melt and CO2: An experimental study at 550–950°C and 1 bar

Geochimica et Cosmochimica Acta, Vol. 60, No. 11, pp. 1963-1973, 1996 Copyright © 1996 Elsevier Science Ltd Printed in the USA. All rights reserved 0016-7037/96 $15.00 + .00

Pergamon

PII S0016-7037(96) 00072-5

Oxygen isotope partitioning between rhyolitic glass/melt and CO2: An experimental study at 550-950°C and 1 bar J. M. PALIN,* S. EPSTEIN, and E. M. STOLPER Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA (Received July 14, 1995; accepted in revised form February 22, 1996)

A b s t r a c t - - O x y g e n isotope partitioning between gaseous CO2 and a natural rhyolitic glass and melt (77.7 wt% SiO2, 0.16 wt% H2Oto~) has been measured at 550-950°C and approximately 1 bar. Equilibrium oxygen isotope fractionation factors (O~CO2_rhyolit e = (180/160)CO2/(180/160)rhyolite) determined in exchange expe,riments of 100-255 day duration are:

T(°C)

1000 In O/co2_rhyolite

550 650 750 850 950

5.08 4.62 3.99 3.71 2.95

--+ 0.13 + 0.14 +-- 0.19 -----0.19 _+ 0.16

These values agree well with predictions based on experimentally determined oxygen isotope fractionation factors fiar CO2-silica glass (Stolper and Epstein, 1991 ) and CO2-albitic glass/melt (Matthews et al., 1994), if the rhyolitic glass is taken to be a simple mixture of normative silica and alkali feldspar components. The results indicate that oxygen isotope partitioning in felsic glasses and melts can be modeled by linear combinations of endmember silicate constituents. Rates of oxygen isotope exchange observed in the partitioning experiments are consistent with control by diffusion of molecular H20 dissolved in the glass/melt and are three orders of magnitude faster than predicted for rate control solely by diffusion of dissolved molecular CO2 under the experimental conditions. Additional experiments using untreated and dehydrated (0.09 wt% H2Ototal) rhyolitic glass quanitatively suppo:rt these interpretations. We conclude that diffusive oxygen isotope exchange in rhyolitic glass/melt, and probably other polymerized silicate materials, is controlled by the concentrations and diffusivities of dissolved oxygen-bearing volatile species rather than diffusion of network oxygen under all but the most volatile-poor conditions. 1. INTRODUCTION

ment that equilibrium oxygen isotope partitioning has been attained and measured. Despite their critical importance for interpreting oxygen isotope ratios in igneous rocks, experimental calibrations of oxygen isotope partitioning involving molten and glassy silicates are limited, covering thus far only silica glass, albitic glass and melt, and basaltic melt (Muehlenbachs and Kushiro, 1974; Matsuhisa et al., 1979; Stolper and Epstein, 1991; Matthews et al., 1994). In this paper, we report the results of an experimental study of oxygen isotope exchange between gaseous CO2 ( - 1 bar) and rhyolitic glass and melt at 550-950°C. We have employed the gas-solid exchange technique originally developed by O'Neil and Epstein (1966) and later modified for studies of oxygen isotope fractionation between CO2 and silica glass, albitic glass/ melt, and crystalline albite (Stolper and Epstein, 1991; Matthews et al., 1994). The essence of the technique is to equilibrate oxygen isotopes between a small amount of CO2 gas and a large reservoir of a silicate phase. For sufficiently large ratios of silicate to gas, the oxygen isotope ratio of the silicate is unchanged by exchange with the gas, and the CO2-silicate isotope fractionation factor can b e obtained from measure-

Knowledge of equilibrium oxygen isotope partitioning and of the mechanisms and rates of isotopic exchange between natural solids and fluids is essential for interpreting 180/~60 variations in earth materials. Exchange experiments provide the most reliable means for calibrating equilibrium oxygen isotope partitioning. Thele have been several recent studies of oxygen isotope partitioning involving many of the major rock-forming minerals under nominally anhydrous conditions (Chiba et al., 1989; Clayton et al., 1989; Chacko et al., 1991; Matthews et al., 1994; Rosenbaum, 1994). Although employing a range of plocedures, these experiments have yielded results that are consistent with one another and with reduced partition function ratios calculated by Kieffer (1982) to within 0.2%o or better for all minerals except quartz. Such agreement involving different experimental methods by independent laboratories provides a strong argu-

* Present address: Research School of Earth Sciences, The Australian National University, Canberra, ACT 0200, Australia.

1963

1964

J.M. Palin, S. Epstein, and E. M. Stolper Table 1. Stable Isotope Values of Starting Materials ~13C

1(3

5180

1(~

n

0-gas

-30.15

0.04

+0.39

0.25

13

20-gas

-12.93

0.09

+20.60

0.16

20

42-gas

-0.47

0.10

+42.44

0.13

16

Glass Buttes rhyolite

+7.76

0.16

8

RSQ quartz

+8.47

0.24

10

Explanation: l o = 1 standard deviation of multiple measurements; n = number of measurements; RSQ = Caltech Rose Quartz standard, accepted 8180 = +8.45 for NBS-28 8180 = +9.6.

ments o f the final 180/160 of the CO, gas and the initial 180/160 of the silicate phase. Reversals of the experimental results and estimates o f the extent o f oxygen isotope exchange between CO2 and silicate can be obtained via multiple runs that are identical except for the initial 180/160 o f the CO2 gas (Matthews et al., 1994). Our previous work has shown that polymerized glasses and melts are well suited for study by this technique (Stolper and Epstein, 1991; Matthews et al., 1994). Silica glass and albitic glass/melt do not crystallize or change composition (other than dissolving several ppm o f CO2) under the experimental conditions. The relatively open structures o f these materials permit rapid oxygen isotope exchange to take place. In addition, because the reduced partition function ratio o f CO2 gas is one of the most accurately known (e.g., calculations o f Bottinga, 1968, Richet et al., 1977, and Chacko et al., 1991, differ by less than 0.1%o over the temperature range o f this study), accurate reduced partition function ratios of silicate glasses and melts can be derived from the equilibrium fractionation factors obtained in these experiments. 2. METHODS

2.1. Starting Materials Three cryogenically purified CO2 gases, having 6180 values of approximately +0, +20, and +42 were used in the experiments (oxygen and carbon isotope ratios are given in the usual delta notation in permil deviations relative to V-SMOW and V-PDB reporting standards, respectively; e.g., 6 ~80~,,,oe = 1000 (( t8O/160 )sample/ (~80/~60)V_SMOW -- 1)). The carbon and oxygen isotope ratios of each gas, as taken through the gas loading procedure described below, are listed in Table 1 with uncertainties ( Icr standard deviation, as used throughout this paper unless otherwise noted). The natural rhyolitic glass used in this study was collected from Glass Butte, an extrusive volcanic dome in Oregon, by Phil Ihinger and Mike Carroll. It is devoid of phenocrysts and vesicles, but does contain opaque microlites. The b~sO of the rhyolitic glass was determined by conventional oxygen isotope extraction methods using F2 gas as the fluorinating reagent (Taylor and Epstein, 1962) and is listed in Table 1. The major element and normative compositions of the glass are given in Table 2. Of the total oxygen, 96 mol% comprises normative silica ( 39 tool% ) and alkali feldspar (57 mol% ) with 4 mol% in other components, principally anorthite and corundum. The glass contains low total dissolved water (H2Ototal = 0.16 _+ 0.03 wt%, dominantly in the form of hydroxyl, OH - ), and negligible dissolved CO2 (Blank et al., 1993). It has been used as a starting material in recent studies of CO, and H,O solubility, speciation, and diffusion and carbon and hydrogen isotope partitioning in rhyolite (Dobson et al., 1989; Ihinger, 1991; Blank, 1993; Blank et al., 1993).

Samples of the glass were crushed in a steel mortar. The fragments were sieved in high purity methanol, repeatedly rinsed, and optically examined and judged to be free of metal contaminants. They were then oven-dried at 110°C in air and stored in a desiccator. Before use in partitioning experiments, several batches ( 2 - 4 g) of the 45-75 /.tm size fraction of the rhyolitic glass were heated in the presence of 23-39 #tool of CO,_ at 600°C and - 1 0 3 bar for 3 - 6 h to oxidize possible organic contaminants introduced during sample preparation. This heat treatment produced no significant change in either the b~3C or amount of CO2, consistent with the absence of carbon contamination. In addition to the partitioning experiments, a series of experiments to examine the rate of CO2-glass exchange as a function of dissolved water content were conducted at 650°C. These experiments were done on splits of the 160-210 #m size fraction that were either partially dehydrated (i.e., heated at 600°C under vacuum (P < 10 ~ bar) for 3 days) or untreated (i.e., taken directly from the desiccator without any heating, even to remove organic contaminants as described in the previous paragraph). Measurements before and after the dehydration treatment by Fourier transform infra-red spectroscopy 1FTIR) indicate that the average water content of single glass fragments decreased from 0.16 to 0.09 wt% H20~ou~.

2.2. Experimental Techniques Capsules were made of either Pt tubing welded shut at one end (0.5 cm outer diameter, 12-15 cm long; cleaned in 3 N HC1 and

Table 2. Chemical and Normative Composition of Rhyolitic Glass

SiO2 TiO2 AI203 FeO MnO CaO

MgO K20 N a20 H2Ototal Total

wt % 77.70 0.07 13.00 0.38 0.04 0.52 0.05 4.19 4.08 0.16 b 100.20

mol O % a 83.36 0.06 12.31 0.17 0.02 0.30 0.04 1.43 2.12 0.20 100.00

Q

36.43

39.2

An Ab Or En Fs Wo IIm Co Total

2.58 34.51 24.75 0.12 0.58 0.07 0.13 0.81 99.99

2.4 34.0 23.0 0.1 0.4 0.1 0.1 0.8 100.0

Explanation:

Oxide wt% from Dobson

et al. (1989). a mol O % in oxide = 100 • moles O in oxide / moles O total, b from Blank (1993), +_0.03 wt % (la).

Oxygen isotope fractionation between glass and CO2 distilled H20, and annealed at 950°C in air for 1 h) or silica glass tubing closed at one end (0.9 cm outer diameter, 8 - 1 0 cm long; cleaned in prepared Chromerge® (H2SO4 + K 2 C r 2 0 7 ) and distilled H20, and dried at 110°C overnight). For each partitioning experiment, 250-320 mg of the heated 45-75 #m size fraction of glass were weighed into a capsule For the rate experiments, 80-90 mg of either the untreated or dehydrated 160-210 ~zm size fraction of the glass were used. The open end of the loaded Pt or silica glass capsule was attached to a vacuum line using Cajon Ultra-Tort® adapters, evacuated (P < 10 -2 bar) overnight, and gently heated to remove air and moisture adsorbed on the capsule walls and glass fragments. A 20-35 #mol ~diquot of CO2 gas was cryogenically purified and its quanitity mea~,ured using a mercury manometer (estimated uncertainty _+0.5 #mol ). This gas was transferred into the tip of the capsule using a liquid nitrogen bath. The capsule was then sealed approximately 3 cm from its top while still under vacuum and immersed in liquid nitrogen. In the case of Pt capsules this was accomplished using a pinch-off device (POD No. 375, CHA Industries) (Komarneni et al., 1979) followed by arc welding. Silica glass capsules were sealed with a hydrogen-oxygen torch. The end of each capsule containing rhyolitic glass (always -< 1 cm in length) was placed in the hotspot of a wire-wound horizontal furnace. Five furnaces held ~Lttemperatures of 550, 650, 750, 850, and 950°C were used. Temperatures were controlled with Eurotherm Model 808 temperature controllers connected to Type K or S thermocouples in the furnace hotspol:. Temperature gradients were less than 5°C over the length of the capsule containing the glass sample. Hotspot temperatures were independently monitored with a second Type K or S thermocouple adjacent to the capsule and typically varied by less than 3°C over the course of an experiment. Several capsules were typically run simultaneously in each furnace. Calculated internal CO2 pressures ranged from 0.8 to 1.5 bar depending on capsule dimensions and temperature. Run durations were between 5 h and 255 days. Samples were quenched by removing them from the furnace and allowing the~aato cool in air. After removal from the furnace and cooling, the capsule was attached to a vacuum line (P < l0 -~ bar) and opened. Platinum capsules were pierced using an assembly made from a bored-out Nupro® high vacuum valve with the stem tip replaced with a sharp steel needle. Silica glass capsules were cracked within an assembly made from Cajon Ultra-Torr® adapters and flexible stainless steel tubing. Condensable gases from the capsule were collected in a trap immersed in a liquid nitrogen bath. The amounts of noncondensable gases (either CO or air based on characteristic pale blue or reddishpurple discharge with a Tesla coil) were estimated using a calibrated volume and thermocouple vacuum gauge. After pumping away the noncondensable gases, the liquid nitrogen bath was replaced with a dry ice + " M 1 7 " (furfural: C5H402) slurry bath (T ~ -115°C) in order to vaporize CO2 but r~tain any H20 present. The COz was transferred to a mercury manometer, the quantity measured, and then transferred into a glass sample tube for subsequent stable isotopic analysis. The cold bath was ~hen replaced with a warm water bath and the amount of H20 was estimated using the calibrated volume and thermocouple vacuum g~tuge. All oxygen and carbon isotope compositions were measured on CO2 gas using a Fiimigan 252 isotope ratio mass spectrometer. The precision of an individual measurement on this instrument is estimated to be better than or equal to _+0.10 for 6180 and _+0.09 for 6 ~3C, based on multiple analyses of a standard gas over the course of this study. Procedural blanks were determined with experiments in which 20-35/zmol of CO2 gas were loaded into empty capsules and run for 3-111 days at 25-850°C. The results of these experiments are given in Table 3. Blank runs i~nPt capsules (n = 3) had quantitative CO2 yields (98 _+ 1%), contained minor amounts of CO (-<0.03 #mol) and H20 (-<0.02 #mol), and exhibited small (-<0.06) shifts in 613C and 6 tso of CO2, indicating no extraneous carbon or oxygen was introduced into the capsule during the gas loading procedure. Blanks determined for silica glass capsules (n = 9 ) also had quantitative CO2 yields (100 _+ 2% ), contained small amounts of CO (0.14 _+ 0.04 #mol) and H20 (0.06 _+ 0.03/zmol), exhibited the variations in 6180 of CO2 expected in response to exchange with the capsule walls, and had virtually unchanged 613C of CO2 (+0.06 _+ 0.07).

1965 3. RESULTS

3.1. G e n e r a l S t a t e m e n t The results of all experiments are listed in Tables 4 and 5. Data from two experiments exhibiting low CO2 yields (runs MP-98 and MP-187 ) are considered unreliable and are not discussed further. In the thirty-two remaining runs, the average CO2 yield was 101 _+ 2% with an average of 0.3 _+ 0.3 # m o l noncondensable gas and 0.1 _+ 0.1 # m o l H20. Negative shifts in the 613C values of COz over the course of the experiments (average 2x Iac = - 0 . 2 0 _+ 0.49, with four runs having A 13C _< - 1 ) suggests variable reduction of the gas to CO or exchange with small amounts of a C-bearing contaminant m a y have taken place. Rhyolitic glass run products exhibited no change in index of refraction or petrographic evidence of crystallization. The glass fragments visibly sintered in runs of 3 days at 850°C and 100 days at 750°C, and flowed in runs of 100 and 120 days at 850 and 950°C. In all other runs, the sizes and shapes of grains did not appear to change. These observations are consistent with the finding of Bacon ( 1 9 7 7 ) that the glass transition for rhyolite with low water content lies between 800 and 900°C. W e infer, therefore, that samples run above 800°C were supercooled melts under the conditions of the experiments whereas those run below this temperature remained glasses.

3.2. Partitioning Experiments Results of the CO2-rhyolitic glass/melt partitioning experiments are listed in Table 4. In most cases, two or three capsules were run simultaneously at the same temperature using CO2 gases with initial 6180 values of +42, +20, and +0. Thus, each set of experiments was reversed in the sense that the final oxygen isotope ratio of the gas was approached from both higher and lower initial values. Typical runs with 300 m g of rhyolitic glass (15.5 # m o l O2/mg) and 23 # m o l of CO2 had a glass-to-gas oxygen ratio of about 200. This is sufficient to allow CO2 gases with initial 6180 values of + 4 2 and + 0 to converge to within 0.2 per mil after complete gas-glass oxygen isotope exchange. All of the experimental sets converged to within 0.6%~, indicating that at least 30% of the oxygen in the rhyolitic glass underwent isotopic exchange with the gas. This extent of oxygen isotope exchange corresponds to a m i n i m u m exchange penetration depth of 3.5 # m for the 4 5 - 7 5 # m size fraction, which is calculated assuming the exchanged portion of an individual grain can be approximated by a spherical shell (e.g., M a t t h e w s et al., 1994). Data obtained from each set of runs were interpreted with the aid of the following expression for oxygen isotope mass balance between CO2 gas and rhyolitic glass/melt (cf. Northrop and Clayton, 1966; H a m z a and Broecker, 1974): NR Nco2 18 final -(6 6 Oco 2 = Otco2-RN~xchanged -NR 18

18

initial

Oco~ - 6

initial

+ Otco2-R(6 OR

18

final

Oco~)

+ 1000) -- 1000

(1)

in w h i c h R refers to rhyolitic glass/melt, aco2_ R is the oxygen isotope fractionation factor between CO2 and the e x c h a n g e d

1966

J.M. Palin. S. Epstein, and E. M. Stolper T a b l e 3. P r o c e d u r a l B l a n k s T(°C) Duration

Run #

Capsule

Final CO2 (p.mol)

Yield

(days)

CO2 (%)

NC a (gmol)

H20 (p.mol)

~.13C CO2 b

Initial 818 0 CO2

Final 818 0 CO2

Intercept $180 CO2 c

1000 In c~ (CO2 - Silica Glass) d

25

111

MP-182

Pt

24.7

97

<0.02

<0.02

-0.01

42.44

42.49

N/A

N/A

850

3

MP-114

Pt

26.7

99

0.02

<0.02

-0.03

20.60

20.59

N/A

N/A

850

3

MP-115

Pt

26.7

99

0.03

<0.02

-0.06

20.60

20.58

N/A

N/A

650

5

MP-86 MP-87 MP-90 MP-92

Silica Silica Silica Silica

26.3 23.1 23.1 23.7

101 97 100 99

0.05 0.10 0.07 0.07

0.04 0.10 <0.02 0.03

+0.03 +0.06 -0.06 +0.06

42.44 42.44 20.60 0.39

39.64 39.56 19.58 2.06

13.35 +_2,69

1.85 +_2.67

650

60

MP-117 MP-116

Silica Silica

28.8 27.2

100 96

0.14 0.14

0.04 <0.02

+0.02 +0.21

20.60 0.39

19.08 4.12

14.63 _+0.77

3.11 _+0.79

850

3

MP-88 MP-89 MP-91

Silica Silica Silica

26.3 25.3 27.3

99 104 100

0.22 0.24 0.22

0.10 0.10 0.07

+0.04 +0.09 +0.06

42.44 20.60 0.39

34.44 18.96 4.14

14.79 0.67

3.27 _+0.70

Explanation: a NC = total gas not condensable in liquid nitrogen trap; unless otherwise indicated, mainly CO. b z~13C CO2 = Final 8130 002 o Initial 813C 002. c Interpolated by weighted-error least-squares regression using Eqn.(1). d c¢ (CO2-Silica Glass) = (8180 CO2 + 1000)1(8180 Silica Glass + 1000), where 8180 Silica Glass = 11.48, estimated l o uncertainty calculated by propagating errors in 8180 of CO2 and silicate. N/A = not applicable.

portion of the rhyolitic sample at the end of the experiment (aco2 R = (180/160)CO2/(180/160)R), NR is the total number of moles of oxygen in the rhyolitic glass/melt, Nco2 is the total number of moles of oxygen in COz, and N ~ xcha"ged is the number of moles of oxygen in rhyolitic glass/melt which exchanged oxygen isotopes with CO2 at the conclusion of the experiment. This equation is derived from a simple mass balance equation for Jso and ~60 assuming that ( 1) CO2 gas is isotopically well mixed at all times, (2) rhyolitic glass/ melt can be considered in terms of unexchanged and equilibrated portions, and (3) total moles of oxygen is equal to total moles of ~60 (i.e., ~O/160 < 1). A set of experiments run at the same temperature and for the same length of time, but with CO2 gases of different initial 6~O values define a linear relation where intercept (b) and slope (m) values are given by b

= aco,-R(6 18ORinitial + 1000) -- 1000

(2)

and NR m

=

Olco2_ R N~xchanged .

(3)

Slope and intercept values for each set of experimental data were obtained by York-type error-weighted least-squares linear regression (G. Ravizza, pers. commun.). CO2-rhyolitic glass/melt fractionation factors and fractions of rhyolitic oxygen exchanged were calculated from these values by rearranging Eqns. 2 and 3. Measured and estimated l~r uncertainties in the initial 6 ~80 values of each CO2 gas and the rhyolitic glass, the amount of glass (+0.005 g) and CO2 gas (_+0.5 #mol) loaded in each capsule, and the final 6~80 of CO2 (_+0.10) were propagated through all calculations. The CO2-rhyolitic glass/melt oxygen isotope fractionation factors so determined are listed in Table 4. Only data from runs of 100-255 days duration at each temperature are considered to represent equilibrium values because, as revealed by inspection of Table 4, shorter duration experiments at

650-850°C yield anomalously large CO2-rhyolite oxygen isotope fractionations. When the results of experiments at 650°C and lasting 5 to 255 days are plotted versus duration (Fig. 1), it is evident that the data approach a constant value with time, which we take to be the equilibrium fractionation factor. The selected equilibrium CO2-rhyolitic glass/melt fractionation factors are plotted in terms of 1000 In ac<, R versus T 2 in Fig. 2. Also shown are curves illustrating experimental fractionation factors for CO2-silica glass and CO2-albitic glass/melt (Stolper and Epstein, 1991; Matthews et al., 1994) and a predicted curve for CO2-rhyolitic glass/ melt based on these previous studies and weighted for the proportions of oxygen contained in normative silica and alkali feldspar (which together constitute 96% of the total oxygen in the rhyolitic glass). The measured equilibrium COz-rhyolite fractionation factors agree with the predicted fractionation curve to within 0.3%~ between 650 and 950°C and to within 0.6%0 at 550°C. The cause of anomalously large CO2-rhyolite oxygen isotope fractionations in experiments of short duration is not clear. Although all these experiments were run in silica capsules, it is unlikely that this is the principal cause of the effect, as analytical blanks (Table 3) indicate that isotope exchange with the walls of the silica glass capsule could account for no more than 1% of the silicate oxygen exchanged. This amount of exchange would produce at most a +0.05%c shift in the final 6 ~80 values of the CO2 gas, much smaller than the 1-2%~ discrepancies which are observed. Moreover, we have encountered similar effects in short duration experiments involving crystalline albite and quartz in Pt capsules (Matthews et al., 1994; and unpubl, data). Because when previously observed, such anomalously large oxygen isotope fractionations were restricted to experiments in which the calculated isotopic exchange penetration depth was less than 0.4 #m, they were ascribed to some type of surface correlated phenomena. However, the large amount of oxygen isotope exchange in the experiments reported here ( -->30%) suggests penetration depths of several micrometers,

255

5

60

120

255

5

100

3

120

100

550

650

650

650

650

750

750

850

850

950

MP-160 MP-159 MP-158

MP-133 MP-132

MP-101 MP-100 MP-99

MP-162 MP-161

MP-110 MP-97 MP-96

M P- 186 MP-185

MP-135 MP-134

MP-106 MP-111

MP-95 MP-94 MP-93

MP-183

MP-157 MP-158 KAD 4 ~

Run #

Pt Pt Pt

Pt Pt

Silica Silica Silica

Pt Pt

Silica Silica Silica

Pt Pt

Pt Pt

Silica Silica

Silica Silica Silica

Pt

Pt Pt D*

Capsule

299 302 299

308 208

312 297 314

311 290

300 296 305

247 259

309 301

321 294

309 302 301

244

304 311 ~,~n

Wt. Silicate (mg)

26.8 26.8 24.2

23.7 23.1

25.3 22.1 25.3

27.9 23.2

30.4 24.2 25.8

25.2 21.4

23.6 23.1

34.0 27.1

23.1 22.1 24.2

21.4

23.6 25.7 '~ ~

Final CO2 (l~mol) CO2

102 106 106

103 104

100 103 98

100 104

98 101 98

104 102

100 103

101 99

100 98 104

102

105 102 ~'~

Yield (%)

0.05 0.14 0.03

>5h _>8 h

0.46 0.44 0.43

0.14 0.02

1.1 0.75 0.87

0.05 0.10

0.03 >7h

0.73 0.51

1.2 0.46 0.49

<0.02

0.02 0.12 0.15

<0.02 <0.02 0.02

0.02 0.02

0.22 0.26 0.22

<0.02 <0.02

0.05 <0.02 0.07

<0.02 0.10

<0.02 <0.02

<0.02 0.03

<0.02 0.24 0.24

0.02

<0.02 <0.02 -'~ '~'~

NC a H20 ( p . m o l ) (Fmol)

-1.38 -0.60 +0.21

-1.13 +0.25

-0.39 -0.13 +0.17

-1.00 +0.21

-0.22 +0.11 +0.35

-0.38 +0.09

-0.46 +0.32

-0.26 +0.36

-0.14 +0.01 +0.22

+0.08

-1.45 -0.58 +0.17

A13C CO2 b

42.44 20.60 0.39

42.44 0.39

42.44 20.60 0.39

42.44 0.39

42.44 20.60 0.39

20.60 0.39

42.44 0.39

42.44 20.60

42.44 20.60 0.39

0.39

42.44 20.60 0.39

Initial ~18O CO2

11.79 _+0.11

13.49 _+0.03

12.37 _-_-_-_K). 13

12.47 -+0.11

12.72 -+0.11

14.44 _+0.01

12.96 -+0.02

Intercept 8180 CO2 ¢

11.13 10.87 10.62

11.62 11.47

10.74 _+0.01

11.51 _+0.11

12.62 1 2 . 6 2 _+0.05 12.56

12.88

12.19 11.66

13.75 13.60 13.35

12.45 12.27

12.61 12.42

13.21 12.83

14.81 14.53 14.23

12.81

13.25 13.08 12.81

Final 8180 CO2

46 _+1

N/A

67 +15

44 _+12

65 _+11

N/A

N/A

40 +15

35 _+1

N/A

47 -+5

(%)

NReXCh/ NR d

2.95 _+0.16

3.71 _+0.19

4.81 _+0.16

3.99 _+0.19

5.67 _+0.16

4.56 +0.20

4.67 -+0.19

4.91 -+0.19

6.61 +0.16

5.01 _+0.20 g

I

i

II

2.95 _+0.16

3.71 -+0.19

3.99 _+0.19

4.62 +0.14

_+0.13

O,UO

1000 In c( 1000 In (CO2--Rhyolite) (CO2-Rhyolite) all e equilibrium f 5.15 i _+0.16

Explanation: All runs with 45-75 I~m sized rhyolite, a NC = total gas not condensable in liquid nitrogen trap; unless otherwise indicated, mainly CO. b A13C 002 = Final 8130 CO2 - Initial 813C CO2. ¢ Interpolated by weighted-error least-squares regression using Eqn.(1) in text. d Fraction of rhyolite oxygen exchanged based on Eqn.(3) in text. e ~ (CO;t-Rhyolite) = (8180 CO2 + 1000)/(8180 Rhyolite + 1000), estimated l a uncertainty calculated by propagating errors in 8180 of CO2 and rhyolite, f c¢ (CO2-Rhyolite) from sets of runs of 100 day duration or longer, averaged in cases of two such sets of data. g Single run, uncertainty assigned, h NC mostly N2.

100

550

T(°C) Duration (days)

Table 4. Results of Partitioning Experiments

,9

r~

5"

O

©

1968

J.M. Palin, S. Epstein, and E. M. Stolper Table 5. Results of Exchange Rate Experiments Duration 5 hrs

5 days

255 days

Initial Final Wt. H20 total H20 total Silicate (wt %) (wt %) (mg)

Run #

Final

CO2

CO2

Yield

(p.mol)

(%)

NC a H20 (p.mol) (/.tmol)

A13C CO2 b

Initial 81aO COlt

Final 8180 CO~

(10-18m2s-l) c

Doxy

MP-192 D MP-195 U

0.09 0.16

N/D N/D

90.4 79.8

21.6 25.2

98 100

<0.02 <0.02

0.18 1.4

-0.29 -0.23

42.44 42.44

30.15 25.32

9 :t: 2 32 _+ 8

MP-193 D MP-196 U

0.09 0.16

N/D N/D

80.1 82.3

26.3 21.0

102 100

<0.02 <0.02

0.48 >2

-0,13 -0.25

42,44 42.44

19.40 16.03

9 _+ 2 31 + 9

MP-194 D MP-197 U

0.09 0.16

0.009 0.02

85.7 82.5

25.1 21.4

98 102

<0.02 0.22

<0.02 0.07

-0.10 -0.17

42.44 42.44

14.33 13.29

4 + 1 36 + 10

Explanation: All runs with 160-210 p.m size fraction of rhyolite in Pt capsules at 650°C. a NC = total gas not condensable in liquid nitrogen trap; unless otherwise indicated, mainly CO. b &laC CO2 = Final 813CCO2 - Initial 813C CO2, ©Apparent diffusion coefficient for oxygen (Doxy)calculated as described in text. D = dehydrated, U = untreated, N/D = not determined.

almost certainly ruling out the predominance of such effects. Thus, although the asymptotic approach of the results to a constant value in the long duration experiments gives us confidence that we have achieved equilibrium oxygen isotope partitioning, we do not understand why the shorter duration experiments, in which the amount of exchange appears to have been very large, have not achieved equilibrium.

3.3. Exchange Rate Experiments The change in the oxygen isotope composition of C O 2 in exchange experiments in which isotope equilibration with the silicate oxygen reservoir is incomplete can be used to estimate apparent oxygen diffusion coefficients (Stolper and Epstein, 1991 ). If the initial gas composition is sufficiently far from the equilibrium composition (e.g., the experiments in which the initial 6~80 of the CO2 is +42), the apparent oxygen diffusion coefficient based on such an analysis is relatively insensitive to the CO2-silicate oxygen isotope fractionation factor. Thus, although as described above, short

8

, , , ,

, , ,

•--= o

I

T (°C) 750 I

550 I

Albitic Glass / Melt -,,.,,..~J

7

~"

6

~ ¢:r-

23

Rhyolitic Glass/Melt

~

/. / / "

silica (3)

e-.

tr"

950 I

7

i i i I I

650°C m

duration experiments yield anomalous fractionations, these results can still be used to constrain rates of oxygen isotope exchange. It is important to point out that diffusion and first-order reaction are difficult to distinguish as rate-controlling mechanisms using bulk exchange data alone (Cussler, 1984). Other experimental studies of oxygen isotope exchange between 02 gas and amorphous silica have shown that both diffusion of molecular 02 and first-order reaction between mobile 02 and network oxygen contribute to the overall exchange rate (Cawley and Boyce, 1988; Kalen et al., 1991). Because of uncertainties regarding the actual diffusing species for oxygen, the precise geometry of the glass grains, and the assumption of local isotope equilibrium between oxygen-bearing species in the rhyolite, we use the term apparent oxygen diffusion in the following discussions. Apparent oxygen diffusivities were calculated for short-

6

O

silica (2)

5

............ * ............ , ................................. , ....... Pt (2)

4 i

0

~

I

100

~

~

Pt(2)

~

~

I

200

,

,

,

-

~

"~'"" Silica Glass

-

1

,

300

Run Duration (day) FIG. 1. CO2-rhyoliticglass oxygen isotope fractionations determined in experiments at 650°C (Table 4) plotted vs. run duration; uncertainties are ___la and the kind of capsule and number of runs for each time are indicated next to the data symbol. Anomalously large values observed in short duration experiments approach equilibrium fractionation factors after about 100 days. The dashed line indicates the equilibrium oxygen isotope fractionation factor obtained by averaging the results of the 120 and 255 day experiments (Table 4, final column).

I

0.5

~

~

~

~

I

,

1.0 106 T-2 (K -2)

,

~

~

I

1.5

FIG. 2. COz-rhyoliticglass/melt oxygen isotope fractionation factors (reported as 1000 In aco2-r) determined in this study from runs of 100-255 day duration (Table 4, final column) plotted vs. T-2; uncertainties are _ hr. Also shown are oxygen isotope fractionation factor curves for CO2-silicaglass (Stolper and Epstein, 1991) and CO2-albiticglass/melt (Matthews etal., 1994) and a predicted curve for CO2-rhyolitic glass/melt based on these previous studies and weighted for the proportion of oxygen contained in normative silica (41%) and alkali feldspar (59%), which together constitute 96% of the total oxygen in the rhyolitic glass starting material.

Oxygen isotope fractionation between glass and CO2 T (°C) 850 I

1300 I o -15 Tt D

;°

t

Rhyolitic Mell

¢xl

g

650 I 1 bar C02

o

r-t Albitic Glass / Mel

_o -20

Silica Glass I

0.6

=

I

0.8

a

I

1.0

~

I

1.2

103 T -1 (K -1) FIG. 3. Arrhenius plot of apparent oxygen diffusion coefficients in rhyolitic glass/melt determined at 650-850°C and 1 bar CO2 results of short duration partitioning experiments (3-5 days, from Table 4); uncertainties are estimated to be less than or equal to the size of the data symbols (solid diamonds). The value at 850°C is a minimum because of sintering, which increased the grain size during the course of the run. Also shown are experimental determinations of apparent oxygen diffusion in the presence of CO2 at 1 bar for rhyolitic melt (Muehlenbachs and Kushiro, 1974), silica glass (Stolper and Epstein, 1991), and albitic glass/melt (Matthews et al., 1994) with 1~ uncertainties indicated.

term experiments involving rhyolitic glass/melt at 650 to 850°C using a series approximation for diffusion in a sphere (Crank, 1975). The resuks are plotted versus T -1 in Fig. 3 where they are compared with previous determinations of apparent oxygen diffusivil:y made in the presence of CO~ at 1 bar for rhyolitic melt (lVluehlenbachs and Kushiro, 1974), silica glass (Stolper and ]Epstein, 1991), and albitic glass/ melt (Matthews et al., 1994). This analysis assumes that the radii and surface areas of glass grains remained constant over the course of the experiments. We believe this is a reasonable assumption as the samples were of a restricted size range (45-75 #m) and were annealed at 600°C prior to the experiments. This clearly was not the case for experiments at 850°C, where simering increased the grain size of the glass during the experiment and the apparent oxygen diffusivity calculated is a minimum value. In previous experiments of the type described here, Stolper and Epstein ( 1991 ) and Matthews et al. (1994) inferred that dissolved CO2 was the principal carrier of oxygen in the silicate. The higher apparent oxygen diffusivities in rhyolitic glass/melt compared with silica glass shown in Fig. 3 is consistent with this sugge.stion because Ar, a neutral atom of similar size to CO2, is lolown to diffuse more rapidly in the former (Carroll, 1991; C~xroll and Stolper, 1991; Roselieb et al., 1992). However, the much higher apparent oxygen diffusivity in rhyolitic compared to albitic glass/melt cannot be readily explained in this way because Ar solubility and diffusivity are very similar in these two silicate glasses (Carroll, 1991 ). The results suggest that an oxygen-bearing species other than or in addition to CO2 was involved in the oxygen isotope exchange process. The most likely candidate

1969

is the small amount of H20 contained in the natural rhyolitic glass starting material used in this work (HzOtot,~ = 0.16 wt%; Table 1 ), but not detected in the synthetic silica and albitic glasses used in the earlier work. A series of exchange experiments with CO2 gas of 6180 = +42 and untreated (initial H2Ototal = 0.16 wt%) and dehydrated (initial H2Ototal = 0.09 wt%) rhyolitic starting materials were performed at 650°C for 5 h, 5 days, and 255 days to test the role of dissolved water in the rhyolitic glass on the rate of oxygen isotope exchange with CO2. The results of these experiments are listed in Table 5. Relatively large amounts of H20 (1.4 and > 2 #mol, corresponding to X (CO2) -< 0.95 in the gaseous run products) were measured in the Pt capsules at the conclusion of 5 h and 5 day runs with untreated rhyolite, consistent with partial dehydration of the glass during the course of the experiments. The capsule of the 255 day run with untreated rhyolite contained virtually no H20 in the vapor in spite of the glass having lost 88% of its initial dissolved water as determined by FTIR analysis (final H 2 O t o t a I = 0.02 wt% ), an indication that over time HzO in the vapor was reduced and H2 escaped from the capsule. The amounts of H20 vapor found in capsules at the conclusion of runs with dehydrated rhyolite were similar to those found for the partitioning experiments listed in Table 4. As with the untreated sample, approximately 90% of the initial dissolved water was lost from the dehydrated rhyolitic sample after 255 days (final H2Oto~a~= 0.009 wt%). The regular decrease in 6J80 of CO2 with increasing run duration illustrated in Fig. 4a reflects progressive oxygen isotope exchange between CO2 gas and rhyolitic glass. From Fig. 4a, it is clear that this process occurred more rapidly in the case of the experiments involving untreated rhyolitic glass than those with the dehydrated glass. This difference cannot be attributed to isotopic exchange between CO2 and variable amounts of H20 in the vapor during quenching of the runs. Such a process would enrich the JSO content of CO2 in experiments involving untreated rhyolitic glass relative to those with the dehydrated glass because of larger amounts of dehydrated water in the former and the positive CO2-H20 oxygen isotope fractionation factor (Chacko et al., 1991; Rosenbaum, 1993), opposite to the observed results. The above results prove that the rate of oxygen isotope exchange was systematically different between the two starting materials, even in runs lasting 255 days. Although a variety of properties could produce such effects (i.e., surface area, reactivity, or molecular structure), the most straightforward explanation is the difference in initial dissolved water content. This conclusion is independent of any model used to quantify the rate of oxygen isotope exchange. The data trends illustrated in Fig. 4a are, nonetheless, consistent with control of oxygen isotope exchange by diffusion in the glass. Apparent oxygen diffusion coefficients calculated for each run are listed in Table 5 and plotted in Fig. 4b. The results for runs with untreated rhyolite are indistinguishable from those derived from short duration partitioning experiments at 650°C (Doxy = 2 - 5 × 10 -17 m 2 S l ) ; this is the expected result as both sets of experiments utilized starting materials with similar initial water contents. In contrast, experiments with dehydrated rhyolitic glass as the starting material yield

1970

J.M. Palin, S. Epstein, and E. M. Stolper

Duration (hr)

101

102

(a)

103

104

35 40

30

6o

~

20

8o

15 10

IO0 II1

I

I I IIIIII

I

I IIIIIll

I

I I IIIIll

I

I I IP

II I

I

I I IIIII I

I

I I IIIII I

I

I I IIIII I

I

I I I I

-16

(b)

Untreated 0.16 wt.% H201o~

I t/)

-17 o

121 o

Dehydrated 0.09 wt.% H2Ole~l

-18 10-1

~

45-75 ktm

]

160-210 ~m

1

100 101 102 Duration (day)

FIG. 4. (a) Final 6180 values of CO2 from exchange rate experiments with untreated (diamond, initial H2Otota] = 0.16 wt%) and dehydrated (inverted triangle, initial H2Otota] = 0.09 wt% ) rhyolitic glass at 650°C as a function of run duration (see Table 5). Also shown are predictions of a model for diffusion-controlled isotope exchange with an initial 6180 CO2 value of +42.44, an equilibrium CO2-rhyoliticglass fractionationfactor of 4.62 (Table 4), a rhyolite to CO2 oxygen ratio of 55, spherical grains with mean diameters of 185 #m, and apparent oxygen diffusivities ranging from 10 -]5 to 10 -18 (m2 s-l). The final 6 1 8 0 of CO2 at 100% equilibration is + 12.97 for these parameters. (b) Apparent oxygen diffusivities in rhyolitic glass at 650°C determinedin exchangerate (closed symbols, Table 5) and short duration partitioning (open symbol, Table 5) experiments with estimated uncertainties indicated.

apparent oxygen diffusion coefficients that are lower by a factor of 4 to 9 ( O o x y = 3 - 11 × 10 -18 m 2 s - l ) . 4. DISCUSSION

4.1. Oxygen Isotope Partitioning in Rhyolitic Glass and Melt As noted previously, the reduced partition function ratio for CO2 vapor is one of the most accurately known (Bottinga, 1968; Richet et al., 1977; Chacko et al., 1991 ). In combination with experimental data on oxygen isotope partitioning between CO2 and a condensed phase, this reduced partition function ratio can be used to derive those of the condensed phase. Using this approach, Stolper and Epstein (1991) and Matthews et al. (1994) showed that the reduced partition function ratios of silica glass and albitic glass/melt could be

fit by making small adjustments (on the order of 2 - 3 % ) to the expressions for the reduced partition function ratios for quartz and albite given by Clayton and Kieffer (1991). In Fig. 2, we show that the COz-rhyolitic glass/melt oxygen isotope fractionation data of this study can be described well between 650 and 950°C if we assume that the reduced partition function ratio for rhyolitic glass/melt is approximated by a linear combination of the reduced partition function ratios for silica glass (Stolper and Epstein, 1991 ) and albitic glass/melt (Matthews et al., 1994), if these are weighted in proportion to the oxygen contained in normative silica and alkali feldspar. Silica and albitic glasses are appropriate models for oxygen isotope partitioning in rhyolitic glass as all are fully polymerized silicate networks based on an average ring size of six tetrahedra (Taylor and Brown, 1979). Albitic glass is a plausible model for normative alkali feldspar in rhyolitic glass because of the lack of significant oxygen isotope fractionation between albite and albitic glass (Matthews et al., 1994) and between albite and potassium feldspar under natural and experimental conditions (O'Neil and Taylor, 1967; Taylor, 1968). Divergence between predicted and experimental values at 550°C may reflect uncertainties due to the large extrapolation required in the high temperature results (->750°C) of Matthews et al. (1994) on albitic glass or the possibility of more complex partitioning behavior in silicate glasses at lower temperatures. The good agreement between measured and predicted fractionation factors shown in Fig. 2 indicates that with respect to oxygen isotope partitioning, felsic glasses and melts are well modeled as ideal mixtures of their major silicate components, silica and feldspar. Natural felsic glasses containing quartz phenocrysts can be used to test both the applicability of our experimental results on rhyolite and the validity of the mixing model described above to other compositions. Figure 5 shows a compilation of measured oxygen isotope fractionations between quartz phenocrysts and coexisting unhydrated felsic glass in rhyolite from Yellowstone (Hildreth et al., 1984), granodiorite from Crater Lake (Bacon et al., 1989; Bacon, 1992), and dacite from Mt. Lassen (Taylor, 1968 ). These data are plotted versus temperatures which were calculated on the basis of compositions of coexisting Fe-Ti oxides as given in the original references for the Yellowstone and Crater Lake samples or calculated from oxygen isotope fractionations between quartz and plagioclase phenocrysts for the Mt. Lassen samples. Also plotted are equilibrium oxygen isotope fractionation curves for silica glass, rhyolitic glass/melt, albitic glass/melt (coincident with albite), and anorthite relative to quartz. The quartzglass pairs from the Yellowstone rhyolites have oxygen isotope fractionations that are consistent with our experimental calibration. Partially fused granodiorite blocks from Crater Lake and dacites from Mt. Lassen show somewhat larger quartz-glass fractionations, consistent with their less silicarich normative compositions. Our data can also be used to evaluate oxygen isotope ratios of volcanic gases in equilibrium with rhyolitic melts. The most abundant magmatic volatiles are H20 and CO2, and Fig. 6 shows equilibrium oxygen isotope fractionation curves for CO2, H20, and their mixtures relative to rhyolitic glass and melt as functions of T -2. This figure emphasizes

Oxygen isotope fractionation between glass and CO2

T (°C) 950 i

1

550

750 i

.[~

/

0 0

' ~"

oo o

Silica Glass

Quartz Rhyolitic Glass / Molt

-1

AIbit0

-2 -3

, 0.0

0.5 106 T-2

,

Anorthite

1.0

1.5

(K-2)

FIG. 5. Measured oxygen isotope fractionations between quartz phenocrysts and coexisting glass matrix in rhyolites from Yellowstone (squares) (Hildreth et al., 1984), in partially fused granodiorite blocks from Crater Lake (triangles) (Bacon et al., 1989; Bacon, 1992), and in dacites from Mt. Lassen (circles) (Taylor, 1968) plotted versus temperature (as T-a). Temperatureswere taken from the original source (ba~,;edon compositions of coexisting FeTi oxides) for samples from Yellowstone and Crater Lake and were calculated using oxygen isotope data for quartz and plagioclase phenocrysts and fractionation fac,tors of Clayton et al. (1989) for samples from Mt. Lassen. Equilibrium oxygen isotope fractionation curves for silica glass, rhyolitic glass/melt, albitic glass/melt (coincident with albite), and anorthite relative to quartz are based on experimental calibrations for CO2-rhyolitic glass/melt from the present work, CO2-silicaglass (Stolper and Epstein, 1991; Matthews et al., 1994), CO2-albite and albitic glass/melt (Matthews et al., 1994), CO2-calcite (Chacko et al., 1991; Rosenbaum, 1994), and calciteanorthite and calcite-quartz (.Claytonet al., 1989). the significant difference between CO2-rhyolitic melt and H20-rhyolitic melt oxygen isotope fractionations at magmatic temperatures. As discussed by Matthews et al. (1994), a prediction based on these relations is that the bulk oxygen isotope ratio of volcanic gas evolved from rhyolitic magma should decrease as the H20/CO2 ratio increases through progressive degassing. Note also that the CO2-rhyolitic melt and H/O-rhyolitic melt oxygen isotope fractionation factors are opposite in sign and increase in magnitude with decreasing temperature. As a consequence, CO2-rich gases exsolved from a rhyolitic magma would have the potential to generate 180-depleted alteration zones as they migrate down temperature through crystallized portions of the same body. H 2 0 rich gases would tend to have the opposite effect. We speculate that zones of focused degassing above rhyolitic magma chambers may exhibit progressive changes in mineralogic and stable isotopic alteration as a consequence of these relations. Early formed alteration zones should reflect passage of H/O-poor gas compositions and have 6 xsO values slightly lower than that of the magma, whereas later formed alteration zones should record H20-rich gas compositions and higher 6180 values.

4.2. Oxygen Diffusion in Rhyolitic Glass and Melt Apparent oxygen diffusivities in rhyolitic glass/melt determined from our short duration ( 3 - 5 day) partitioning and exchange rate (5 h - 2 5 5 day) experiments are compared in

1971

Fig. 7 with a model for diffusion of single oxygen-bearing species (molecular CO2 in Fig. 7a; molecular H20 in Figure 7b) which is locally in isotopic equilibrium with network oxygen of the glass (Zhang et al., 1991a). As shown in Fig. 7a, the apparent oxygen diffusivities are approximately 3 orders of magnitude higher than expected if molecular CO/ had been the sole diffusing species, based on data for CO2 solubility, speciation, and diffusion in rhyolitic glass and melt (Blank, 1993; Blank et al., 1993). On the other hand, Fig. 7b shows that the apparent oxygen diffusivities are close to predicted values for rhyolitic glass/melt with 0.1 wt% H2Otota!in which molecular H20 is the mobile species responsible for the transport of oxygen through the glass or melt (Zhang et al., 1991b). This is similar to the water content of the rhyolitic glass starting material used in our experiments and, therefore, lends support to the hypothesis that oxygen transport is mainly accomplished by diffusion of mobile water molecules. It is important to emphasize that at the low total water contents of our samples, most of the water is dissolved as hydroxyl ( O H - ) (Stolper, 1989; Silver et al., 1990) and is effectively immobile for oxygen transport (Zhang et al., 1991b). For example, at 650°C, rhyolitic glass with 0.1 wt% H2Ototal is estimated to contain only 63 ppm of dissolved molecular H20 using parameters of Zhang (1991b) in the equilibrium speciation model of Stolper (1989) and Silver et al. (1990). This small quantity of molecular H20 is nevertheless capable of transporting the bulk of the oxygen because its diffusivity in the rhyolitic glass (DH2o ~ 10 -12 m 2 s - x) is several orders of magnitude faster than that of other oxygen-bearing species (OH , molecular CO/, CO 2-, or network 0 2- ) in the glass under the conditions of the experiments. The hypothesis that oxygen transport in our experiments is controlled by the mobility of hydrous species (i.e., molecular H 2 0 ) is further supported by systematic differences in the

T (°C) 950 '

6

750 '

550 ' C02

5

0.8

I,,-

>" rr

4

0.6

, (n

3

(~

2

¢}

1

--= (3 oo

-1

o

•i--

0.4

X(C02) 0.2

0

~

-2

H20

'

0.0

0.5 106 T-2

1.0

1.5

(K-2)

FIG. 6. Equilibrium oxygen isotope fractionation curves for CO2, H20, and their mixtures relative to rhyolitic glass/melt plotted as a function of T-2 based on experimentalcalibrations for CO2-rhyolitic glass/melt from the present work, CO2-albite (Matthews et al., 1994), and H20-albite (O'Neil and Taylor, 1967; Matsuhisa et al., 1979; Matthews, 1980). Mixed H20-CO2gas compositionsare given as molar X(CO2) = CO2/(CO2 + H20).

1972

J.M. Palin, S. Epstein, and E. M. Stolper

T (°C)

850 I

750 I

(a)

650

I fco2 (b)

d 8' -2o

that small amounts of H20 and/or CO2 may control diffusive exchange of oxygen isotopes in silicate materials under both experimental and natural conditions. As discussed at greater length by Zhang et al. (1991a), this hypothesis is supported by linear relations between apparent oxygen diffusivity and H20 and CO2 fugacity found in oxygen isotope exchange experiments involving quartz (Giletti and Yund, 1984; Farver and Yund, 1991; Sharp et al., 1991 ) and alkali feldspar (Farver and Yund, 1990). Interpretation of these data has been complicated by inadequate quantitative relations for the solubility, speciation, and diffusivity of important oxygen-bearing volatile species in silicate minerals.

(b)

5. CONCLUSIONS

5 I

-15

i~

o.1

O.Ol -

o ~,/o "20

I

0.8

,

I

,

1.0

I

1.2

103 T -1 (K -1) FIG. 7. Arrhenius plots of apparent oxygen diffusivity in rhyolitic glass/melt at 650-8500(2 and 1 bar CO2 determined in short duration partitioningand exchange rate experimentswith untreated (diamond, initial H2Ototal = w t % , Table 4 and 5) and dehydrated (inverted triangle, initial H2OtotaI = 0 . 0 9 w t % , Table 5) rhyolitic glass; uncertainties less than or equal to size of data symbols. (a) Contours show predicted apparent oxygen diffusion coefficients for the case of dissolved molecular CO2 as the sole diffusing species in the glass or melt as a function of CO2 fugacity based on experimental data for CO: solubility and diffusion in rhyolitic glass and melt (Blank, 1993; Blank et al., 1993). (b) Contours show predicted apparent oxygen diffusivities for the case of dissolved molecular H20 as the sole diffusing species in the glass or melt as a function of total dissolved water content based on experimental data for H20 solubility, speciation, and diffusion in rhyolite (Stolper, 1989; Silver et al., 1990; Zhang et al., 1991b).

rate of oxygen isotope exchange in dehydrated rhyolitic glass relative to untreated glass in exchange rate experiments (Fig. 4). The untreated glass had 3 to 5 times more dissolved molecular H20 than the dehydrated glass at 650°C according to calculations based on total dissolved water contents of glass samples measured by FTIR before and after exchange experiments (Table 5). This difference compares favorably with the 4 to 9 times greater apparent oxygen diffusivity in the untreated glass relative to the dehydrated glass estimated from these runs (Table 5). These findings agree with the predictions of the model of Zhang et al. ( 1991 a) for diffusion of a multispecies component such as oxygen; i.e., apparent diffusivity is equal to the sum of the diffusive fluxes (the product of concentration and diffusivity) of all species. In the present study, the diffusive flux of oxygen carried by dissolved molecular H20 appears to have dominated those of all other oxygen species and thus controlled the rate of oxygen isotope exchange between CO2 gas and rhyolitic glass and melt. The important implication of these studies is

1 ) Exchange experiments conducted at 550-950°C for 100 days or longer yield equilibrium oxygen isotope fractionation factors between CO2 and rhyolitic glass and melt. Oxygen isotope partitioning is well described by predictions based on experimentally calibrated CO2-silica glass (Stolper and Epstein, 1991 ) and CO2-albitic glass and melt (Matthews et al., 1994) fractionation factors and the proportions of oxygen in the most abundant normative constituents, silica and alkali feldspar. Oxygen isotope partitioning in felsic glasses and melts can thus be modeled by linear combination of endmember silicate constituents. The experimental data are consistent with oxygen isotope fractionations observed between quartz phenocrysts and coexisting unaltered glass in felsic volcanic rocks. 2) Exchange experiments of short duration ( 3 - 5 days) exhibit anomalously large CO2-rhyolitic glass and melt oxygen isotope fractionations. These effects are similar to those observed in previous studies using similar techniques and remain to be explained definitively. 3) Rates of CO2-rhyolitic glass oxygen isotope exchange in the experiments are three orders of magnitude faster than predicted if diffusion of dissolved molecular CO2 in glass were the rate-controlling mechanism. These rapid exchange rates are, however, consistent with control of oxygen isotope exchange by diffusion of molecular H20 dissolved in the glass. Comparison of the rates of oxygen exchange between CO2 gas and untreated and dehydrated rhyolitic glasses support this interpretation quantitatively. We conclude that, as proposed by Zhang etal. (1991a), diffusive exchange of oxygen isotopes in rhyolitic glass and melt, and probably other polymerized silicate materials, is controlled by the diffusive fluxes of dissolved oxygen-bearing volatile species rather than diffusion of network oxygen under all but the most volatile-poor conditions.

Acknowledgments--This work was supported by DOE grant DEFG03-85ER13445 (to EMS) and NSF grant EAR 9303975 (to SE and EMS). We thank A. Matthews for refining the experimental procedures and S. Newmanfor generouslyproviding the bTIR analyses. JMP thanks C. Martin, J. Blank, G. Holk, E. Dent, J. Kubicki, J. Beckett, S. Newman, and C. Tacker for stimulating discussions and assistance in the lab. P. Ihinger, P. Richet, and J. Rosenbanm provided thorough and helpful reviews. Division of Geological and Planetary Sciences contribution 5538. Editorial handling: B. E. Taylor

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