -- H2S sulfur > basalt-seaThe rate constants, r. obtained from our expenwater oxygen-could lead to more rapid reequilibramental systems can be used to compute the time retion ofthe sulfur isotopes, thus giving shorter apparent quired to achieve a given fraction of isotopic exchange times. (F). A genera! expression relating the (W/R) mole ratio, Another example of our calculations IS shown in density of the rock (p), average number of moles of Fig. 10, which illustrates the change with time in the oxygen in the rock (X,), the grain radius (a), the degree fi’*O of basalt and seawater at two different (W/R),,,, of isotope exchange (F), and time (t) was derived by ratios, I and 50. and two grain sizes, 0.0 I and 0.135em COLE et al. (1983) for fluid interaction with a spherical (see COLE, 1983, for details). Other conditions used in grain in a closed system where the computations-i.c., r. p, ,U,, etc.--are summarized in the figure caption. This figure may be used 60 eva! [ = -In (1 -- F)(W/R)L,uc, (8) uate the oxygen isotopic exchange reactions which may 3[1 +(W/R)]ry !(V4 take place in the mid-ocean ridges. We have selected ( !Oe4 is a factor used to convert cm* to m’). The results 350°C as the temperature of interaction based on the of calculations that used this expression are given in maximum vent temperatures measured for sea !loor Fig. 9 where the time required to attain 90 percent hot springs at 2 1"N on the East Pacific Rise (EDMOND isotopic exchange is plotted as a function of tempercl al.. 1982). Although the chemistry of al! springs ature for different grain radii (a = 0.5,O. 1,0.01.0.00 1 sampled to date on the Galapagos Rift at 8h”W and cm) and water/rock mass ratios (W/R = I. 25). the East Pacific Rise at 2l’N indicate a water/rock Figure 9 indicates, for example. that a system with mass ratio for the hydrothermal systems of about one a = 0.01 cm and (W/R),, = 1 requires approximately (MOTTL, 1983) we have calculated two cases--~ t’. 1880 years to reach 90 percent oxygen isotopic ex- “rock-dominated” with a (W/R) = I and a “waterchange at 200°C but only about 22 years at 500°C. dominated” case where (W/R) = 50 (by mass) An increase in the grain radius by one order of magIn the low (W/R) case (a = 0.05 cm. dashed Imes). nitude increases the time needed for 90 percent equil- both the fluid and the basalt d’*O values shift toward ibration--i.e.. 18,800 and 220 years for 200°C and one another. Since the seawater contains the majority 5OO”C, respectively. However, an increase from 1 to of the oxygen in this interaction (2: I), its isotopic com25 in the water/rock mass ratio produces only a slight position will shift less than the basalt--i.r.. 1.W&O shift increase (-0.2 log units) in the time required to achieve in Seawater versus 2.0% shift in the basalt. isotopic 90 percent exchange. equilibration for this system (i.e., F’> 0.99) is attained In the case of drillhole 8 at Reykjanes, COLE (1980) after approximately 900 years. Conversely. at d estimated F values that range from 0.75 to over 0.90, W’/Rhn,, = 50 and a = 0.05 cm (solid line), the 6’Y) suggesting the presence of isotopic disequilibrium. of the basalt shifts significantly, on the order of 3%0, These values were determined by taking into account: whereas the seawater remains unchanged. Equilibrium (1) the role of mixing between meteoric water and sea- for this system is achieved after about 1200 years (F’ water-as detailed by 01 AFSSON and RILFY ( 1978). z 0.99). For systems with a grain size of0.0 1 cm ant! 3.3% at 1000 m (Tmeasured = 200°C); 2.6L at 1162 m (T measured = 2 10°C); 3.5% at 1220 m (T measured = 225°C); 2.3%~ at 1280 m (T measured = 240°C): and 3.4’% at 1370 m (T measured = 250°C). If we assume that the current geothermal fluid was responsible for the alteration and subsequent isotopic exchange, then our experimental fractionation data would indicate that above about 1250 m, isotopic exchange took place at temperatures 50 to 75°C hotter than measured in the borehole. Conversely, if the current thermal fluid was not responsible for isotopic exchange, then our experimental data would predict more negative 6”O values (-3 to -5%) for the paleo-fluids that equilibrated with altered rock at the measured temperatures (assumed to be constant through time). COLE ( 1980, 1983) considered a third alternative, wherein the observed oxygen isotope fractionations between altered basalts and a mixed seawater-meteoric water at Reykjanes are the result of isotopic disequilibrium (discussed in more detail below).

1535

Basalt-seawaterisotopic exchange T(T) 6

500

400

300

xx)

I

I

I

I

I

0

,/I 4.0

f.2

( i.4

i.0

I f.6

I

I

2.0

2.2

24

r03/T(Kl

FIG. 9. Time to 90% oxygen isotopic exchange wsus temperature for the basalt-seawater system. These curves were calculated from Eqn. (10) described in the test. The rate constants used in Eqn. (10) were taken from Eqn. (7). The grain radius (a, in cm) and the (W/R)_ ratio were varied from 0.001 to 0.5 cm and I to 25, respectively.

a W/W,, ratio of I (dotted line), oxygen isotopic equilibration is more rapidly attained, on the order of 200 years. Coarser materials are also common in the MORB environments. In particular, when first chilled, the basalt glass exhibits a more-or-less continuous fabric broken by cooling fractures and pillow margins, with shards on the size of several centimeters (reviewer’s comment). Oxygen isotopic equilibration during alteration of this coarse material will be slow. For example, at conditions where T = 35O”C, (W/R),, = 1 and a = 5 cm, oxygen isotope equilibrium is achieved after about 90,000 years. An increase in a to 10 cm results in an increase in the equilibration time to 180,000 years. Hydrothermal activity along active spreading centers is thought to be episodic due to multiple injections of magma at depth, self-sealing of the hydrothermal plumbing system, or both (MOGUL, 1983). Estimates of the duration of these episodes range from a few tens of years to a few thousand years. Assuming that the simplified calculations presented above are applicable to natural systems-i.e., those with relatively slow flow

rates-then there appears to be sufficient time for significant isotopic shifts to occur in the rock and seawater. Additionally, if the episodes of hydrothermal activity last as long as a few thousand years as suggested by LISTER (1974) and MOTTL ( 1983), then there is ample time for both oxygen and hydrogen isotopic equilibration to take place between the finer-grained material and seawater. An oxygen isotopic shift of 1‘SO or more in the vent fluids sampled prior to mixing would suggest a low (W/R) environment that has undergone extensjve reaction over a modest period of time prior to rapid transfer of fluid to the sea floor. CATHLES( 1983) has used the oxygen isotope rate data for basalt-seawater given by COLE( 1980) to model the hydrothermal system responsible for massive sulfide deposition in the Hokuroku Basin of Japan. The rate data were used to predict the changes in the 6’*0 values of fluids and rocks as the convective fluid-flow regime evolved through time. CATHLES(1983) also used the rate data to help constrain the thickness of the reaction zone near a fracture. Although the rates he used differ by about one order of magnitude (too high) compared to those described above, his selection of E,, is relatively close--i.e., 15 versus 11.5 kcal/mole. Therefore, as TP-time change, the magnitudes of change in 6”O his model calculated should remain approximately the same, but the absolute values will vary.

CONCLUSIONS For the first time, the systematics of oxygen and hydrogen isotopic exchange between rock and water has been investigated experimentally. Isotopic redistribution of both oxygen and hydrogen during basaltseawater interaction is controlled by alteration reactions similar to those observed in natural systemsi.e., formation of smectite, tremolite-actinolite, chlorite and talc. The abundances of these phases commonly increased with an increase in temperature, time, and the amount of glass in the starting basalt. In general, depletion of ‘*O and enrichment of D in altered basalts were observed at temperatures ranging from about 300” to 500°C. The magnitudes of isotopic shift in the rock accompanying alteration also increased with increasing temperature and time, and to a lesser extent with glass content. In most cases, the final measured b”O and 6D of the rock do not represent equilibrium values. In the case of the 6”O data, the Northrop-Clayton partialexchange method yielded apparent equilibrium fractionation factors between altered basalt and seawater of approximately 3.5, 2.0 and OSk at 300”. 400” and 5OO”C, respectively. These values compare closely to those estimated from calculations using equilibrium modal mineral assemblages (e.g., mass transfer). These same partial-exchange oxygen isotope data were observed to follow closely with those predicted from a first-order rate model. Rate constants for oxygen

E). R. Cole, M. J. Mottl and H. Ohmoto

1536

0

I 200

-1

I

400

I

____I-._

600 800 TIME lyead

..._i (000

1200

!40(,

FIG. 10. The change in the lr”O values for hypothetical basalt-seawater interaction. Calculations HC~C carried out at 350°C (r = 1.43X 10m9moles of oxygen/m’ of solid surface/se@ for the following combinatiilnc of grain radius (a) and water/rock mass ratio: a = 0.05 cm, (W/R),, = 1 (dashed lines); u = 0.05 : m z I (dotted lines). The initial #‘O values of the hasalr (W/R),, = 50 (solid iines); a = 0.01 cm. (W/R),, and seawater were 5.75% and 0.00/w,respectively. Density of the rock (p) was set at 3 gee ‘. and YR= 0 t)?.% moles oxygen g-’ of basalt.

isotopic exchange

between altered basalt and seawater

range from 10-9.5to 1O-8.omoles of oxygen/m’ of solid surface/set for temperatures of 300” to 500°C; an activation energy of 11.5 kcal/mole was calculated from these rate data. Using a simple closed-system model (for oxygen isotopes), we estimate the duration of fluidrock isotopic equilibration in MORB-type systems to be on the order of a few hundred to a few thousand years. Hydrogen isotope fractionation between altered basalt and seawater ranged from about -74%0 at 300°C. -62%0 at 4OO”C, to -480/w at 500°C. These fractionations are probably attributable to exchange between seawater and secondary phases such as smectite. chlorite, and amphibole which are enriched in Fe. wish to thank Drs. Karen Von Damm, David Wesolowski. Ed Drummond. and Don Palmer for comments and suggestions that greatly improved the manuscript. The authors would also like to thank Drs. H. P. Schwartz and J. Welhan for their reviews of this paper. Completion of this paper could not have been realized without the typing talents of Ceci Steele and Betty Benton. This study was supported by grants from The National Science Foundation. EAR 76-03724, EAR 80-07839 and EAR 85-08379 to Ohmoto, OCE-8 I-10913 to Mottl. and The Division of Engineering and Geosciences. Office of Basic Energy Sciences. U.S. Department of Energy under contract DE-ACOS84OR21400 with Martin Marietta Energy Systems. Inc. for Cole. Acknowledgements-We

Editorial

handling: H. P. Schwarc7

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1538

I>. R. Cole, M. J. Mottl and H. Ohmoto

&rn-~~ = A%+o. At&o 2 &&CT. Cj) Gph-“fl= -3.35( IO’/T) + 0.03( lob/T’), regression

WE~OLOWSKI D.

(h) (9

(1984) Geochemistry of tungsten in scheelite deposits: The skarn ores at King Island, Tasmania. Ph.D thesis, The Pennsylvania State University, 431~.

APPENDIX

(a)

(b)

(c) (d)

(e)

(f) (g)

of data given by WESOLOWSKI(1984) for andradite (analog for sphene). (k) Aom+, = -4.66( lO’/T) -t-3.73( iO’/7’“). rcgress~on of data given by O’NEIL and TAYLOR( 1969) AL&o = -9.56(10’/7) + -4.86(106/T’! regression of data given by Srrzuo~r anli &S’I ~ih ( 1976). (1) A&+n,o = A”~-np. AD Ph14_Hzo = - 14.26( IO-‘/T) -- 3.?8( IO”/1 ). regression of data given by SUZUOKIand EPSTEIN( 1976). Cm) A~0,.np = -5.09(103/T) f 3.52( 106/T2). regression of data given by KULLA ( 1979). A~,,,.,,, = I .87( lO’/T) - 8.77(106/T’), regresston 01 data &en bv KULLA t 1979). = -3:47( IO-‘/r) t 4.‘89(106/p), rqrressior~ (n) $?“:nh& of data given by LLOYD ( 1968).

A: ISOTOPE FRACTIONATION FACTORS’

= -5.79( lO’/T) + 3.47( 106/T2). regression of data given by ONUMA o ul. ( 1972) and COLE (1985). A&,.w20= -2.5.19(10’/7) + 0.11(106/7‘2). regression of data given by TAYLOR(I 979). A&, = -6.99(103/T) + 4.97(106/T’). regression of data given by MAT-~HEWS e( ol f 1983) for zoisite (epidote analog). %lo = -35.9% for 300”-500°C. -8.4L at 200°C. data from GRAHAMet (11.I 1980). $&rJ = -2.99(103/T) + 3.26( 10”/T2), regression of data given by MATSUHISAet al. ( 1979). A:~~.H+J= -4.79( lO’/T) + 4.5 1(IO’/T’). regression of data given by MATTHEWSer al. (1983). A!LI.H?o= -29.0%0 for 300” to 500°C. GRAHAMn al. j 1984). L$,,,_~~ = A&.np, MATTHEWSet ul. (I 983). Trrm_HPL -2 I .7%0for 300” to 500°C. GRAHAM c/ al. (1984). A&_np = -3.97(10)/r) + 4.49(106/T’), regression of data given by MATSUHISAet al. (1979). A&+nlo = -6.45(103/T) + 1.08( 106/T2). regression of data given by BECKERand CLAYTON(1976). GLfl

(0)

$i”“-Hfl- @h&o.

~&l”“.“~ = - 1.75( Id/?-) -- 7.02( 106/1-)). regressron

-

of data given by TAYLOR(1979) for serpentine (analog). (P) A:n.np = -3.72( IO-‘/T) t 2.7 1(lob/T’), regresston of data given by MATSUHISAet al. (I 979) = -82.55(103/T2) + 17.51(106/‘7? (s) A:nlnn,u-~20 regression of data given by SUZUOKIand EPSI ~IY (1976).

-..-.-_---_..

--_.

’ These equations are appropriate only for the expertmental temperature range reported in the original study. 7”is in “K.