Oxygen isotope fractionation between analcime and water: An experimental study

Gemhimica ef Comwhimica

Acia Vol. 54, PP. 1359-1368

copv&ght 5 MMPqamon Press pk.Printed in U.S.A.

Oxygen isotope fractionation between analcime and water: An experimental study HARALDURR. KARLSON'** and ROBERTN. CLAYTON'-* ‘Department of the Geophysical Sciences, The University of Chicago, 5734 S. Ellis Avenue, Chicago, IL 60637, USA 2Emico Fermi Institute, Department of Chemistry, The University of Chicago, 5640 S. Ellis Avenue, Chicago, IL 60637, USA (Received December 30, 1988; accg~~~din r~v~ed~rrn rebury

16, 1990)

Abstract-The fractionation of oxygen isotopes between natural analcime ( - 100 pm) and water has been determined at 300, 3.50, and 400°C at fluid pressures ranging from 1.5 to 5.0 kbar. Isotope ratios were obtained for the analcirne framework, the channel water, and bulk water. Analcimes from Surtsey ( 145’C), DSDP Hole 417A (30 to 55”C), and Guam (25’C) were used to constrain the fractionation factors below 300°C. Analcime channel water exchanged completely with external water in all runs. Although some retrograde exchange may have occurred during quenching, the results indicate that the channel water is depleted in “0 relative to bulk water by a constant value of -5L, nearly independent of temperature. Analcime is the first hydrated mineral found to have water of hydration depleted in “‘0. Analcime framework oxygen exchanged 80,90, and 96% at 300°C for 4 12 h, 350°C for 178 h, and 400°C for 120 h, respectively. Equilibrium Ai80 ( %I) are as follows: 2.9 (4OO”C), 4.5 (35O”C), and 5.8 (3OO’C) for the experimental runs and 12.2 ( 145°C) and 24.2 to 28.2 (30-55’C) for the empirical data. The an~cime-water fm~ionation curve is within experimental error of that of ~lcite-water. The exchange had little effect on grain morphoiogy and does not involve recrystallization. This is the fastest exchange observed for a silicate. The rapid exchange rates indicate that zedites in active high-temperature geothermal areas are in oxygen isotopic eiprilibrium with ambient fluids. Once calibrated, zeolites may be among the best low-temperature oxygen isotope geothermometers. yielded the largest range in zeolite framework oxygen values ( 6 “0 r) . The 6 “0 r values of analcime can be measured with an accuracy of +0.2!%0 (KARLSSON, 1988; KARLSSON and CLAYTON, 1990). These observations suggest that an analcimawater geothermometer might be very useful (IbW.SsoN et al., 1985, 1988; KARLSON, 1988). Finally, analcime has the narrowest channels of any zeolite (2.6 A). If analcime exchanges its oxygen isotopes at low temperatures, more open zeolites might exchange even more readily.

INTRODUCTION OXYGEN ISOTOPEGEOTHERMOMETRY hasbeenusedovera wide range of temperatures, although it is most sensitive at low temperatures. With the exception of carbonates (EPSTEIN et al., 1953; O'NEIL et al., 1969) and quartz (CLAY'IONet al., 1972; MATSUHISAet al., 1979), however, few fractionation curves have been calibrated ex~~rnen~y with success at tem~tu~ lower than 3OO”C,so there are few calibrated mineral-pair geothermometers below 4OO’C (for compilations see, e.g., CLAYTON, 198 1; FAURE, 1986). This is a consequence of slow diffusion rates of oxygen through minerals, which make it difficult to achieve equilibrium in experiments at low temperatures. Hence the interior of many minerals is effectively rendered inaccessible to oxygen isotope exchange. However, the zeolite structure is highly porous and is penetrated by a network of channels in which the exchangeable cations and water molecules reside. Therefore the oxygen atoms of the zeolite framework should be easily reached from the outside. The great majority of framework oxygen atoms form the walls of channels, and those that do not are only one site away. Each framework oxygen atom is effectively a surface atom. Hence the kinetics of oxygen isotope exchange in zeolites should be controlled not by lattice diffusion but rather by surface reactions such as adsorption, exchange, and desorption. In order to test the feasibility of using zeolites as low-temperature thermometers, the oxygen isotope fiactionation between analcime and water was studied. Analcime was chosen because it has the widest inferred range of thermal stability of any zeolite (5 to 650°C) and has correspondingly

PREVIOUS WORK We know of no published studies of isotopic exchange rates and equilibrium fractionation factors for either framework or channel water oxygens in natural zeolites. There are some very scanty data on reaction rates of framework oxygen in synthetic zeolites that give an indication of the exchange rates that could be expected in natural zeolites. The exchange of oxygen atoms has been measured between synthetic zeolites and HZO, COZ, and O2 ( SAXENAand TAYLOR, 1963; ANTOSHINet al., 1971; PERI, 1975; GENE-BEet al., 1980; VON BALLMOOS,198 1; VON BALLMO@and MEIER, 1982). Only the exchange with water will be discussed here. The first attempt made to measure the exchange of oxygen between a zeolite and water was that of SAXENAand TAYLOR ( 1963) for Linde 5A. They found no measurable exchange at room temperature, but reported that 75% of framework oxygens exchanged with water in 2 h at 250°C. In a preliminary study, PERI ( 1975) stated that synthetic faujasite exchanged 3 1% of its oxygen in 1 h at 100°C. The most extensive work was that of VONBALLMOOS( 198 1) and VON BALLMOOS and MEIER ( 1982 ) , who measured the rate and extent of oxygen isotope exchange between NHe-forms of ZSM-5 and of mordenite with water. They reported that after 2 days

* Present address: Planetary science Branch, Code SN2, NASA Johnson Space Center, Houston, TX 77058, USA. 1359

li. K. Kartsson and R. N. (‘layton

1360

externally in a well near Ihe bottom of the bomb while the other

= 37) had exchanged 2.4% of its total oxygen and the mordenite (Si/AI = 5.6) had exchanged 4.2%. After 20 days the ZSM-5 had exchanged 6% of its total oxygens. at 95°C the ZSM-5 (Si/Ai

thermocouple was located internally in contact with rhe top of the capsule. In each case the temperature ditrerence between the two thermocouples was less than 3°C. Runs lasted from 5 to 17 days and were quenched (7’ :s 50°C) in I to 3 min. At the end of each run the capsules were weighed to check for leaks. If the weight change was less than r0.5 mg, the capsule was punctured and between 2 and 6 mg ofexchange fluid immediately extracted with it rnicrosyringc and introduced into a vacuum line. where the water was reacted with BrF, at 300°C for 90 min to liberate 0: (O’NEIL.and FI~S~I~IU, 1966:

SPECIAL CONSIDERATIONS OF ZEOLITE FRACTIONATION FACTORS

The isotopic compositions of an anhydrous mineral (m) and the enveloping fluid (w) can be related to one another by a single fractionation factor: CY,_,. In the case ofa zeolite. however. there are two inde~ndent fractionation factors:

MATSUHISAet al., 1979). As ;t precaution. the oxygen extraction

line was flushed with 3 to 4 mg of known water, whose diXO vaiue was similar to that expected from the run water. prior fo reacting a water from a run. This was done in order to minimize possible memory effects from earlier water samples. Finally, the capsule was cut open and run-products washed with distilled water into a beaker and allowed to dry at least overnight at room temperature IR air. In order to test the possibility of retrograde exchange during dying. two samples of the analcime used in the hydrothermal experiments were sealed in test tubes along with liquid water of known isotopic composition (al80 = -23.8%/on)for periods of I h and 80 days, respectively. Neither framework (13.2%) nor channel water (-16.8%~) ;,‘*O values changed significantly as a result of room-temperature drying. The analcime run-products were loaded into aspeciallv built dehydration vessel and dehydrated under vacuum to 450°C -in the manner dcscribed by KARLSS~N( 1988) and by KARLSSONand (‘IL\\ TOM(in prep.). The channel water was collected and analyzed for oxygen isotopes as described above. Oxygen was extracted from dehydrated analcime by the method Of CI.A\‘TONand MAYEUA( I963 ). The Oz liberated from channel water and dehydrated anaicime WE converted to COz and isotopically analysed using a triple collecting mass spectrometer: fi”O values are reported relative to SMOW

(1) and ( ‘80/‘60)f af-cw

=

( 180,160)ru

where cw is the channel water, w is the ambient fluid and f is the framework. The fractionation factor between the framework and the ambient fluid. afmw. can be related to af+ and LU,_, as follows:

EXPERIME~AL

Analcime-Water:

Fractionation

METHODS

Factors and

Exchange Rates

1. Metho&

The results of the oxygen isotope exchange between analcime and water are summarized in Tables I. 2, and 3. It is clear from the data in Table 1 that considerable exchange has occurred at temperatures as low as 300°C over a period of 17 days. The equilibrium fractionation factors and the degree of exchange for the channel water and the analcime framework are displayed in Tables 2 and 3, respectively. The apparent fractionation factors f flu,)and extent of exchange were calculated according to the method of NORTHROP and CLAYTON (1966). C~unne~ ~~,ut~r-~t~iktr,Llterfiatriunatio. The exchange was at or near completion over the tern~~ture range investigated (Table 2). The apparent equilibrium fmctionations range from -5.8%@at 300°C to -6.7% at 400°C (Table 2). An excetfent linear fit was obtained on a plot of (inn, - InLu‘) versus Intu, for the 300°C samples even though these were run at pressures ranging from 1.5 to 3.0 kbars. Thus. pressure had no measurable effect on the fractionation factor

The fractionation of oxygen isotopes between analcime and water was studied at temperatures from 300 to 400°C and at pressures ranging from I .5 to 5 kbars. Analcime is within its stability field at these pressures and temperatures ( MANGHNANI, 1970; KIM and BURLEY, IQ7I ). Approximately 20 mg of solid (IMNH-7522, 200 to 325 mesh; for sample identification see KARLSON and CLAYTON, 1990) and 200 mg of water of known isotopic composition were sealed together in a gold capsute (L: 3.5 to 4.4 cm, OD: 5 mm, LD: 4 mm}. Two kinds of devices were used: a cold-seal bomb and an internally heated pressure vessel, for pressures up to 3 and 5 kbar, respectively. Temperatures were read to 1.1“C with chromel-alumel thermocouples. Two ihermocouples were used simultaneously in all runs except one mid-seal run. In runs in the internally heated pressure vessel the thermocouples were placed 6 mm apart at the sample end of the capsule. In the cold-seal runs one thermocouple was situated

Table 1. Run condirions

Run#

and results of oxygen

Tcmperanrre&ssure Tii CC) (hrs) Wu)

isotope exchange

Initial S1soC’,

or

(?W 0,

G

experimcnrs

Final81&W) C%i @wow or

5‘.Anal

300

1.8

41.5

12,2

-16.8 -23.9 -11s

62.Anal

300

3.0

412

12.2

-16.8

16.7

for analcime-water

20.0

MAnal

300

I.5

423

12.2

-16.8

5.5

11.2

52.Anal

350

1.9

178

12.2

-16.8

-23.9

-15.2

-28.3 -22.x 9.0

Ma&d Balance(*) @L*IO) 0.1t 0.3

15.S

0.6 f 0.3

-0.2

5.9

-0.4 f 0.3

-211.4

-22.6’3’

58.Anal

350

1.9

189

12.2

~16.X

167

19.8

9.4

16.1

0.1 + 0.3

6oAnal

4cHl

5.0

120

12.2

-16.8

16.7

19.4

9.2

16.7

-0.6 + 0.3

64.Allal

400

5.0

120

12.2

-16.8

-23.9

-17.7

-28.0

-21.8

-0.4 + 0.3

(‘)Analyticai pruxton i1).2% (lo). WBulk initial 8180 - bulk final 8180, when 6% bulk = n (S’8of) + m. (8’*Ocw) + r (818Oy), D, III and r wz the III& fractions of franwork. channel water and exchange water. respectively. The yield for framework oxygen was taken as 14.9 runoUms. (3)Na masured but calculated by asruming periect mass balance.

1361

Oxygen isotope fractionation by zeolites

Run P

Tcmpmmrc

Prcssun

Tvnc

(‘a

(kb=)

mm

ldlna Initial (+lo)(‘) 1.2 f 0.3

415

62-Anal

3M)

3.0

412

-33.6 t 0.3

-6.4 f 0.3

7oAllal

300

1.5

423

-225 + 0.3

-6.1 f 0.3

52.Anal

350

1.9

178

7.2 f 0.3

-6.0 f 0.3

S&Anal

3S0

1.9

189

-33.6 i 0.3

-6.5 f 0.3

6aAMl

4Oil

5.0

120

-33.6 f 0.3

-7.4 f 0.3

64-Anal

4m

5.0

120

7.2 f 0.3

-6.3 f 0.3

deviation obtained by prcpgation

of Qmrs acmding

Run %

fractionation facts

Tanpaaturc

Rrssun

Thxx

(‘0

War)

fhrs)

to the rules &en

w

97*3

&&JJ

99*3

u

97*i

in BEV[NGION (1969).

and menu of exchange for analcimc hamewofk and bulk waer

1031,a Initial

Final

(*la)(‘)

WJ)

S4Aml

300

1.8

41s

36.3 f 0.3

300

3.0

412

-4.5 f 0.3

7OAml

300

1.5

423

6.6 f 0.3

52-Anal .58-Anal

3S0 350

1.9 I.9

178 189

36.3 f 0.3

7.s4 f 0.3

-4.5 i 0.3

3.6 f 0.3

60-Anal 66Anni

400 40

5.0 5.0

-4.5 i 0.3 f 0.3

4.2 f 0.3

120 ml

obtained by pmpagatim of mm

36.3

11.6f

EXChangcd F.quiiibrium (*la)

62-Anal

[I)S~~~~

W) (Ha)

In spite of the possibility of retrograde exchange during quenching, the results obtained for channel water-bulk water fractionation are still interesting. The channel water in analcime is depleted in ‘*O relative to water (Fig. 1). This is opposite to the behavior observed in crystalline hydrates and clathrates, whose measured A’*( hw - w) (hw is hydrate water) are positive, ranging from - + 1 to - t lo%, at least at tem~mtu~ below 57°C (Fii 3 and Table 5; for references see Table 5). anv_* for analcime was not determined ex~~rnen~ly for temperatures below 300°C. However, the 6’80, values of three natural analcimes whose formation conditions can be reasonably well constrained give an approximation to the sign and magnitude of the fractionation at lower temperatures. These samples are the analcimes from Surtaey, DSDP-417A, and from Guam, whose 6’*0, values were -4.9%0, -7.4%0, and -9.2%, respectively (KARLSON, 1988; KAR LSSON and CLWKIN, 1990). Since the Surtsey analcime formed in present-day seawater (6’rO = 0.0%) at 145 f 5”C, one obtains A I*(cw - w) = -4.9960 at that temperature ( KARLSON and CLAYTON, 1990). Formation conditions for analcime in sample DSDP4 17A are not as well constrained as those of the Surtsey sample. However, seawater-derived porewater with which the analcime equilibrated probably had S’*O between 0 and -4% (LAWRENCE and GIESKES,1981; BOHLKE et al., 1984). The temperatures can be estimated from the al80 value of the associated calcite (27.4%; see KARLSSON,1988; KARISSON and CLAYTON, 199O)andthe~l~tewater oxygen isotope geothermometer of O’NEIL et al. ( 1969), assuming that the calcite was in isotopic equilibrium with porewater. Such an estimate yields temperatures of 30 to 55°C. The DSDP sample therefore places A ‘*(cw - w) for analcime between -3.4% (30°C) and -7.4% (55’C). No oxygen isotope correction was made for the “salt effect” in either the Stutsey or DSDP fluids since studies have shown this effect is barely detectable in normal seawater (EPSTEIN and MAYEDA, 1953; CRAIG and GORWN, 1965; SAFER and CAT, 1972; TRUESDELL,19741.According to YURT~EVER and CAT ( 1981) the mean annual temperature and isotopic composition of precipi-

between the channel and bulk water at 300°C. Figure 1 shows the relationship between temperature and LX,_,. Although the fractionation factor becomes slightly more negative with increasing temperature, ail three data points lie within 20 of each other. In light of the fast self-diffusion coefficient observed for water in anaicime (see, e.g., DYER and MOLYNEUX, 1968), it is possible that measured fmctionations were “frozen in” at some lower temperature or at a so-called closure temperature below the run temperature. The closure temperature is defined as the temperature where effective isotopic exchange stops (D~DSON, 1973; HART, 1981). In order to test the likelihood of retrograde oxygen isotope exchange between the channel water and the bulk water upon quenching, it is necessary to know how fast the channel water in the 100 pm diameter grains equilibrates with the enveloping fluid. The rate of exchange depends on the selfdiffusion rate of water in analcime. The self-diffusion rate is not known above 95°C (DYER and MOLYNEUX, 1968)and the data for temperatures below 95’C must therefore be extrapolated to higher temperatures. Such extrapolations are valid only if the activation energy, E,in the Arrhenius equation [D = 4 exp( - E/RT)]remains constant. In other words, the diffusion mechanism must be the same over the entire temperature range over which diffusion coefficients are extrapolated. With these sho~comin~ in mind, the effusion rates for water in analcime were calculated at 50” intervals from 100 to 400°C. These diffusion rates were then used in conjunction with the equation x2 = 2Df (see footnote in Table 4) to estimate how rapidly an analcime grain 100 pm across would reach equilibrium with surrounding water at the respective temperatures. The equations and results are displayed in Table 4. When these results are compared with the quench rates shown in Fig. 2, it appears that the actual channel water-water fiactionations were preserved from the 300°C run and perhaps even from the 350°C run. For example, the 350°C run was cooled down to 3OO“Cin roughly IO sec. According to Table 4 it would take 79 to 260 set for the channel water to exchange with the bulk water in the temperature range 350 to 300°C.

Table 3. Oxygen iwtcp

(flu)

-5.6 ?: 0.3

1.8

(&uKIzA

Equilibrium

(*lo)

300

S4-Anal

Exchanged

Final

(46) (*la)

0.3

4.4f0.3

u

82f4

4&sAu

WI*2

zA#LQL?

%zt6

5.2 f 0.3

2.6 i: 0.3

acccsding to the mie.6given in BEvliWlW

f1969).

1361

H. R. Karlsson and R. N. Clayton Temperature “C 500 400 350 300 I

0 ‘4

I

I

I

ClNClLClME CH*NNEL WATER-WATER

I

111 d -5-

r2n

1(=1oI[,oY--887 cc

-6 -7 _ -8 0.0

I 0.5

1.0

1.5

2.0

2.5

3.0

3.5

lo6T-’ FIG. 1.The temperature dependence of the analcime channel water-water fractionation factor for oxygen isotopes. Line A is the bestfit regression line through the data points. The extrapolation of this line is shown stippled. Curve B is drawn to show that there must be at least one minimum above 400°C since the fractionation curve

must go to zero at infinite temperature.

tation on Guam are nearly constant. Assuming that the Guam analtime was in isotopic equilibrium with local meteoric water (-5.3%) at 25°C yields A’*(cw - w) = -4.9%. Data derived from natural analcimes are thus consistent with negative fractionations at low temperatures. The analcime channel water-water fractionation thus obtained is shown in Fig. 3. Included for comparison are the A’*(hw - w) in other hydrated crystalline substances and the equilibrium fractionation of oxygen isotopes between liquid water and ice. It is clear from the figure that the distribution of oxygen isotopes between analcime channel water and water is distinctly different from that of any previously determined oxygen isotope partitioning between a hydrous substance and liquid water. The reason for this unusual behavior in the channel water of analcime must lie in the bonding and crystallographic environment of the channel water, Table 5 compiles all observed A ‘*(hw - w) values among crystalline hydrates (chlorides, sulfates, a clathrate, and ice), along with the hydration number, the temperature range over which fractionation has been observed, and the method used to obtain the partitioning. Analcime differs structurally from other crystalline hydrates in two ways. First, it has the lowest hydration number. Second, the water molecules in analcime at room temperature are disordered and have either extremely weak or no hydrogen bonding with the framework or one another (&RRAW et al., 1972; MAZZIand GALLI, 1978). This contrasts sharply with crystalline hydrates such as gypsum that have strong hydrogen bonds. The nature of hydrogen bonds may in fact be the major difference between water in zeolites and in other crystalline hydrates such as sulfates or chlorides. According to recent neutron diffraction studies, hydrogen bonds in zeolites are among the weakest observed in crystalline hydrates ( ARTIOLI, 1985; KVICK, 1986). Framework-bulk water fractionation. The results of the oxygen isotope exchange between the analcime framework and bulk water are summarized in Tables 1 and 3 and in Figs. 4 through 6. The exchange was extremely rapid and increased with temperature. Nearly complete exchange was obtained in five days at 400°C and roughly 80% in I7 days at 300°C. This is remarkably fast, especially in light of the large grain size (- 100 pm) used during runs. By comparison, MATSUHISA et al. ( 1979) obtained only 50% oxygen isotope exchange between natural albite finer than 2 pm and water at 400°C and 15 kbar in 7 days. The equilibrium oxygen isotope fractionation factors were calculated for all run temperatures using the method of NORTHROPand CLAYTON(~~~~). The 3 runs at 300°C yielded an interpolated fractionation at 5.8% (Fig. 4), although individual pairs ofruns give fractionations ranging

from 4.9 to 6.3. corresponding to deviations of the data points from a straight line on a (Incu, -.- Inru,) versus lnai diagram ofabout 0.5%(,. Uncertainties ofthe order of tO.S%o probably apply to all the extrapolated fractions. The temperature dependence of the analcime framework-water isotopic fractionation is shown in Figs. 5 and 6 along with some other pertinent, experimentally determined mineral-water fractionations. Also included in Fig. 6 are the oxygen isotopic fractionations deduced from the two natural analcimes that were discussed above in the section on the channel water-water fractionation: namely the Surtsey (6’80f= 12.2%0)andDSDP-417A(6’*Of =24,2%0)samples,which yield rl’“(f - w) values of 12.2%0at 145 + 5°C and 24.2to 28.2% at 30 to 55”C, respectively. The analcime framework-water oxygen isotopic fractlonations lie between those of calcite-water (O’NEIL et al., 1969) and quartz-water ( MATSUHISAet al., 1979). At 400°C the analcime fractionation is identical to the calcite fractionation of O’NEIL et al. ( 1969) and the albite fractionations of MATSUHISAet al. ( 1979) and MATTHEWS ( 1980). At 350 and 300°C. however, our analcime fractionation is higher than that of either calcite or albite and lies approximately halfway between the calcite and quartz fractionation curves. A straight line can be fitted through the analcime data points (line A in Fig. 5). The slope of this line is steeper than any of the other lines shown in Fig. 5, suggesting that the analcime fractionation is more temperature dependent than the plagioclase feldspars, calcite, or even quartz. However. when the analcimes from Surtsey ( 145 + 5°C) and DSDP417A (30 to 55°C) are also included, all the data can be satisfied with the calcite-water fractionation curve (Fig. 6). To conclude, the present experimental and empirical data on the analcime framework-water fractionation suggest that analcime-water fractionation below 400°C is similar to that of calcite-water. The d “0 relationship between associated natural analcime-calcite pairs discussed by KARLSSON(1988) and KARLSSONand CLAYTON( 1990) also supports this conclusion. EXCHANGE

MECHANISM

Introduction

Detailed studies of the kinetics and mechanisms of oxygen isotope exchange between minerals and fluids during hydrothermal experiments show that the exchange generally occurs in two stages. In the first stage, there is rapid isotopic exchange at the mineral surfaces which in turn is followed by a second stage of slower bulk exchange (e.g., MATTHEWS et al., 1983a,b,c). SEM-photomicrographs taken of the reactant grains before and at various stages of isotopic exchange have been used to discern the processes controlling the bulk exchange. Such information is vital to an accurate interpretation of measured exchange rates. Observations of morphological features indicate that the bulk exchange is governed by one

Table 4. Approximate

Temperature OC

equilibrium

times for water diffusion in analcnne

Equilibrium

Diffusion Ccefficienc(‘J cm2/,

loo

I .57

150

2.38 x lo-‘”

200

2.03 x 10-9

250

1.1s x 10-8

set

min

Tin&) hrs

x lo-”

days 9

1s 7 18

300

4.80 x 10-8

260

350

I.59

x 10-T

79

400

4.43 x 10-T

28

(~K?alculatcd from the expression: D = D&q@ZJRT), when Del.52 x IO-’ cmz/,C; the activatica energy. is 17.03 kcal/mol; and R. the Gas Constant, is I.985 cavK’mOi. Walculated solving (X)3 = 2 Df for I. where X is grain radius self-diffusion coefficient of water given in the table. X is 50 pm.

E.

in cm and D is the

1363

Oxygen isotope fractionation by zeolites

HYDRO~ER~L RUNS

f\

04 0

30

60

90

120

150

180

Time elapsed (seconds) FIG. 2. Cooling history during quenching of hydrothermal runs. The symbols represent the following: (3 m 400°C; 0 35O’C; 0 3OO’C. The 400°C runs were done in an internally heated pressure vessel and the remaining runs in cold-seal rod bombs. Runs were quenched nearly isobarically.

of two processes: either ~lution-pr~ipi~tion Or effusion (MATSWHISAet al., 1978; MATTHEWS et al., 1983a,b,c). During a solution-ptecipitation process, or so-called “Ostwald ripening,” larger grains grow at the expense of smaller grains. Smaller grains dissolve and the material recrystallizes on the St&aces of larger grains, resulting in a decrease in the overall surface area and lowering the total free energy of the system (MATTHEWSet al., 1983b). During a bulk diffusion process, isotopic exchange takes place at the interface between the crystal and the fluid with little or no change in morphology of the reactant grains. In all minerals studied so far, measured exchange rates with water have been governed by the second stage of bufk exchange, i.e., either by diffusion or by ~lution-p~ipi~tion followed by diffusion at later stage. The initial surface reaction has not been observed. Methods Starting materials and run produets were investigated using a SEM to delineate the exchange mechanism.

EDX-system showed that the laths had the same chemical com~sition as their host grains. The crystallites were probably formed when the material adhering to the original reactant grains dissolved during the runs and was redeposited during the quenching of the run. Quench crystals have frequently been observed in hydrothermal exchange experiments by other workers. Such crystals can pose a problem in exchange experiments if they constitute a considerable fraction of the product material, since their isotopic composition may be controlled by kinetic isotope efI’ects associated with precipitation rather than by equilibrium isotope effects (see, e.g., MATSUHISAet al., 1979). In the present study, however, the contribution of quench crystals to the analcime-water fmctionation can safely be ignored since they were volume~~ly insignificant in the run products. Discussion Exchange mechanism and kinetics The excellent preservation of reactant grains during the exchange as revealed by the SEM-photomic~phs excludes a ~lution-pr~ipi~tion process as the mechanism of oxygen isotope exchange between analcime and water. The idea proposed by COLE and OHMOTOf 1986 )-that redeposited and original materials could have similar morphologies and are thus in~stinguishable~n, however, not be ruled out. Whatever the actual m~hanism of exchange was, it must have involved a reaction or reactions that led to substantial exchange of the oxygen isotopes but conserved the original grains at the same time. The preservation of the original grains indicates that the zeolite framework was also left intact even though framework bonds must have been broken in order for the exchange to take place. Since all the oxygen atoms in

Tbmperahrre “C soos8o280

‘“II

10

so

so

28

a 7

1

*20

Results Photomicrographs taken of the run materials before and after exchange are shown in Fig. 7. The starting material consisted of m 100 pm diameter grains with sharp edges and conchoidal fractures (Fig. 7a) with some smaller grains adhering to the surface. After exchanging with water at 308°C for 17 days, which resulted in 80% exchange, very little change was observed (Fig, 7b); the grains were virtually unafhected. Only the smalbr grains on the crystal faces had disappeared; the sharp edges and conchoidal fmctums were perfectly preserved. Similarly small changes were observed at 350°C after 90% exchange (Fig. 7~). The greatest change occurred in the 4OO*C runs, which had completely exchanged (Fig. 76). However, even in this case the changes were only minor. The grain edges were slightly rounded and small lath-like crystal&es (m~mum siz -3 X 63 pm) appearedon the surface of the reactant grains. Semiquantitative analysis with the

FIG. 3. The temperature dependence of the analcime channel water-water oxygen isotope fractionation, along with other pertinent fmctionations for various hydrous substances. The fractianation lines shown are aa follows: A. analcime (same as in Fig. 1); B. analcimebest-fit regression line through all the analcime dam; C. mirabilite; D. gypsum; E. camallite. The symbols denote fractionations for the following: 0 analcime; 0 tachyhydrite; fh epsomite; l carnallite; 8 CuS045H20 and bischofite; Clgypsum; n rn~bi~~; •9 iw, (9 &&rate. The data for analcime are shown in Table 2 and for the remainder in Table 5.

H. R. Karlsson and R. N. Clayton

I364

Substance

bischofite cam&e tachyhydrite mirabilite

CO~pOsXiO~

Uydr.

No.@)

Temp.

A18%w w

MethwV

Refewnce

H2OiM

C”C)

(%I

MgC12. 6&O

f,

IO to40

R

synthesis

HORITA (1989)

MgC12 KC1 .6H20

1

lot040

x

synthw

HORITA

‘&Cl2 2MgC12 .12H20

4

2s

Y

synthws

HORITA (1989)

Na2SOJ.

IOH

(1989)

5

0 and 25

1.5

synthesis

STEWART(lY74)

grpsum

CaS04

2HzO

2

171057

4

synthesis

GONRANnNl

ep%Xllite

MgS04.

Pi20

7

25”

8.4

CUSOJ

5H30

F

25

98

clafhmte

THF@)

ice anal&

17H~O H20

NaAiSizO6.

Hz0

1

and FONIES (1963)

synthesis

HE~N?J.NCER and Go12 (I975P)

4

2.7

synthesis

DAVIDSON 8r uf

0

3.0

synthesis

O’NEIL 1196%

-5.5

exchange

Thiswork

2sto4Oo

f1972)

SoFER and GAT

i 1U83)

(:)Number of water molecuies per cation. %echnique used todetennine the fractionation, Where method was not specifxd a blank space has bzen left. c’rToral crystal warer. This sulfate contains three types of isotopically distinct water molecules ~MAIWALD and HEINZINGER. 1972: HEINZMGER and G&Z, 1975).

the analcime framework are essentially surface atoms, as discussed earlier, the exchange process can be considered a surface reaction in the sense that rate of exchange is controlled by sorption-exchange-desorption rather than diffusion or recrystallization processes. A plausible exchange mechanism involves the formation of transient hydroxyl groups involving a hydrogen atom or proton from the channel water and a framework oxygen. The fo~ation of transient structural hydroxyls as the means of exchanging framework oxygen atoms without destroying the mineral structure has been proposed frequently, dating back as early as OBLAD et al. ( 1953 ), who studied the rate of oxygen isotope exchange between water and a silica-alumina oxide at 450°C. The formation of transient hydroxyl groups was also proposed by MCNJLSON and ROBERTS ( I96 1) as a mechanism for solution of water in silica glass and as the catalyst for Si /Al disordering and oxygen isotope exchange in feldspars ( DONNAY et al., 1959: WYART et al., 1959). All these workers assumed that the water dissociated either just before or just after entry, pr~ucing protons and hydroxyls which then diffused as such in and out of the crystals. However, O’NEIL and TAYLOR ( 1967 ) , who studied oxygen isotope exchange chemistry of feldspar% argued against such an exchange process: “it is difficult to envision the isotopic exchange in these systems as resulting from simple solid-state diffusion . , and somehow causing rupture and reforming of Si-0 and AI-O bonds. . . . This is to say nothing of the problems of size and charge of the diffusing species” (pp. 1433- 1434 ) . They offered an alternative mechanism involving a fluid film that migrates through the feldspar, dissolving unexchanged mineral in the front of its path and recrystallizing an identical but exchanged mineral in its wake. Since the reactant mineral and the product mineral would be the same, such a m~hanism would not be detected by morphological analysis. The source of water is not a problem in the case of analtime-it is supplied by the channel water. A possible exchange scheme involving the formation of transient hydroxyls in analcime is shown schematically in Fig. 8. The scheme is analogous to that suggested by VON BALLMOOS ( 198 1) to explain oxygen isotope exchange in the synthetic zeohte ZSM5 discussed earlier. The exchange rate observed for the analcime framework is the fastest observed in natural silicates to date. It is much

greater than that observed for albite at 400°C and 1.5 kbar ( MATSUHISAet al., 1979). Although few data are available on the relative oxygen isotope exchange rates for other silicates, the existing data can give some interesting information. For example, from the results of MATSUHISAet al. ( 1979) and MATTHEWS et al. ( 1983a,b,c) it is possibie to construct the following order for decreasing exchange rates in silicates: feldspars > quartz > wollastonite > jadeite 2 hedenbergite = diopside > zoisite. MATTHEWS et al. ( 1983a) noted that the exchange rates were highly dependent on the exchange mechanism and attributed the rate sequence above to differences in the exchange mechanism and the rate of recrystallization. That is, the solutionprecipitation coupled with slow recrystallization rates were the dominant processes in zoisite and the pyroxenes while diffusion and rapid rec~s~li~tion were the dominant processes in quartz. The feldspars showed little signs of recrys~lli~tion and their exchange rates were governed by diffu-

15 10 -

_\\

f &$ =581-122(+-Ai3 -821%exchanged

5

(1 51

0 -5 -

h=,Gwlnu (#I Kbar

130)

I t2o

->a -30

I

I

I

I

-25

-20

-15

-10

I

I -5

0

5

lOOO(lna,-hai)

FIG. 4. Estimationof the analcimeframework-water fractionation for oxygen isotopes at 300°C using the partialexchangemethod of NORTHROP and CLAYTON (1966). The intercept of the line with the Y-axis givesthe equilibrium fractionation and the inversEof the slope the extent of exchange. Note that the points deviate slightly but not significantly from a straight line. Numbers in parenthesesshow pressures (kbar) for each run.

Oxygen isotope fra~onation Temperahire “C 500

400

330

300

230

IO 3 i

ANUCIYE

3_

I f20

FRAYEWORK.WATER

rE( i

B. ::

(O’Neil elill..1968~

II,

3-

*

2_ lo-1 -2

, 0.0

, 0.5

,

, 1 .o

.

,

,

1.5

,

*

2.0 104

, 2.3

*

,

,

3.0

( 3.5

, 4.0

T-*

FIG. 5. The temperature dependence of the analcime frameworkwater oxygen isotope fractionation. Experimentally determined values. Line A is a best-fit regression line through the data. The fractionation lines for other silicate-water systems and calcite water are

shown for comparison.

sion. Analcime can now be added to this list as having the fastest exchange rate thus far observed for a silicate. For analtime, however, exchange rates are determined not by diffusion or solution-precipitation, but rather by sorption-exchangedesorption. Because of the absence of recrystallization in analcime, this mineral is ideally suited for studying the kinetics of direct oxygen isotopic exchange in silicates. Preservation of oxygen isotope ratios in zeolites and porous silicates From the analcime results, we can draw inferences about other zeolites and porous minerals, assuming that their exchange rates are also controlled by rapid “surface reactions.” Zeohtes that have been exposed to waters hotter than - 300°C cannot be expected to retain their original oxygen isotope ratios but may have undergone complete isotopic exchange even if they appear “fresh.” Thus zeolites in high-temperature geothermal areas or low-grade metamorphic areas would rapidly exchange their oxygen isotopes with ambient fluids. In this respect the zeolites are analogous to the feldspars, which undergo subsolidus exchange with little or no physical signs of alteration (TAYLOR and FORESTER, 197 I; O’NEIL and TAYLOR, 1967; O’NEIL, 1987). Zeolites would be most likely to retain original oxygen isotope ratios in low-temperature geothermal areas and sedimentary or soil en~onmen~. It is possible that some zeolites in low-temperature geothermal areas (< 1SOYI) have retained their original oxygen isotope ratios, especially if they were formed in an aging geothermal system. This study may also shed new light on the exchangeability of oxygen isotopes in other minerals with open structures, such as the feldspathoids nepheline, Na3K[ A14Si40161, and leucite, K[ AISi,06]. Roth these minerals have open crystal structures which are either closely related (as in the case of nepheline) or identical to that of analcime, Na[AlSi, 061. Hz0 (as in the case of leucite). The structural similarity

by zeolitcs

1365

between these minerals and analcime suggests that they may also readily undergo oxygen isotope exchange at low temperatures. This explains why nepheline crystals from the Honolulu Volcanic Series, Hawaii, have abnormally high 6 ‘*O values ranging from 5.7 to 10.2%0 as compared to those of pyroxene, which are 5% (CLAYTON and O’NEIL, 1978, unpubl. results). The present work supports the conclusion that the rocks containing the nepheline have been altered at low temperatures. Leucite crystals from potassium-rich volcanic rocks from a variety of localities world-wide also have anomalously high 6 “0 values, ranging from $8 to 10.6%~according to TayIor and coworkers (TAYLOR et al., 1979a,b; TAYLORand SHEPPARD, 1986). The interpretation of the d “0 values of these leucites lies at the focal point of an ongoing debate between Taylor and coworkers on one hand (TAYLOR et al., 1979a,b; TAYLORand SHEPPARD,1986; TURI et al., 1986) and HOLM and MUNKSGAARD(1982, 1986) on the other concerning the origin of some potassic lavas from the Roman Province of central Italy. Taylor and his coworkers claim that these lavas were generated by a combination of processes of mantle metasomatism and crustal assimilation while Holm and Munksgaard contend that the lavas originated from melting of an inhomogeneous mantle. Taylor and his coworkers based their conclusion on the 6”O values of minerals resistant to oxygen isotope exchange, while Holm and Munksgaard relied on 6 “0 values of whole-rock lava samples. Taylor and coworkers argued that the use ofwhole-rock samples was invalid because they were invariably altered. To support their conclusion Taylor and coworkers pointed to high 6’*0 values of leucite phenocrysts. Their basic premise was that leucite could easily undergo subsolidus oxygen isotope exchange without showing any physical sign of alteration. The results obtained in the present study on the exchange rate between analcime and water support the premise, held by Taylor and coworkers, that leucite could suffer subsolidus oxygen isotope exchange.

Temperature “C 35

,

300330230 I I

150 I

I

50 I

25 I

0 I

(

ANALCIHE FRAYEWORK-WATER

0:

, 0

, 2

,

, 4

*

, 6

(

,

8

.

, 10

, 12

i

14

lo6 T-* FIG. 6. The temperature dependence of the analcime frameworkwater oxygen isotope fractionation. Experimentally and empirically determined fiactionations. The figure illustrates that the analcime

framework-water fdonation is very similar to that of calcite-watir. Line A is the same as in Fig. 5.

1366

H. K. Karlsson

and R. N. Clayton

FIG. 7. Scanning electron photomicrographs of starting materials and run products from analcime-water exchange experiments. The magnification is the same in all the photos. a) Starting material. Note jagged edges, conchoidal fractures and small grains adhering to the surfaces of the larger grains. b) Run #62-ANAL; 4 I2 h at 300°C. Although these grains are -80% exchanged, there are very small changes in the physical appearance compared with the starting material. Delicate features like fractures and sharp grain edges are well preserved. The only noticeable change is the disappearance of the small surface grains. c) Run #52-ANAL: 178 h at 350°C. After -90’;: exchange the original material is still amazingly well preserved. Edges appear slightly more rounded. Noticeable are small flakes that pepper the grain surfaces. d) Run #60-ANAL; 120 h at 400°C. Although the exchange is nearly complete. -96% the original grain features still persist. Grain edges appear slightly duller. Note the appearance of the needle-shaped crystals on some of the reactant grains. They are probably a quench product (zeolite P?) The figure shows that these new crystals are volumetrically very insignificant.

CONCLUSIONS

Oxygen isotope fractionation has been measured experimentally in the system analcime-water in the temperature range 300 to 400°C and extended on the basis of empirical data down to 30 or 55°C. Two fractionation factors were determined, one for the channel water-bulk water (tie_.,) and another one the framework-bulk water ((~r_~). The channel water is depleted in I80 by 5 to 6%0 relative to bulk water, and the fractionation is nearly independent of temperature over the entire stability range of analcime. Because of this small temperature dependence, oxygen isotope analyses of channel water in analcime may yield information about the isotopic composition of ambient fluids. Analcime

is the first mineral whose water of hydration is significantly depleted in “0 with respect to external water. Reasons for this unusual behavior in analcime are not clear, but the explanation may be linked to differences in hydrogen bonding in these minerals. Analcime is notable in lacking hydrogen bonds. The partitioning of oxygen isotopes between analcime framework and water are similar to those of calcite-water. Hence, temperature or fluid estimates from analcime 6180 values can be calculated from the calcite-water fractionation cullle. Probably the most exciting results from this study are those of the exchange rate and mechanism. At the highest temperature, 4OO”C, exchange was virtually complete in five days.

Oxygen isotope fractionation by zeohtes

4

5

6

FtG. 8, A diagram depicting a possible reaction scheme for the isotopic exchange of framework oxygens in a zeolite. The sketches show a portion of the zeolite framework and 1 channel water molecule. During the reaction which leads to the exchange of a channel water ‘*O and a framework i60, temporary hydroxyls are formed (step #3). Chemical bonds in the sketches are represented with lines. Permanent bonds am whole lines, while temporary or weak bonds are shown as stippled lines. Note that the exchange leaves the framework unchanged.

The exchange occurred with no physical signs of recrystallization. The exchange rate for analcime is the fastest observed

for a silicate thus far. Although zeolites in geothermal systems may form met&ably during c~~li~tion from fluids, they will rapidly reach isotopic ~uilib~um. The extremely rapid exchange rate for analcime suggests that zeolites forming in active geothermal systems are in isotopic-and hence perhaps also in chemical-equilibrium with ambient fluids. In conclusion, once calibrated, zeolites could be among the best low-temperature oxygen isotope geothermometers. Acknowledgments-This research was a part of the senior author’s PbD dissertation and was funded by NSF grant EAR 8616255 to R. N. Clayton. This work could not have been done without the expert technical advice of Toshiko Mayeda. We are indebted to Jonathan Peterson for his help in setting up and monitoring the hydrothermal experiments and to Professor Julian R. Goldsmith for allowing us to use his hyd~~e~~ equipment. We thank James Eason for doing an excellent job in proofreading the manuscript. Reviews by Drs. A. Matthews, J. Ho&a, and S. Savin were helpful, Editorial handling: J. R. O’Neil

REFERENCES ANTOSHIN G. V., MINACHEVKH. M., SEVASTJANOV E. N., KON-

DARATJEVD. A., and NE&Y C. Z. ( 1971) Investigation of oxygen mobility in synthetic zeolites by isotopic exchange method. In Molcular Sieve Zeoliles--I; Adv. Chem. Ser. 101, pp. 5 14-5 19. ARTIOLIG. ( 1985) Structural studies of the water molecules and hydrogen bonding in zeolites. Ph.D. disseriation, Univ. of Chicago. VON BALLMOOSR. I 198 1) The U~8-exchanee~~hod in Zeolite Chemistry. Texte zur chkmie und chemiet&hnik, Salle-SauerIander. VON BA~OOS R. and MEIER W. M. ( 1982) Oxygen-18 exchange between zeolite ZSM-5 and water. .I. Phys. Chem. 86,2698-2700. BARRERR. M. and DENNYA. F. (1964) Water in hydrates. Part I. Fractionation of hydrogen isotopes by crystallization of salt hydrates. J. Chem. Sot., 4677-4684. BEVINGTONP. R. ( 1969) Data Reductionand Error Analysis~orthe PhysicalSciences. McGraw-Hill. B~HLKE J. K., ALT J. C., and MUEHLENBACHS K. ( 1984) Oxygen isotope-water relations in altered deepsea bar&s: Low-temperature mineralogical controls. Phys. Chem. Minerals 21,67-77. CLAYTONR. N. ( 1981) Isotopic thermometry. In Therm~ynamics

1367

o~~inerals and Melts (eds. R. C. NEWTON,A. NAVROTXY, and B. J. WOOD), Vol. 6, Chap. 5, pp. 85-109. Springer-Verlag. CLAYTONR. N. and MAYEDAT. K. ( 1963) The use of bromine pentatluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta 21,43-52. CLAYTONR. N., O’NEIL J. R., and MAYEDAT. K. ( 1972) Oxygen isotope exchange between quartz and water. J. Geophys.Res. 77, 3057-3067. COLE D. R. and OHMOTOII. ( 1986) Kinetics of isotopic exchange at elevated temperatures and pressures. In Stable Isotopesin High TemperatureGeologicalProcessesfeds. J. W. VALLEY,H. P. TAYLORJR., and J. R. O’NEIL), Vol. 16, Chap. 2, pp. 41-90. Mineral. Sot. Amer. CRAIG H. and GORDON L. I. ( 1965) Deuterium and oxygen-18 variations in the ocean and marine atmosphere. Syrup. Mar. Geochim.; Univ.Rhode Island Publ. 3, 277-374. DAVIDSOND. W., LEAISTD. G., and HE&SER. (1983) Oxygen-18 enrichment in the water ofclathrate hydrate. Geochim.Cosmochim. Acta 47,2293-2295. D~DXIN M. H. ( 1973) Closure temperature in cooling geochronological and petrological systems. Phys. Chem. Minerals 40, 259274. D~NNAY G., WYARTJ., and SABATIERG. ( 1959) Structural mechanism of thermal and compositional transformations in silicates. Z. Kristallogr.112, 161-168. DYER A. and MOLYNEUXA. ( 1968) The mobility of water in zeoiites-I. Self-diffusion of water in analcime. J. fnorg. Nucl. Chem. 30,829-837. EPSTEINS. and MAYEDAT. K. ( 1953) Variations of 0” content of waters from natural sources. Ge~him. Cosmochim. Acta 4,213224. EPSTEINS., BLJCHSBAUM R., LQWENSTAMH. A., and UREY H. C. ( 1953) Revised carbonate-water isotopic temperature scale. Bull. Geol. Sot. Amer. 64, 1315- 1326. FAUREG. ( 1986) Principlesof Isotope Geology.J. Wiley & Sons. FERRARISG., JONES D. W., and YERKESSJ. ( 1972) A neutrondiffraction study of the crystal structure of analcime, NaAISi70nH,0. Zeitschr. Kristalloar.135. 240-252. GEN%E k., ANDERSONT. F.,and F~P~AT J. J. ( 1980) Study of oxygen mobility in some synthetic faujasites by isotopic exchange with CO*. J. Phys. Chem, 84,3562-3567. GONFIANTINIR. and FONTESJ. C. ( 1963) Oxygen isotopic fractionation in the water of crystallization of gypsum. Nature ZOO,644646. HART S. R. ( 1981) Diffusion com~n~tion in natural silicates. Geeshim. Cosmochim. Acta 45,2?9-29 1. HEINZINGERK. and GO’rz D. ( 1975) A method to determine crystallization temperatures of hydrated crystals by intracrystalline oxygen isotope effects. Earth Planet. Sri. Lett. 27, 2 19-220. HOLM P. M. and MUNKSGAARDN. C. ( 1982) Evidence of mantle metasomatism: An oxygen and strontium isotope study of the Vulsinian District, Central Italy. Contrib.Mineral.Petrol.80, 367378. HOLM P. M. and MUNKSGAARDN. C. ( 1986) Reply to: A criticism of the Holm-Munksgaard oxygen and strontium isotope study of the Vulsinian District, Central Italy. Earth Planet. Sci. Lett. 78, 454-459. HORITA J. ( 1989) Stable isotope fractjonation factors of water in hydrated saline mineml-bone systems. Earth Planet.Sci. Lett. 95, 173-179. KARLSON H. R. ( 1988 ) Oxygen and hydrogen isotope geochemistry of zeolites. Ph.D. thesis, Universitv of Chicaao. KARL~SONH. R. and CLAYTON R. N. ( 1990) Oxygen and hydrogen isotope geochemistry of zeolites. Geochim Cosmochim. Acta 54, 1369- 1386 (this issue). KARLSSONH. R., CLAYTONR. N., and MAYEDAT. K. ( 1985) Analtime-a potential geothermometer? (abstr.) . Geol. Sot. Amer. 17, 622. KARL~SONH. R., CLAYTONR. N., and MAYEDAT. K. ( 1988) Fractionation and kinetics of oxygen isotope exchange between analtime and water (abstr.). Eos 69, 527. KIM K.-I. and BURLEYB. J. ( 197 1f Phase equilibria in the system

H. R. Karlsson and R. N. Clayton

136X

NaAlSi,O*-NaAISiO.,-Hz0 with special emphasis on the stability of analcite. Canadian J. Earth Sci. 8, 3 1l-337. KVICK A. ( 1986) Hydrogen bonding in zeolites. 7‘rans. .lmer. ( ‘rJ..ylaliogr. Assoc. 22, 91- 106.

LAWRENCEJ. R. and GIESKESJ. M. ( 1981 ) Constraints on water transport and alteration in the oceanic crust from the isotopic composition of pore water. J. Geuphys. Res. 86, 7924-7934. MAIWALI)B. and HEINZINGERK. (1972) Isotopieeffeckte in den Hydratstrukturen von festem und geliistem Kupfersulfat. Z. &‘ufwforsch. 27a, 8 19-826. MANGHNANIM. H. ( 1970) Analcite-jadeite phase boundary. Phys. Earth Planet. Int. 3, 456-46 I. MATSUH~SAY., GOLDSMITHJ. R., and CLAYTONR. N. (1978) Mechanisms of hydrothermal crystallization of quartz at 250°C and I5 kbar. Geochim. Cosmochim. Acta 42, 173-182. MATSUHISA Y., GOLDSMITHJ. R., and CLAYTONR. N. ( 1979) Oxygen isotopic fractionation in the system quartz-albite-anorthitewater. &whim. Cosmochim. Acia 43, I I3 I-I 140. MA’r7HEWSA. ( 1980) Influences of kinetics and mechanism in metamorphism: A study of albite crystallization. Geochun. Cotmochim.

Actu 44, 387-402.

MATTHEWSA., GOLDSMITHJ. R., and CLAYTONR. N. (1983a) Oxygen isotope fractionations involving pyroxenes: The calibration of mineral-pair geothermometers. Geochim. Cosmochim. Acta 47, 63 I -h44. MATTHEWSA., GOLDSMITHJ. R., and CLAYTONR. N. ( 1983b) On the mechanisms and kinetics of oxygen isotope exchange in quartz and feldspars at elevated temperatures and pressures. &ol. Sot. .dmrr. Bull. 94, 396-4 12. MATTHEWSA., GOLDSMITHJ. R., and CLAYTONR. N. (1983~) Oxygen isotope fractionation between zoisite and water. C&whim. Cosmochim.

Acta 47,645-654.

MAZZI F. and GALU E. ( 1978) Is each analcime different? Amer. Minerul. 63, 448-460.

C., and MCKAGUE H. L. ( 1981) Zeolite diagenesis below Pahute Mesa, Nevada Test Site. C/ay.y C/u>

MONUJRE G. K., SURDAM R.

Minerals 29, 385-396. MOULSON

A. J. and ROBERTSJ. P. (1961) Water in silica glass. Sot. 57, 120% I2 16. NORTHROPD. A. and CLAYTONR. N. ( 1966) Oxygen-isotope fractionations in systems containing dolomite. J. Geol. 74, l74- 196. OBLAD A. G., HINDIN S. G., and MILLS G. A. ( 1953) Dynamic structure of oxide cracking catalyst. J. Amer. Chem. SM. 75,4096Truns. Faraday

4098. O’NEII. J. R. (1968) Hydrogen and oxygen isotope fractionation between ice and water. J. Phys. Chem. 72, 3683-3684.

R. ( 1987) Preservation of H, C, and 0 isotopic ratios in the low temperature environment. In Stnbie Isotope Geochemistr? ql‘l,ow Temperature Fluids (ed. T. K. KYSER), Vol. 13. Chap. 2, pp. 85-128. Mineral. Assoc. Canada. O’Nrll. J. R. and EPSTEINS. ( 1966) A method for oxygen isotope

O’NEIL. J.

analysis of milligram quantities of water and some of its applications. J. Geophw Res. 71, 4956-496 I. O’NEIL J. R. and TAYLORH. P.. JR. ( 1967) The oxygen isotope and cation exchange chemistp of feldspar. Amer. ~Mineral. 52. 14141437. O’NEIL J. R.. CI.AY~O~~’ R. N., and MAYEDAT. K. ( IY69) Oxygen isotope fractionation in divaient metal carbonates. J. Chew Physics 51.5547-5558.

PER] J. B. ( 1975 ) Oxygen exchange between ClH02 and “acidic” oxide and zeolite catalysts. J Phys. Chem. 79, 1582-I 588. SAXENAS. C. and TAYLORT. I. (1963) Isotope exchange of the oxygen atoms in molecular sieve adsorbents with water and carbon dioxide. J. Inorg. Nncl. Chem. 25, 261-270. SOFER Z. and GAT J. R. ( 1972) Activities and concentrations of oxygen- I8 in concentrated aqueous salt solutions: Analytical and geophysical implications. Earth PlaneI. Sci. Left. 15, 232-238. STEWARTM. K. ( 1974) Hydrogen and oxygen isotope fractionation during crystallization of mirabilite and ice. Geochim. Cusmochim. Acta 38, 167-172.

TAYLORH. P., JR., and FORESTERR. W. ( 197 I ) Low-0i8 igneous rocks from the intrusive complexes of Skye, Mull, and Ardnamurchan, Western Scotland. J. Petrol. 12, 465-497. TAYLORH. P., JR.. and SHEPPARDS. M. F. (1986) Igneous rocks: processes of isotopic fractionation and isotope systematics. In Stable Isotopt~~ in High Temperalurc Geological Processes ( cds. J W. VALLEY, H. P. TAYLORJR., and J. R. O’NEIL), Vol. 16. Chap. 8, pp. 227-27 I. Mineral. Sot. Amer. TAYLORH. P., JR.. GIANNETTI B.. and TURI B. (1979a) Oxygen isotope geochemistry of the potassic igneous rocks from the Roccamonfina volcano, Roman comagmatic region, Italy. Earth Planei. Sri. Lctl. 46, 81-106. TAYLORH. P.. JR., TURI B., and CUNDARIA. (1979b) “O/‘“O and chemical relationships in K-rich volcanic rocks from Australia. East Africa, Antarctica, and San Venanzo-Cupaello, Italy. Earth Planet. Sci. Lest. 69, 263-276.

TRUESDELLA. H. ( 1974) Oxygen isotope activities and concentrations in aqueous salt solutions at elevated temperatures-Consequences for isotope geochemistq. Earth Planet. Sci. I.ett. 23,387396.

TURI B., TAYLORH. P., JR., and FERRARAG. ( 1986) A criticism of the Holm-Munksgaard oxygen and strontium isotope study of the Vulsinian District, Central Italy. Earth Planei. Sri. Lett. 78, 447-453.

WYARTJ., SABATIER G., CURIENH., DUCHEYLARD G., and S~VERIN M. ( 1959) Gchanges isotopiques des atomes d’oxyghe dans les silicates. Bull. Sot. Franc. Mink Crist. 82, 387-389. YURTSEVERY. and GAT J. R. ( 198 I ) Atmospheric waters. In Stable Isotope Hydrology:

Deuterium and Oxygen-18 in the Water C’.vcle

(eds. J. R. GAG and R. G~NFIANTINI): Tech. Rept. Ser. 210, Chap. 6. pp. 103-142. Intl. Atomic Energy Agency.