Oxygen isotope fractionation factors involving cassiterite

0 downloads 0 Views 156KB Size Report
tion at St. Michael's Mount, Cornwall, England. Inst. Mining. Metallurgy Trans. B 88, 57– 60. Northrop D. A. and Clayton R. N. (1966) Oxygen isotope fractionation.
Geochimica et Cosmochimica Acta, Vol. 69, No. 5, pp. 1301–1305, 2005 Copyright © 2005 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/05 $30.00 ⫹ .00

doi:10.1016/j.gca.2004.09.0020

Oxygen isotope fractionation factors involving cassiterite (SnO2): II. Determination by direct isotope exchange between cassiterite and calcite GUIXING HU,1,3 ROBERT N. CLAYTON,1,2,3,* VENIAMIN B. POLYAKOV,4 and SERGEY D. MINEEV4 1 Enrico Fermi Institute, University of Chicago, Chicago, IL 60037, USA Department of Chemistry, University of Chicago, Chicago, IL 60637, USA 3 Department of the Geophysical Sciences, University of Chicago, Chicago, IL 60637, USA 4 V. I. Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences, Kosygin Street 19, Moscow, 119991, Russia 2

(Received January 6, 2004; accepted in revised form September 9, 2004)

Abstract—Direct oxygen isotope fractionation between cassiterite and calcite has been investigated experimentally at 15 kbar with temperature ranging from 800 to 1000°C. Combined with the quartz-calcite fractionation measured with the same technique (Clayton et al., 1989), the calcite-cassiterite and quartzcassiterite oxygen isotope fractionations can be expressed as: ⌬CaCO3⫺SnO2 ⫽ 4.14 ⫻ 106T⫺2 ⌬SiO2⫺SnO2 ⫽ 4.52 ⫻ 106T⫺2 at temperatures higher than 600°C. These calibrations are in good agreement with those obtained from heat capacity and X-ray resonant data. The theory of using X-ray resonant data to calculate the reduced partition function ratios of transition elements that have at least one Mössbauer-sensitive isotope is evaluated with the current experimental result. Copyright © 2005 Elsevier Ltd quartz-cassiterite) fractionation factors. A recent study indicates that mineral-pair fractionations determined with this hydrothermal exchange technique may not be reliable because of the effect of dissolved minerals on the isotope partitioning behavior of fluids (Hu and Clayton, 2003). The same study also demonstrated that the use of carbonate rather than water as the common isotope exchange medium provides correct mineral-pair fractionations. In this study, we experimentally calibrated fractionation factors for cassiterite by direct oxygen isotope exchange between calcite and cassiterite under anhydrous conditions. These data can be combined with the existing experimental database of mineral-calcite fractionation factors (Clayton et al., 1989; Chiba et al., 1989; Chacko et al., 1996) to arrive at a large number of fractionation factors involving cassiterite, including the quartz-cassiterite fractionation factor applicable to the genesis of tin deposits. These experimental data can also be compared to fractionation factors for cassiterite calculated in the study of Polyakov et al. (this issue, paper 1 of this series). The calculations use a new method suggested by Polyakov and Mineev (2000), which enables calculation of reduced partition function ratios of light elements, such as oxygen and sulfur, in compounds that consist of two elements, one of which has a Mössbauer-sensitive isotope. Cassiterite is such a compound in that it consists of two elements, Sn and O, and one of those elements has a Mössbauer-sensitive isotope (119Sn). Thus, the experimental data reported in the present study provide an independent test of the new Mössbauer-derived theory for calculating fractionation factors described in paper I.

1. INTRODUCTION

There are two objectives of this study: (1) evaluate the theory of using Mössbauer spectroscopy or inelastic X-ray resonant data to calculate the reduced partition function ratios of transition elements that have at least one Mössbauer-sensitive isotope (Polyakov and Mineev, 2000), and (2) provide reliable calcite-cassiterite and quartz-cassiterite oxygen isotope fractionation factors using an anhydrous carbonate-exchange technique. As a major tin-bearing mineral, cassiterite plays a crucial role in the geochemistry and economic geology of tin. Hydrothermal fluids are very important in the formation of tin deposits. Oxygen isotopes have proven to be a useful tool for providing temperature information and determining the origin of geological fluids. However, the application of oxygen isotopes in the study of tin deposits has been limited due to the lack of reliable oxygen isotope fractionation factors involving cassiterite. Hydrothermal experimental studies (e.g., Sushchevskaya et al., 1985; Zhang et al., 1994), empirical calibration (Kelly and Rye, 1979; Patterson et al., 1981; Borshchevskiy et al., 1983; Sun and Eadington, 1987; Alderton, 1989; Strauch et al., 1994) and a semiempirical calculation (Zheng, 1991) have previously been used to calibrate the oxygen isotope fractionation factors involving cassiterite. However, a range of results was reported by these studies, with each study neither confirming nor disproving earlier results. It should be stressed that most previous experimental measurements of oxygen isotope fractionation factors involving cassiterite were based on separate mineral-water isotope exchange experiments, which were then combined to yield mineral-mineral (e.g.,

2. EXPERIMENTAL METHODS The starting cassiterite is a commercial product of “Alfa Aesar” (referred as the AA sample in this paper). It was heated at 1000°C for 73 h before being used for the experiments. A loss in weight of 3.84% was observed, resulting from water release. No weight change was

* Author to whom correspondence should be addressed (r-clayton@ uchicago.edu). 1301

1302

G. Hu, R. N. Clayton, V. B. Polyakov, and S. D. Mineev Table 1. Oxygen isotopic compositions of starting material.

Mineral

␦ OSMOW (‰)

SD (1␴) of ␦18OSMOW (‰)

Cassiterite (AA) Calcite (HCC) Calcite (WC4)

14.54 (4) 17.72 (2) 14.10 (2)

0.15 0.09 0.06

18

NOTE: Number in parentheses indicates the number of replicate analyses.

removed from the second portion of the experimental product by first ultrasonically agitating it for 5 min in 2 N acetic acid solution and then letting it react with the acid solution for 1 h at room temperature. The solid residue was recovered by washing repeatedly with distilled water and then dried. The recoveries of cassiterite were larger than 99%. The oxygen isotope composition of cassiterite was then measured with the conventional bromine pentafluoride method. Control experiments showed that cassiterite did not change its oxygen isotope composition in the acetic acid treating process. X-ray studies showed that there was no change in mineral structure between starting minerals and experimental products. 3. EXPERIMENTAL RESULTS

detected on further heating. An X-ray study confirms the cassiterite structure of the AA sample. The average grain-size of AA sample is ⬃10 –15 ␮m. The heated AA sample was used both in this study and in determining the heat capacity (Paper I). The two starting calcites are natural samples. They were prepared by grinding in an agate mortar, followed by sedimentation in distilled water for 30 min. Scanning electron microscopy (SEM) study showed that the size of the calcite recovered from the suspended fraction that was used in this study ranges from 1 to 10 ␮m. Starting calcite and cassiterite were mixed with a 1:1 oxygen atomic ratio under reagent-grade acetone in a small agate mortar three times to dryness. Then the mixture was baked at 300°C for 5 h. Oxygen isotope compositions of the starting materials are shown in Table 1. The oxygen isotope composition of cassiterite was measured by the conventional bromine pentafluoride method (Clayton and Mayeda, 1963) and the oxygen isotope compositions of calcite samples were measured with the standard phosphoric acid method (McCrea, 1950). All cassiterite-calcite exchange experiments were carried out in a 1.91-cm piston-cylinder apparatus with a graphite heater and a NaCl pressure medium. Temperature was controlled and measured with W-3% Re vs. W-25% Re thermocouples. The accuracy of temperature measurements is estimated to be better than ⫾5°C. Pressures were measured with a Heise Bourdon tube gauge with an accuracy of about ⫾200 bar. In each experiment, a charge of 25 to 30 mg mixture of starting materials was used. The exchange experiments and the analyses of experimental products followed the procedure developed by Clayton et al. (1989). After cutting open the platinum capsule, experimental products were recovered, crushed and ground. A small portion of the product was reacted with 100% phosphoric acid for 1 d at 25°C for determination of the oxygen isotope composition of calcite. Calcite was

The cassiterite-calcite exchange experiments were conducted at 15 kbar from 800 to 1000°C. The experimental conditions and results are shown in Table 2. The isotope exchange reached equilibrium in runs at 900°C. The partial exchange results at 1000°C and 800°C are shown on the Northrop-Clayton plots (Fig. 1) (Northrop and Clayton, 1966). The extent of exchange is 85% at 800°C and 91% at 1000°C, so that errors in the extrapolation procedure to derive the equilibrium fractionation factors cannot be large. For comparable grain sizes and exchange duration, the extent of exchange between calcite and cassiterite is larger than that between calcite and rutile (Chacko et al., 1996), which has the same structure as cassiterite. This suggests that cassiterite is less isotopically refractory than rutile. The analytical uncertainty for the fractionation factor at all temperature is within ⫾ 0.20‰ (1␴). The fractionations as a function of temperature are shown in Figure 2. At sufficiently high temperature, the fractionation between two anhydrous phases is linear in the graph of ln␣ vs. 1/T2 (T in K), passing through the origin (Bigeleisen and Mayer, 1947). “Sufficiently high” is defined by hc␻/KT⬍2, where ␻ is the highest vibrational frequency in the two phases. A linear extrapolation to 600°C of the high temperature results in anhydrous carbonate-exchange experiments may be reasonable (Clayton et al., 1989). The regression line of the current experimental results is shown in Figure 2. Also shown in Figure 2 is the calculated calcite-cassiterite fractionation (Paper I) over the same temperature range. A best-fit straight line to the experimental data, forced

Table 2. Cassiterite-calcite exchange at 15 kbar. Run no.

T (°C)

t (h)

Calcite

Final ␦calcite (‰)

Final ␦cassiterite (‰)

⌬fa (‰)

␦18OMBb (‰)

Exchangec (%)

1 2 3

1000 1000 1000

72 72 72

WC4 HCC WC4

15.45 17.40 15.52

12.97 14.60 13.01

2.45 2.76 2.47

⫺0.11 ⫺0.13 ⫺0.05

91

2.73

4 5 6

900 900 900

120 120 120

WC4 HCC WC4

15.88 17.50 15.92

12.69 14.49 12.54

3.15 2.96 3.33

⫺0.03 ⫺0.13 ⫺0.09

100

3.15

7 8 9

800 800 800

183 183 183

WC4 HCC WC4

15.75 17.70 15.73

12.85 14.38 13.00

2.86 3.30 2.69

⫺0.02 ⫺0.09 ⫹0.05

85

3.36

a

⌬f ⫽ Final value of 1000 ⫻ ln





1⫹

1⫹

␦18Ocalcite 1000



␦18Ocassiterite



.

1000 ␦18OMB ⫽ Material balance ⫽ ␦18Omean products ⫺ ␦18Omean reactants. Value should be 0 ⫾ 0.2‰. c Exchange extent and equilibrium fractionation factors are calculated on the basis of the method of Northrop and Clayton (1966). b

⌬ec (‰)

18

O fractionation between cassiterite and calcite

1303

Fig. 2. Oxygen isotope fractionation between calcite and cassiterite.

involved the fractionation between cassiterite and quartz. Clayton et al. (1989) measured the fractionation between quartz and calcite. Their results were confirmed by direct quartzcalcite exchange in both anhydrous and three-phase hydrothermal systems (Hu and Clayton, 2003). Combining the current cassiterite-calcite exchange results and the quartz-calcite exchange results of Clayton et al. (1989), the fractionation between quartz and cassiterite can be constructed and is shown in Figure 3. The quartz-cassiterite fractionations by Mössbauerderived calculation (Paper I), hydrothermal exchange experiments (Zhang et al., 1994), empirical calibration (Borshchevskiy et al., 1983; Alderton, 1989) and semiempirical calculation (Zheng, 1991) are also shown in Figure 3 for comparison.

Fig. 1. Partial-exchange (Northrop and Clayton, 1966) plots for oxygen isotope exchange between calcite and cassiterite. Figure 1a shows the result at 1000°C and 15 kbar for 72 h. Figure 1b shows the result at 800°C and 15 kbar for 183 h. ⌬i ⫽ 1000ln␣Calcite-Cassiterite at the start of the experiment. ⌬f ⫽ 1000ln␣Calcite-Cassiterite at the end of the experiment. The y-intercept gives an extrapolated equilibrium fractionation of 2.73‰ and ⫺1/slope gives an exchange extent of 91% at 1000°C in Figure 1a. The y-intercept corresponds to an extrapolated equilibrium fractionation of 3.36‰ and ⫺1/slope corresponds to an exchange extent of 85% at 800°C in Figure 1b.

to go through the origin, has a slope of 4.14. The two lines agree well within experimental and calculation uncertainties. Most of previous studies on the oxygen isotope of cassiterite

Fig. 3. Oxygen isotope fractionation between quartz and cassiterite.

1304

G. Hu, R. N. Clayton, V. B. Polyakov, and S. D. Mineev 4. DISCUSSION AND SUMMARY

4.1. Oxygen Isotope Fractionation Involving Cassiterite As shown in Figure 3, the agreement between the quartzcassiterite fractionations obtained by carbonate-exchange method (labeled as Experiment) and by Mössbauer-derived calculation (labeled as Paper I) is good. Over the temperature range from 600 to 1000°C, the difference between the two lines is ⬍0.15‰, within the experimental uncertainty. The current quartz-cassiterite fractionation results are far from the results of previous studies on this system. Zhang et al. (1994) measured oxygen isotopic exchange in the quartz-cassiterite-water system from 250 to 500°C, using amorphous SnO2 and silica gel as starting materials. The exchange experiments were conducted in hydrothermal autoclaves with a gold lining. The fractionations at 400°C and 500°C were measured by direct exchange in a three-phase hydrothermal system. The fractionations at 370°C, 310°C and 250°C were derived by combining the fractionations in separate quartz-water and cassiterite-water systems, using silica gel and amorphous SnO2 as starting materials, respectively. Some limitations of the low-pressure hydrothermal exchange technique were discussed by O’Neil et al. (1969) and Clayton et al. (1972) such as (1) large temperature gradient across the long sample tube causes large uncertainty in the determination of exchange temperature, (2) premature exchange during the slow heating period may not be erased in exchanges later at relatively low temperature, (3) quench product may add error to the measurements of experimental products. Furthermore, kinetic fractionation may occur in exchange experiments using noncrystalline starting materials. The kinetic effect may be so large that it drives the fractionation past equilibrium (Matsuhisa et al., 1978). Even though water with different oxygen isotope compositions was used as starting material, the exchange may pass equilibrium and approach the equilibrium from the same direction. Moreover, Hu and Clayton (2003) showed that the combination of two mineralwater fractionations may not give reliable fractionations between the two minerals because of the oxygen isotope salt effects of dissolved minerals. Therefore, it cannot be concluded that the quartz-cassiterite fractionations at 250°C, 310°C and 370°C measured by Zhang et al. (1994) reached equilibrium. In the three-phase hydrothermal exchange experiments carried out by Zhang et al. (1994), it is not known whether the exchange rate of quartz-water and that of the cassiterite-water are the same. If the exchange rates are different, the fractionation between the two solid phases may pass equilibrium as observed by Zheng et al. (1999) in the forsterite-calcite-fluid system. The contribution of a kinetic effect in using noncrystalline starting materials further complicates the determination of equilibrium in their three-phase hydrothermal experiments. Mass balance calculation shows that fractionation factors between quartz and water were negative (lower than ⫺20‰) in the final stages of the three-phase exchange experiments conducted by Zhang et al. (1994), demonstrating that the oxygen isotope equilibrium was not achieved in the experiments involving quartz-cassiterite isotope exchange. The two empirical calibrations by Borshchevskiy et al. (1983) and Alderton (1989) gave quartz-cassiterite fractionations much smaller than those measured in the current study.

The data used in their calibration are very scattered. At the same temperature, the quartz-cassiterite fractionations from different natural assemblages used by Alderton (1989) may differ by 2.7‰. Alderton (1989) realized that his regression line was not a perfect fit, due to the lack of complete quartzcassiterite isotopic equilibrium in the natural assemblages, confirming the studies on fluid inclusions (e.g., Moore and Moore, 1979) that indicated that the apparently coexisting quartz and cassiterite actually formed from fluids of different temperature and salinity. Further detailed discussion of the application of current calibration to natural assemblages is beyond the scope of this study. Generally, it is not reliable to calibrate fractionation based on natural samples. On the contrary, the combination of well designed experimental measurements and experimentally confirmed theoretical study should be used to investigate the temperature and material source information and/or any disequilibrium phenomena in natural assemblages. The semiempirical incremental method (Zheng, 1991) also gives quartz-cassiterite fractionation that is much smaller than the current result. It is not surprising that the quartz-cassiterite fractionations calculated by Zheng (1991) differ from the current results, since his calculation on rutile, which has the same structure as cassiterite and was treated in a same manner in the increment method, also gives smaller quartz-rutile fractionation than the experimentally calibrated quartz-rutile fractionation (Clayton et al., 1989; Chacko et al., 1996). The good agreement between the experimental equilibrium fractionation factors obtained in this study and those obtained by the theoretical treatment of heat capacity and Mössbauer data (Paper I) in the temperature range above 600°C allows us to recommend the theoretical calibrations of the equilibrium isotope fractionation involving cassiterite for use at lower temperatures, down to 400 K. 103 ln ␤SnO2 ⫽ 7.176x ⫺ 0.07369x2 ⫹ 0.0008026x3, x ⫽ 106 ⁄ T2 These calculations of oxygen reduced partition function ratios for cassiterite can be combined with previously published calculations of reduced partition function ratios for calcite and quartz (Clayton and Kieffer, 1991) to arrive at calcite-cassiterite and quartz-cassiterite fractionation factors (Figs. 2 and 3). 103 ln␣CaCO3⫺SnO2 ⫽ 4.607x ⫺ 0.3463x2 ⫹ 0.01500x3, x ⫽ 106 ⁄ T2; 10 ln␣SiO2⫺SnO2 ⫽ 4.942x ⫺ 0.2963x ⫹ 0.01150x , 3

2

3

x ⫽ 106 ⁄ T2 . 4.2. Evaluation of the Use of X-ray Resonant Data for Calculating Stable Isotope Fractionation Factors As shown in Figure 3, the calculation based on X-ray resonant data gives quartz-cassiterite fractionations that are in good agreement with the experimentally measured quartz-cassiterite fractionations using the carbonate-exchange technique. The O-isotope ␤-factors given by Polyakov et al. (this issue, paper 1 of this series) are derived from the O-sublattice kinetic energy of cassiterite, which is calculated by subtracting the Sn-sublattice kinetic energy of cassiterite (obtained from X-ray resonant

O fractionation between cassiterite and calcite

data) from the total kinetic energy of cassiterite (obtained from heat capacity data). The O-isotope ␤-factors and Sn-isotope ␤-factors are calculated in exactly the same manner with no assumption involved. Therefore, if the calculated O isotope ␤-factors are correct (as implied by the experiments), then the Sn sublattice kinetic energy determined using the vibration density of states (VDOS) approach must also be correct, and, by implication, the Sn ␤-factors calculated with this approach are correct as well. The excellent agreement between the experimentally determined oxygen isotope fractionation factors reported in the present study and those calculated using the VDOS method strongly supports the validity of this method in the calculation of isotopic reduced partition ratios. 4.3. Summary This study provides new measurements on the oxygen isotope fractionation involving cassiterite using carbonate-exchange experiments. The calcite-cassiterite fractionations measured in this study are in good agreement with calculations on the basis of heat capacity and X-ray resonant data. This supports the validity of the approach using Mössbauer or X-ray resonant data for calculating the reduced partition function ratios of heavy elements that have at least one Mössbauersensitive isotope, along with the ratios in light elements to which they are chemically bound. Acknowledgments—We are particularly grateful to Dr. R. Mendybaev for assistance in establishing and accomplishing the heating procedures of cassiterite samples. We thank Toshiko K Mayeda for assistance in the isotope laboratory. G. Hu especially thanks Toshiko K. Mayeda for her many years mentoring in the laboratory and in daily life and dedicates this work to her. A significant part on the present study was made during the stay of V. B. Polyakov in the Enrico Fermi Institute as a research scholar supported by U.S. National Science Foundation grant EAR 98-15338. Additional NSF support was provided by grant EAR 0126185. Associate editor: T. Chacko REFERENCES Alderton D. H. M. (1989) Oxygen isotope fractionation between cassiterite and water. Mineral. Mag. 53, 373–376. Bigeleisen J. and Mayer M. G. (1947) Calculation of equilibrium constants for isotopic exchange reactions. J. Chem. Phys. 13, 261–267. Borshchevskiy Y. A., Borisova S. L., Zakharova O. Y., Lugov S. F., Makeyev A. M., Podolskiy A. M., and Politov V. K. (1983) Oxygen-isotope systematics of the tin deposits of the northeast USSR. Int. J. Geol. Rev. 25, 107–116. Chacko T., Hu X., Mayeda T. K., Clayton R. N., and Goldsmith J. R. (1996) Oxygen isotope fractionations in muscovite, phlogopite and rutile. Geochim. Cosmochim. Acta 60, 2595–2608. Chiba H., Chacko T., Clayton R. N., and Goldsmith J. R. (1989) Oxygen isotope fractionation involving diopside, forsterite, mag-

1305

netite, and calcite: Application to geothermometry. Geochim. Cosmochim. Acta 53, 2985–2995. Clayton R. N. and Mayeda T. K. (1963) The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta 27, 43–52. Clayton R. N., O’Neil J. R., and Mayeda T. K. (1972) Oxygen isotope exchange between quartz and water. J. Geophys. Res. 77, 3057– 3067. Clayton R. N., Goldsmith J. R., and Mayeda T. K. (1989) Oxygen isotope fractionation in quartz, albite, anorthite and calcite. Geochim. Cosmochim. Acta 53, 725–733. Clayton R. N. and Kieffer S. W. (1991) Oxygen isotopic thermometer calibrations. In Stable Isotope Geochemistry: A Tribute to Samuel Epstein (eds. H. P. Taylor Jr., J. R. O’Neil, and I. R. Kaplan), pp. 3–10. Special Publication No. 3, The Geochemical Society. Hu G. and Clayton R. N. (2003) Oxygen isotope salt effects at high pressure and high temperature and the calibration of oxygen isotope geothermometers. Geochim. Cosmochim. Acta 67, 3227–3246. Kelly W. C. and Rye R. O. (1979) Geologic, fluid inclusion and stable isotope studies of the tin-tungsten deposits of Panasqueira, Portugal. Econ. Geol. 74, 1721–1822. Matsuhisa Y., Goldsmith J. R., and Clayton R. N. (1978) Mechanisms of hydrothermal crystallization of quartz at 250° and 15 Kbar. Geochim. Cosmochim. Acta 42, 173–182. McCrea J. M. (1950) On the isotopic chemistry of carbonates and a paleotemperature scale. J. Chem. Phys. 18, 849 – 857. Moore F. and Moore D. J. (1979) Fluid inclusion study of mineralization at St. Michael’s Mount, Cornwall, England. Inst. Mining Metallurgy Trans. B 88, 57– 60. Northrop D. A. and Clayton R. N. (1966) Oxygen isotope fractionation in systems containing dolomite. J. Geol. 74, 174 –196. O’Neil J. R., Mayeda T. K., and Clayton R. N. (1969) Oxygen isotope fractionation in divalent metal carbonates. J. Chem. Phys. 51, 5547–5558. Patterson D. J., Ohmoto H., and Solomon M. (1981) Geologic setting and genesis of cassiterite-sulfide mineralization at Renison Bell, Western Tasmania. Econ. Geol. 76, 393– 408. Polyakov V. B. and Mineev S. D. (2000) The use of Mössbauer spectroscopy in stable isotope geochemistry. Geochim. Cosmochim. Acta 64, 849 – 865. Sun S. S. and Eadington P. J. (1987) Oxygen isotope evidence for the mixing of magmatic and meteoric waters during tin mineralization in Mole granite, New South Wales. Australia. Econ. Geol. 82, 43–52. Strauch G., Thomas R., and Schidlowski M. (1994) Oxygen isotope fractionation in quartz-cassiterite mineralizations from the Erzgebirge region, Germany. In Metallogeny of Collisional Orogens (eds. R. Seltmann, H. Kämpf, and P. Möller), pp. 218 –223. Czech Geological Survey, Prague. Sushchevskaya T. M., Ustinov V. I., Nekrasov I. Y., Gavrilov Y. Y., and Grinenko V. A. (1985) The oxygen-isotope fractionation factor in cassiterite synthesis (in Russian). Geokhimiya 10, 1513–1516. Zhang L., Liu J., Chen Z., and Zhou H. (1994) Experimental investigations of oxygen isotope fractionation in cassiterite and wolframite. Econ. Geol. 89, 150 –157. Zheng Y.-F. (1991) Calculation of oxygen isotope fractionation in metal oxides. Geochim. Cosmochim. Acta 55, 2299 –2307. Zheng Y.-F., Satir M., Metz P., and Sharp Z. D. (1999) Oxygen isotope exchange process and disequilibrium between calcite and forsterite in an experimental C-O-H fluid. Geochim. J. Cosmochim. Acta 63, 1781–1786.