Oxygen Isotope Fractionation in Carbonate and Sulfate Minerals

Geochemical Journal, Vol. 33, pp. 109 to 126, 1999

Oxygen isotope fractionation in carbonate and sulfate minerals

YONG-FEI ZHENG

Department of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, P.R. China

(Received April 20, 1998; Accepted November 23, 1998)

Oxygen isotope fractionations involving carbonates and sulfates have been controversial for a long time. There are important unresolved conflicts among the results of theoretical calculations, experimental measurements and empirical estimates. In this paper, the increment method is adapted to systematically evaluate oxygen isotope fractionations in the carbonates and sulfates. The following sequence of 180-enrichment in carbonate minerals is obtained: siderite > ankerite > magnesite > _ dolomite > calcite > aragonite > strontianite > cerussite >_ witherite. The sequence of 180-enrichment in sulfate minerals is predicted as follows: anhydrite > celestite > barite > anglesite. The internally consistent fractionation factors for the systems carbonate-water and sulfate-water are acquired for a temperature range of 0 to 1200°C, which are in fair agreement with existing experimental and/or empirical data. The present calculations suggest that dolomite should behave isotopically like calcite; equilibrium fractionation between dolomite and calcite is only 0.56%o at 25°C. Aragonite is predicted to be significantly depleted in 180 relative to calcite; equilibrium fractionation between calcite and aragonite is 4.47%o at 25°C. It is possible that polymorphic transition from aragonite to calcite could proceed through an essen tially intact oxygen structure without isotopic resetting. As a result, the temperature dependence of oxygen isotope partitioning in aragonite could be conveyed to calcite. Oxygen isotope inheritance in calcite formation by the polymorphic transition may be of critical importance in attempts to resolve dilemma involving fractionations in aragonite-calcite-dolomite-water systems.

tionation factors between phases at equilibrium. INTRODUCTION The first calculation of isotopic fractionation for Oxygen isotope data for carbonate assemblages oxygen in calcite was done by McCrea (1950). have been extensively used to understand the mode Bottinga (1968), O'Neil et al. (1969), Shiro and of their origin and the temperature of precipitation Sakai (1972), Kieffer (1982), Chacko et al. (1991), from seawater. It is essential that the interpretation Dove et al. (1992) and Gillet et al. (1996) carried of natural variations in oxygen isotope ratios de out calculations of reduced partition function ratios pends on knowledge of the magnitude and tem for calcite. O'Neil et al. (1969) also calculated perature dependence of isotopic fractionation fac oxygen isotope fractionations in the other divalent tors between minerals and fluids. Urey (1947) metal carbonates such as strontianite and witherite. noted that oxygen isotope fractionations between So did Tarutani et al. (1969) for aragonite, and calcium carbonate and water are temperature de Becker and Clayton (1976) for siderite. Golyshev pendent and thus can be used to determine past et al. (1981) developed a lattice dynamical model oceanic temperatures. Since then numerous ther for calculation of equilibrium isotope fractionations mometric calibrations involving carbonates have in carbonate minerals and concluded that the iso been made on the basis of experimental, theoreti topic properties of carbonates are determined pri cal and empirical studies. marily by the cationic radius, the mass effect be Primarily, statistico-mechanical methods were comes substantial only for ions with large masses developed by Urey (1947) and Bigeleisen and like those of Ba2+ and Pb2+. Mayer (1947) for calculations of isotopic frac As pointed out by Hulston (1978) and O'Neil

109 110 Y.-F. Zheng

(1986), the statistico-mechanical methods are temperature dependence of nonequilibrium frac strictly valid only for ideal gases; further ap tionations, and (6) recognition of diagenetic pro proximations and assumptions are required for cesses that alter the original isotopic compositions. calculations involving solids and liquids. For di A quantum jump in the theoretical calculation valent carbonates, Tarutani et al. (1969) illus of oxygen isotope fractionation factors for solid trated that the results of such calculations depend minerals has been brought about by the increment rather strongly on which of the several published method of Schiitze (1980) and its modificators sets of vibrational frequencies are used. In other (Zheng, 1991, 1993a, 1993b; Hoffbauer et al., words, the choice of different vibrational fre 1994). This method was originally based on the quencies for calcite and aragonite could produce observations of Taylor and Epstein (1962) and a large change in the calculated fractionation fac Garlick (1966) that the degree of 180-enrichment tors. O'Neil et al. (1969) noted that, any change in a set of cogenetic silicate minerals can be cor in the calculated line which would bring it into related with bond strengths in the minerals. It fo agreement with the low-temperature experiments cuses on individual bonds, the relative strengths would create serious disagreement with the high of which are given by integral inoic charges and temperature experiments. The calculation of bond-lengths determined from ionic radii. Cation Kieffer (1982) for calcite and some silicates took mass is included in a form appropriate for diatomic into account appropriate modifications on the molecules. Richter and Hoernes (1988) brought statistico-mechanical method. For sulfate minerals, the approach to public attention by applying it to no theoretical fractionation factor has been pro silicate minerals. posed in the literature. Zheng (1991, 1993a) has modified the incre Experimental calibrations of oxygen isotope ment method principally in the following three fractionation factors have been carried out for most aspects by: (1) adding a factor of low temperature carbonates (e.g., McCrea, 1950; Epstein et al., correction with its temperature variable to calcu 1953, 1964; Northrop and Clayton, 1966; O'Neil lation of mineral /3-factors, causing mineral frac et al., 1969; Tarutani et al., 1969; Fritz and Smith, tionations to go from straight lines to more realistic 1970; Matthews and Katz, 1977; Carothers et al., curves in standard fractionation plots; (2) intro 1988; Clayton et al., 1989; Chacko et al., 1991; ducing coupling coefficients to calculation of Scheele and Hoefs, 1992; Rosenbaum, 1994; Kim normalized 180-increments for different cation and O'Neil, 1997) and sulfates (e.g., Lloyd, 1968; oxygen bonds, allowing a detailed consideration Kusakabe and Robinson, 1977; Chiba et al., 1981). of the effect of crystal structure on oxygen isotope Empirical estimates have been presented for partitioning in minerals; and (3) adding a factor CaCO3 polymorphs (e.g., Sommer and Rye, 1978; of mineral/water interaction with mineral I-18O Grossman and Ku, 1986; Patterson et al., 1993; variable to calculation of mineral-water a-factors, Sharp and Kirschner, 1994; Thorrold et al., 1997). giving a further correction to curves involving Unfortunately, most of these calibrations disagree water. As a result, the I-180 indices of individual significantly with each other, particularly for sys minerals are rigorously governed by the funda tems containing calcite, aragonite and dolomite. mental crystal structure and chemical composition, A number of complications of the method have and thus they are a quantitative derivation of the been recognized; the most important one is Kim physico-chemical properties of the minerals. This and O'Neil (1997): (1) validity of the original is especially useful in providing a basis for un calibrations, (2) influence of polymorphic form and derstanding the behaviors of oxygen isotope frac chemical composition of CaCO3 shell or cement tionation in the minerals which have the same on equilibrium fractionation factors, (3) kinetic structure types. Consequently, the modified in effects of both biological (vital) and physical ori crement method has been rationalized to be an gin, (4) reliability of acid fractionation factors, (5) independent approach for calculating oxygen iso 0 in carbonate and sulfate minerals 111

topic fractionation in solid minerals as a function lationship between vibrational frequency and re of statistico-mechanical and crystal structural ef duced mass. fects. For carbonate and sulfate minerals, the major Using the modified increment method, Zheng part of thermodynamic isotope factors arises from (1991, 1992, 1993a, 1993b, 1996, 1997, 1998) has the internal vibrational modes of bonds in complex calculated oxygen isotope fractionation in metal anions. The remaining contributions to the ther oxide and hydroxides, wolframate, silicate and modynamic isotope factors come from the external phosphate minerals. The validity of the theoretical modes of metal-oxygen bonds. The 180-enrich calibrations at both high and low temperatures has ment depends not only upon the bond-type in the been verified by existing data derived from: (1) crystal structure of the minerals, but also upon isotope exchange experiments under hydrothermal the mass of metal cations bonded to the complex or anhydrous (i.e., using carbonate as exchange anions and upon the interatomic distances of all medium) conditions (Zheng, 1991, 1992, 1993a, the cation-oxygen bonds. 1993b, 1997; Rosenbaum and Mattey, 1995), (2) Different crystal structures of carbonate and synthesis experiments (e.g., goethite-water: Yapp, sulfate minerals have been dealt with in this study. 1990; rutile-water: Bird et al., 1993; brucite-wa For carbonates, the divalent metal cation M2+ in ter: Xu and Zheng, 1998), (3) theoretical calcula the calcite group (including calcite, otavite, tions by statistico-mechanical methods (e.g., rhodochrosite, siderite, smithsonite, magnesite, magnetite and garnet: Becker and Clayton, 1976; dolomite and ankerite) occupies the octahedral Rosenbaum and Mattey, 1995; Zheng, 1995), (4) sites with 6-fold coordination, whereas M2+ in the bond-type calculations (e.g., talc and serpentine: aragonite group (including aragonite, strontianite, Saccocia et al., 1998), and (5) empirical estimates witherite and cerussite) occupies the triangular on the basis of natural data (e.g., apatites: Zheng, sites with 9-fold coordination. The calcite structure 1996). These successes are encouraging for the is stable for calcium and the smaller metal ions, continued application of the increment method to whereas the aragonite structure is stable for cal carbonate and sulfate minerals. cium and larger ions (Berry et al., 1983). For sulfates, M2+ in anhydrite is in 8-fold coordination whereas M2+ in the barite group (including barite, CALCULATION METHOD AND RESULTS celestite and anglesite) are in 12-fold coordination. The present calculations follow the modified The structural parameters of the minerals are after increment method which has been described Jaffe (1988) and Smyth and Bish (1988). As listed comprehensively by Zheng (1991, 1993a). In in Tables 1 and 2, the coordination numbers of principle, the degree of the 180-enrichment of a cation-oxygen and interatomic distances are dif mineral is quantified by the oxygen isotope index ferent within each group of either carbonate or of the mineral relative to a reference mineral (I sulfate minerals. This enables quantitative differ 180) . The I-L80 index is calculated by summing entiation of oxygen isotope partitioning among the normalized 180-increment for different cation them. oxygen bonds in the mineral structure (i,t_o'). The In the previous calculations for metal oxide, 180-increment is determined by the effects wolframate and silicate minerals, quartz was taken of cation-oxygen bond strength (C,t_0) and cation as the reference mineral (Zheng, 1991, 1992, mass on isotopic substitution (Wct_0). The cation 1993a, 1993b). In the present calculation, calcite oxygen bond strength is defined as a function of is taken as the reference mineral. The same has cation oxidation state (V), coordination number been done for phosphate minerals (Zheng, 1996). (CN,t) and corresponding ionic radii (rct + r0). This presumes a close similarity in crystal structure Substantially, the I-180 index of a mineral results and oxygen isotope partitioning among carbonate, from a synthesis of crystal chemistry with the re sulfate and phosphate minerals. This implies that 112 Y.-F. Zheng

Table 1. 180-increment for cation-oxygen bonds in carbonates

Mineral Bond ret + ro (A) met Wct-o Cct-o i ct-o i ct-o

Calcite Group Calcite C-O 1.281 12.01 1.02477 1.04085 0.02547 1.0000 Cav1CIUO3lu Ca-O 2.360 40.08 1.04234 0.14124 0.00586 1.0000 Magnesite C-O 1.285 12.01 1.02477 1.03761 0.02539 0.9992 Mg VIClll03111 Mg-O 2.102 24.31 1.03537 0.15858 0.00551 0.9697 Siderite C-O 1.287 12.01 1.02477 1.03600 0.02535 0.9988 Fe"IC11I03m Fe-O 2.144 55.85 1.04631 0.15547 0.00704 1.0961 Rhodochrosite C-O 1.287 12.01 1.02477 1.03600 0.02535 0.9988 Mn VICII103111 Mn-O 2.190 54.94 1.04613 0.15211 0.00686 1.0819 Smithsonite C-O 1.286 12.01 1.02477 1.03681 0.02537 0.9990 Znv1Cu1O3n1 Zn-O 2.111 65.37 1.04798 0.15790 0.00740 1.1237 Otavite C-O 1.290 12.01 1.02477 1.03359 0.02529 0.9982 CdvICm03lu Cd-O 2.288 112.40 1.05249 0.14569 0.00745 1.1278 Dolomite C-O 1.284 12.01 1.02477 1.03842 0.02541 0.9994 CaVIMgVI(CIl[rtm)2 Ca-O 2.381 40.08 1.04234 0.14000 0.00581 0.9957 Mg-O 2.084 24.31 1.03537 0.15995 0.00556 0.9741 Ankerite C-O 1.284 12.01 1.02477 1.03842 0.02541 0.9994 CavIFevI(Cm03u1)2 Ca-O 2.371 40.08 1.04234 0.14059 0.00583 0.9974 Fe-O 2.126 55.85 1.04631 0.15679 0.00710 1.1007

Aragonite Group Aragonite C-O 1.282 12.01 1.02477 1.04004 0.02545 0.9998 Ca[xCulOO1v Ca-O 2.528 40.08 1.04234 0.08790 0.00365 0.7892 Strontianite C-O 1.285 12.01 1.02477 1.03761 0.02539 0.9992 SrIXCIEIO3tV Sr-O 2.636 87.62 1.05057 0.08430 0.00416 0.8424 Witherite C-O 1.287 12.01 1.02477 1.03600 0.02535 0.9988 BaIxCIIlO3[V Ba-O 2.807 137.34 1.05381 0.07917 0.00415 0.8415 Cerussite C-O 1.270 12.01 1.02477 1.04987 0.02569 1.0043 Pb lx Cl11031v Pb-O 2.695 207.19 1.05618 0.08246 0.00451 0.8770

Table 2. 180-increment for cation-oxygen bonds in sulfates

, I reL+ ro (A) Met Wet-o Cct-o i ct-o Mineral Bond ct-o

Anhydrite S-O 1.472 32.06 1.03926 1.01912 0.03924 1.0747 Ca VIIISIV04II1 Ca-O 2.471 40.08 1.04234 0.10117 0.00420 0.8466 Celestite S-O 1.474 32.06 1.03926 1.01764 0.03919 1.0745 SrxIISIV041V Sr-O 2.745 87.62 1.05057 0.06072 0.00299 0.7149 Barite S-O 1.478 32.06 1.03926 1.01488 0.03908 1.0740 BaxlhSIV041V Ba-O 2.879 137.34 1.05381 0.05789 0.00303 0.7196 An lesite S-O 1.482 32.06 1.03926 1.01215 0.03898 1.0735 Pbx"S[VO41V Pb-O 2.783 207.19 1.05618 0.05989 0.00327 0.7474

no additional complexity arises when considering bonds in both carbonate and sulfate minerals. All the anionic complexes of carbonate, sulfate and of the calculation procedures closely follow those phosphate minerals. The C-O bond in calcite is summarized by Zheng (1996), who has also illus thus taken as the reference bond not only for C-O trated the suitability of using the C-O bond in bonds in the other carbonates but also for S-O calcite as the reference bond for P-O bonds in bonds in sulfates. The Ca-O bond in calcite is phosphates. taken as the reference bond for the other M-O I-180 indices of the minerals in question are 0 in carbonate and sulfate minerals 113

Table 3. Calculated oxygen isotope fractionations in carbonate minerals (1031na = A x 106/72 +Bx103/T+C)

Mineral (M 16/M18) 312 1_180 quartz-mineral mineral-water C02-mineral

A B C A B C A B C

Calcite Group Calcite 0.91638 1.0000 0.47 -0.10 0 4.01 -4.66 1.71 -1 .71 10.01 -3.45 Magnesite 0.90203 1.0051 0.41 -0.13 0.01 4.07 -4.64 1.72 -1.77 9.98 -3.46 Siderite 0.92706 1.0194 0.25 -0.19 0.05 4.23 -4 .58 1.73 -1.93 9.92 -3 .47 Rhodochrosite 0.92652 1.0153 0.29 -0.17 0.04 4.19 -4.59 1.72 -1.89 9.94 -3 .46 Smithsonite 0.93228 1.0228 0.21 -0.20 0.06 4.27 -4.56 1.73 -1 .97 9.90 -3 .47 Otavite 0.94998 1.0046 0.41 -0.12 0.01 4.07 -4.64 1.71 -1 .77 9.97 -3 .45 Dolomite 0.90977 1.0040 0.42 -0.12 0.01 4.06 -4.65 1.71 -1 .76 9.99 -3.45 Ankerite 0.92208 1.0097 0.36 -0.15 0.03 4.12 -4 .62 1.71 -1 .82 9.96 -3.45

Aragonite Group Aragonite 0.91638 0.9296 0.57 1.03 -0.47 3.91 -5.79 1.92 -1 .61 11.14 -3 .66 Strontianite 0.94199 0.9212 0.58 1.16 -0 .52 3.90 -5 .93 1.95 -1 .60 11.27 -3.69 Witherite 0.95607 0.9071 0.61 1.38 -0.61 3.87 -6 .15 1.99 -1 .57 11.49 -3.73 Cerussite 0.96724 0.9099 0.60 1.34 -0.59 3.88 -6.10 1.98 -1.58 11.45 -3 .72

Table 4. Calculated oxygen isotope fractionations in sulfate minerals (1031na = A X 106/72 + B x 1031T + C)

Mineral (M 16/M 18)312 1_180 quartz-mineral mineral-water C02-mineral

A B C A B C A B C

Anhydrite 0.91791 1.0160 0.29 -0.17 0.05 -0.18 -0.07 0.05 4.19 -4 .59 1.71 Cerussite 0.93805 0.9619 0.52 0.52 -0.25 0.05 0.62 -0.25 3.96 -5 .28 1.82 Barite 0.95070 0.9498 0.54 0.71 -0.33 0.07 0.81 -0.33 3.94 -5 .47 1.86 Anglesite 0.96169 0.9456 0.54 0.78 -0.36 0.08 0.88 -0.36 3.94 -5 .54 1.87

presented in Table 3 for carbonates and in Table of Becker and Clayton (1976) on the magnetite 4 for sulfates. The parameters used are given in water system. Particularly, the theoretical calibra Tables 1 and 2 where interatomic distances are tion of the apatite-water systems can be applied after Smyth and Bish (1988). The weight q was to paleotemperature determinations of ancient employed to deal with different types of cation oceans (Zheng, 1996). These imply that an exten oxygen bonds in silicate minerals (Zheng, 1993a). sion of the methodology to the other low-tem For the present study, it is treated by assuming perature minerals is potentially valid. the coupling coefficient k to be -1 for the C-O and As done by Zheng (1991, 1992, 1993a, 1993b, S-O bonds as well as for the M-0 bonds. Fol 1996) for the metal oxide, wolframate, silicate and lowing the previous assumption by Zheng (1991) phosphate minerals, constraints on the assumptions for low-temperature correction to metal oxides, the have been set by the nature of cation-oxygen bonds same AE value of 1 kJ/mol (which is just to in the carbonate and sulfate minerals and by transform an unitless variable into an energy pa known data from experimental and/or empirical rameter in the term D) is also taken for carbonate calibrations. As pointed out by Zheng (1993a, and sulafte minerals. The correctness of the low 1993b), the advantage of the assumptions is that temperature correction has been verified by the it can be uniformly applied within individual synthesis experiments of Bird et al. (1993) on the mineral groups (having the same structural types rutile-water system and the theoretical calculations and compositional series). The two sets of ther 114 Y.-F. Zheng

modynamic oxygen isotope factors thus obtained priate to the present study on low-temperature for mineral-mineral and mineral-water systems are minerals. The reduced partition function ratios of internally consistent. The disadvantage of the as water are after Hattori and Halas (1982) which sumptions is that it could introduce systematical were based on the calculation of Richet et al. errors in the thermodynamic isotope factors be (1977). The experimental investigation of Horita tween the different mineral groups. The reasonably and Wesolowski (1994) on oxygen isotope frac good agreement of the calculated fractionation tionation between liquid and vapor waters cor factors for the mineral-water systems with those roborates the validity of the assumption by Hattori derived from the experimental and/or empirical and Halas (1982) for this system. The application calibrations, as shown in next section, demon of the reference data results in that both 1031n/3 strates that the assumptions are adequate. Also the and 1031na values calculated are not straight lines C-O bond in calcite is reasonably used as the at low temperatures below about 400°C when reference bond for the S-O bonds in sulfates. plotted against 1/72. Therefore, it is necessary to Hence, the I-180 indices calculated in this study include both 1/TZand 1/T terms in the same frac are quantitatively representative of 180-enrich tionation equation for accurately fitting data at both ment in the carbonate and sulfate minerals. high and low temperatures. According to the principles of the increment The calculated fractionation factors are method (Schiitze, 1980; Zheng, 1991), the I-i80 graphically presented in Figs. I to 9, along with values greater than unity in Tables 3 and 4 imply the existing experimental calibrations and empiri an 180-enrichment in a certain mineral relative to cal estimates. The algebraic expressions of the calcite, whereas the I-180 values less than unity fractionation factors are given in the form of mean a depletion of 180 in the mineral in question 1031na = A x 106/T2 + B x 103/T + C. The values with respect to calcite. The greater the I-18O in of the parameters A, B and C are listed in Table 3 dex of a mineral, the more 180-enriched it is. In for the carbonates and in Table 4 for the sulfates, terms of the size of I-18O indices obtained, there respectively, for the temperature range from 0 to fore, the order of 180-enrichment in the carbonate 1200°C. The fractionation factors between the minerals can be predicted as follows: smithsonite carbonate or sulfate mineral pairs can be deduced > siderite > rhodochrosite > ankerite > magnesite by combining the parameters for the 1031na _> otavite >_ dolomite > calcite > aragonite > equations of quartz-mineral fractionations in strontianite > cerussite >_witherite. The sequence Tables 3 and 4. of 180-enrichment in the sulfate minerals is: an Zheng (1991, 1993a, 1993b) has analysed hydrite > celestite > barite > anglesite. In com various error sources associated with the modified parison with other classes of minerals, carbonates increment method. By using a reference mineral and sulfates are 180-rich because oxygen is to determine the major portion of 1031n/3values for bonded to the small, highly-charged C4+ or S6+ the other minerals, the methodology itself mini The nature of divalent cations also plays a con mizes any inherent errors that are present . Spe siderable role in determining the overall isotopic cifically, the I-180 indices of the minerals within partitioning in the carbonates and sulfates. the carbonate group are internally consistent be As done in Zheng (1991, 1992, 1993a, 1993b , cause they are rigorously governed by the funda 1996), this study applies the reduced partition mental crystal structure. Uncertainties in the cal function ratios of calcite and quartz listed by culated 1031n/3values between the carbonate min Clayton et al. (1989) on the basis of the calcula erals can thus achieve minimum. For the sulfate tion of Kieffer (1982). The reduced partition minerals, it is appropriate to assume the close function ratios of calcite and quartz listed by similarity in crystal chemistry and oxygen isotope Clayton and Kieffer (1991) are valid only at tem behavior to carbonates, as successfully demon peratures greater than 400 K, which are inappro strated for phosphates (Zheng, 1996). In this re 0 in carbonate and sulfate minerals 115

gard, no additional error is implied by taking the 1969; Tarutani et al., 1969; Shiro and Sakai, 1972; C-O bond in calcite as the reference bond for the Becker and Clayton, 1976; Chacko et al., 1991). In S-O bonds in the sulfates. Consequently, the un order to test the accuracy of the present calcula certainties contributed to the thermodynamic iso tions, comparison can be made between the cal tope factors (1031n/3)for the carbonate and sulfate culated fractionations and existing experimental minerals can be reasonably estimated to be within and/or empirical calibrations. ±5% of the factor values (i.e., 20 ± 1.0%0 or 4 ± 0.2%0). Carbonates As shown in Table 3 and Figs. 1 and 2, the present calculations suggest a considerable differ DISCUSSION ence in oxygen isotope fractionations between the The theoretical calculations by means of the calcite group and aragonite group minerals, with modified increment method present a set of self the calcite group being enriched in 180 relative to consistent fractionation factors for quartz-carbon the aragonite group at isotopic equilibrium. The ate-sulfate-water systems. Although the increment same sequence of 180-enrichment was obtained method employs the reduced partition function from the hydrothermal experiments of O'Neil et al. ratios of quartz and calcite tabulated in Clayton et (1969) for calcite, strontianite and witherite. al. (1989) as the reference minerals, which were Tarutani et al. (1969) investigated the effect of primarily calculated by Kieffer (1982) by classical magnesium substitution on oxygen isotope frac statisco-mechanical methods with several addi tionation between CaCO3 and H20 at 25°C and tional assumptions and approximations, the present found that 180 concentrates in magnesium calcite. results are independent of the classical calculations This is concordant with the theoretical prediction (e.g., McCrea, 1950; Bottinga, 1968; O'Neil et al., that magnesite is enriched in 180 relative to cal cite.

§88§ 8 S 8 6 8 ' O 8 L) 88g 0 a Ilk) N 10 wm Q 5 9

c 0 1 0) 8 ~`oc cam 4 CC X,O-~ 7 c 0 yy. ca 6 V~`J Nc 3 0 U 5 o'o O IV cu ca 0 2 v 4 m

3 GaG~ e~`~e c O 1 2

Dolom+tJ 1 /g~~te 0 0 2 4 6 8 10 12 0 106/-2 0 2 4 6 8 10 12 106/12 Fig. 1. Calculated oxygen isotope fractionation fac tors between dolomite and calcite, between calcite and Fig. 2. Calculated oxygen isotope fractionation fac aragonite, and between dolomite and aragonite, re tors between quartz and the common carbonates indi spectively. cated, respectively.

cO 116 Y.-F. Zheng

40 N888 8 S 9 S CD U C) N °e N b 40 t-This calculation 1-This calculation 35 2-Carothers et al. (1988) 35 2-Patterson et al. (1993) /3 3-O'Neil et al. (1969) i 3-O'Neil et al. (1969) / 30 / '1 Ql 30 / / / 2 d)25 25 / 0 20 CU 20 // AR Lco 0 •4 c U 15

c 15 U/ QS C: 0 10 10 3 a 3 5 1 5 Y/ 0 0 0 100 200 300 400 500 0 2 4 6 8 10 12 Temperature (°C) 106/T2 Fig. 4. Comparison of the calculated oxygen isotope Fig. 3. Comparison of the calculated oxygen isotope fractionations for the siderite-water and witherite-wa fractionations for the calcite-water system (dashed ter systems, respectively, with those derived from the curve) and the aragonite-water system (solid curve), hydrothermal experiments of O'Neil et al. (1969) and respectively, with those derived from the hydrothermal Carothers et al. (1988). experiments of O'Neil et al. (1969) on the CaCO3-H20 system (circles denote the complete equilibrium, tri angles denote the partial equilibrium, and squares de note the low-temperature precipitation) and the em of 3.2 to 30.3°C, which are exactly overlapped by pirical calibration of Patterson et al. (1993) for the the low-temperature synthesis data of O'Neil et al. aragonite-water system (dot-and-dash curve). (1969). As depicted in Fig. 4, there is good agreement between the theoretical and experimental calibra Figure 3 shows a comparison of oxygen iso tions involving siderite and witherite. For the tope fractionations calculated for CaCO3 poly strontianite-water system (Fig. 5), the theoretical morphs (calcite and aragonite) with those for the fractionations are in good agreement with the ex CaCO3-H20 systems derived from the hydrother perimental calibrations of O'Neil et al. (1969) at mal experiments of O'Neil et al. (1969) in the temperatures below -25°C, but they are gradually temperature range from 0 to 500°C (the values lower than the experimental values at higher have been recalculated using the C02-H20 frac temperatures. tionation of 1.0412). The fractionations derived For the dolomite-water system (Fig. 5), the from the exchange experiments of both complete present calculations are consistent with the ex equilibrium at temperatures above 200°C and perimental results of Fritz and Smith (1970) and partial equilibrium at temperatures above 60°C are Matthews and Katz (1977), but they are system in good agreement with the theoretical calibration atically lower than the experimental determination of the calcite-water system, while the fraction of Northrop and Clayton (1966) which was made ations obtained by chemical precipitations at 0 and using the partial exchange method. It is noted that 25°C are close to the theoretical calibration of the the experimental data of Northrop and Clayton aragonite-water. Patterson et al. (1993) empiri (1966) were rather scattered relative to the fitted cally calibrated oxygen isotope fractionations be straight line. Furthermore, O'Neil (1986) pointed tween aragonite and water for a temperature range out that the dolomite-water fractionations deter O in carbonate and sulfate minerals 117

40 12 1-This calculation 1-Rosenbaum (1994) 2-Zheng (1993) 2-Matthews& Katz (1977) 35 A 3-Scheele & Hoefs (1992) 3-Fritz& Smith (1970) 'vaa) 10 4-Chacko et al. (1991) 4-O'Neilet al. (1969) . 30 Ca 5-O'Neil& Epstein (1966) U) 5-Northrop& Clayton(1966) U c» M .o a) 25 3 X 8 t0 0 C 0 0 20 C L as 0 05 U ~' o 0 2 15 I 6 C 3`• O 10 \ \4 2 \ C 0 4 4\ \ 5

2 0 0 200 400 600 800 10001200 0 100 200 300 400 500 Temperature (°C) Temperature (°C) Fig. 6. Comparison of oxygen isotope fractionations Fig. 5. Comparison of the calculated oxygen isotope between CO2 and calcite derived from the theoretical fractionations for the dolomite-water and strontianite data of Richet et al. (1977) and Kieffer (1982) com water systems, respectively, with those derived from bined by Zheng (1993a) with the experimental calibra hydrothermal experiments. Solid curves denote the do tions of O'Neil and Epstein (1966), Chacko et al. lomite-water system, and dashed curves denote the strontianite-water system. (1991), Scheele and Hoefs (1992) and Rosenbaum (1994).

mined by the partial exchange technique are the experimental calibrations of O'Neil and Epstein probably greater than the equilibrium fraction (1966), Chacko et al. (1991), Scheele and Hoefs ations. As observed by O'Neil et al. (1969) for the (1992) and Rosenbaum (1994). calcite-water system and by Matthews and A ragonite-calcite water system The behavior Beckinsale (1979) for the quartz-water system, of oxygen isotope fractionations in the system there are also systematic discrepancies between the aragonite-calcite-water has been controversial in results of partially equilibrated and completely last three decades. The precipitation experiments equilibrated systems. The present calculations of Tarutani et al. (1969) and the natural observa predict that equilibrium fractionations between tions of Sommer and Rye (1978) as well as dolomite and water are from 39.09 to 31.78%o at Grossman and Ku (1986) suggest that calcite is the temperatures from 0 to 25°C, being in good depleted in 180 relative to aragonite. In contrast, agreement with the observation of Clayton et al. the natural observations of Epstein et al. (1953), (1968) for a natural fractionation of 35.05%o be Behrens and Land (1972) as well as Horibe and tween sedimentary dolomite and its interstitial Oba (1972) show that calcite is enriched in 180 water. relative to aragonite. Also listed in Table 3 are the fractionation For the CaCO3 polymorphs, it appears in terms equations for the C02-carbonate systems, where of the principles of the increment method (Schiitze, the reduced partition function ratios of CO2 are 1980; Zheng, 1991, 1993a) that the more dense after Richet et al. (1977) and those of calcite after aragonite should be depleted in 180 relative to the Kieffer (1982). As shown in Fig. 6, the theoretical less dense calcite. The present calculations show data combined by Zheng (1993a) for the C02 a significant 180-depletion in aragonite relative to calcite fractionations are in good agreement with calcite, with the equilibrium fractionation between

rT 118 Y.-F. Zheng calcite and aragonite being 4.47%o at 25°C (Fig. citic shells were formed by polymorphic transi 1). The similar sequence of 180-enrichment has tion from the aragonitic shells with little (or no) been suggested for Si02 and A12SiO5polymorphs postdepositional alteration of the isotopic (Zheng, 1993c). According to the electrostatic compositon, or that the calcitic and aragonitic potentials of oxygen sites in the CaCO3 polymor shells were originally precipitated at different phs Smyth (1989) predicted that calcite is enriched temperatures. However, caution must be exercised in 180 relative to aragonite at isotopic equilibrium. when comparing the results of laboratory systems This prediction is reasonably valid because cat to those of natural systems, particularly when ionic mass is the same for both calcite and arago biological reactions are involved. nite. In the hydrothermal experiments of O'Neil As shown in Fig. 3, the experimental determi et al. (1969) at 100°C on the oxygen isotope nations of O'Neil et al. (1969) on the calcite-wa fractionation between CaCO3 and H20, the ex ter system at low temperatures are consistent with trapolated equilibrium fractionation is 18.66% the empirical calibration of Patterson et al. (1993) when using a natural calcite as starting material, on the aragonite-water system and also close to which is considerably greater than a value of the theoretical calculation in this study on the 16.98%o obtained by using a reagent aragonite as aragonite-water system. O'Neil et al. (1969) X starting material. The other carbonates in the ara rayed their run products and no aragonite was gonite group are also depleted in 180 relative to found. It remains questionable why systematic calcite, being consistent with the hydrothermal discrepancies are found in the results between the experiments of O'Neil et al. (1969). experimental, empirical and theoretical calibra Kim and O'Neil (1997) observed that carbon tions. If aragonite does behave unlike calcite iso ates could be precipitated reproducibly in or out topically, a possible explanation is that the low of oxygen isotope equilibrium with the environ temperature runs of O'Neil et al. (1969) would mental solution by varying the concentrations of originally precipitate as aragonite, which subse bicarbonate and cation. They found that the frac quently converted to calcite without changing its tionation factors between synthesis carbonates and 8180. In this regard, metastable aragonite would water measured for solutions with different initial presumably precipitate at first from aqueous solu concentrations differ by as much as 2 to 3 permil tions in the infancy of CaCO3 crystallization and at a given temperature. According to achieve isotopic equilibrium with ambient water McConnaughey (1989), oxygen isotope at low temperatures. Polymorphic transition from disequilibrium between synthetic carbonates and aragonite to calcite would proceed quickly through water occurs when CO2 undergoes reactions to an essentially intact oxygen structure, without re form HCO3-. setting isotopic re-equilibration with the water. In According to the discussion above, the present other words, it only involves the breaking and calculations reasonably predict the magnitude and rebinding of the bond between Ca2+ and carbon direction of the equilibrium oxygen isotope frac ate complex [C03]2 rather than the bond between tionation between calcite and aragonite. In this carbon and oxygen within the carbonate complex. context, the coexisting calcite-aragonite pairs As a result, the calcite could have preserved the showing negative fractionations could be simply oxygen isotope composition inherited from the assemblages in isotopic nonequilibrium, although preexisting aragonite. it is not clear why such nonequilibrium fraction Experimental studies on equilibrium fraction ations have resulted. ation in the CaCO3-H2O system indicate that dis Some marine organisms with aragonitic shells solution-reprecipitation controls the rate of isotope often show very small fractionations (even no exchange between finely dispersed calcium car fractionation at all) when compared to coexisting bonate and aqueous solutions (Clayton, 1959; calcitic shells. This may imply either that the cal Northrop and Clayton, 1966; O'Neil et al., 1969). 0 in carbonate and sulfate minerals 119

On the basis of oxygen isotope exchange kinetics 10 and detailed microscopic examination, Anderson 1-This calculation and Chai (1974) concluded that dissolution and 2-Sharp & Kirschner (1994) 3-Clayton et al. (1989) reprecipitation is the mechanism by which isotope N 8 exchange occurs between calcite and water under c 0 hydrothermal conditions. They also found that the c`n 6 extent of fractional isotope exchange correlates U with the observed extent of recrystallization. cc Clayton (1959) observed that the calcium carbon C) 4 ate in the high temperature experiments underwent

recrystallization during polymorphic transition c c) 0 from aragonite to calcite and hence achieved new '' 2 calcite-water isotope equilibrium under the hy drothermal conditions. This implies the breaking and rebinding of the bonds not only between Ca2+ 0 0 200 400 600 800 10001200 and carbonate complex [C03]2 but also between Temperature (°C) carbon and oxygen within the carbonate complex during the polymorphic transition and recrystalli Fig. 7. Comparison of oxygen isotope fractionations between quartz and calcite derived from the empirical zation. However, it is not clear whether the cal calibration of Sharp and Kirschner (1994) with the ex cium carbonate in the low temperature experiments perimental data of Clayton et al. (1989) and the theo of O'Neil et al. (1969) experienced the breaking retical calculations for the quartz-calcite system and rebinding of the bonds not only between Ca2+ (Clayton et al., 1989 after Kieffer, 1982) and the and [C03]2 but also between carbon and oxygen quartz-aragonite system in this study. within the carbonate complex during the poly morphic transition. Further studies on synthetic products by atomic force microscopy and/or other techniques are thus required to clarify this pro in their synthesis experiments that in cases where cess. rutile formed by inversion of less stable anatase, Tarutani et al. (1969) noted that no significant the rutile appears to be at least partially inherited isotope fractionation occurs in conversion from its oxygen isotope composition from the precur aragonite to calcite by heating their samples in sor polymorph. According to a comprehensive vacuum at 410°C. Clayton et al. (1975) observed study of oxygen isotope fractionation in magne from their high-pressure experiments that there is tites, Zheng (1995) suggested that most magne no measurable effect on oxygen isotope fraction tites could have the spinel-type structure in the ation between CaCO3 and H20 at 500°C due to the first stage of their crystallization, although they calcite-aragonite phase change which occurs be are normally observed to occur in the inverse tween 11 and 13 kbar experiments. Negligible spinel-type structure in nature. It was hypothesized isotope fractionation also occurs in other experi that the inverse sp i nel -structural magnetites used ments involving polymorphic transition, e.g., from in isotopic exchange experiments can inherit the graphite to diamond (Kropotova et al., 1967) and temperature dependence of oxygen isotope frac from fl-quartz to a-quartz (Clayton et al., 1972; tionation in the spinel structural magnetite which Matsuhisa et al., 1979), although the theoretical formed by reducing a precursor hematite. calculations predicted considerable isotopic effects An empirical calibration for the quartz-calcite associated with the change in crystal structure geothermometer has been made by Sharp and (Bottinga, 1969; Shiro and Sakai, 1972; Kawabe, Kirschner (1994) on the basis of measured oxygen 1978; Zheng, 1993c). Bird et al. (1993) observed isotope fractionations in greenschist-facies marbles, 120 Y.-F. Zheng

veins composed of cogenetic quartz and calcite, to the quartz-calcite system in the course of dia and various low-grade metamorphic rocks. Figure genesis and metamorphism. However, it is not 7 shows a comparison of their data with the theo clear if such an oxygen isotope inheritance could retical calculations on the quartz-calcite system exist in calcite formation by polymorphic transition (Clayton et al., 1989 after Kieffer, 1982) and the from aragonite during diagenesis and metamor quartz-aragonite system in this study and the ex phism. Further experimental insights are highly perimental calibration of Clayton et al. (1989) on needed to assess the theoretically predicted effect the quartz-calcite system by means of partial ex of isotopic inheritance. change technique. Apparently, the empirical Dolomite-calcite-water system The behavior of quartz-calcite curve of Sharp and Kirschner (1994) oxygen isotope fractionation in the dolomite-cal is discrepant with the experimental data of Clayton cite-water system has been particularly intractable. et al. (1989). However, it is in good agreement Its resolution has a great impact on controversies with the theoretical quartz-aragonite curve, par concerning the origin of dolomites in sedimentary ticularly at low temperatures. It is unclear why rocks. One of the major questions involved is there is such an agreement between the results of whether the dolomite precipitated directly from empirical and theoretical calibrations. seawater or instead replaced sediments composed It hardly supposes that the reduced partition of calcium carbonate (e.g., Ingerson, 1962; Epstein function ratios of calcite and quartz tabulated by et al., 1964; Weber, 1965; Katz and Matthews, Clayton et al. (1989) after the theoretical calcula 1977; Land, 1980; Hardie, 1987). Many geological tion of Kieffer (1982) are problematic at low evidences suggest a direct chemical precipitation temperatures, because the reasonable fractionation of dolomite in marine evaporite basins, yet no one factors between various minerals and water have has been able to demonstrate experimentally such been obtained by means of the modified increment a process, and modern sediments comparable to method. If the experimental calibration of Clayton the massive strata of high-purity dolomite have et al. (1989) and the present calculation are as not been observed. On the other hand, the dolo sumed to be correct, a potential interpretation is mite is considered a diagenetic replacement of that the calcites investigated by Sharp and preexisting calcium carbonate. However, the cause Kirschner (1994) would be mostly the product of and the mechanism of the alteration process re polymorphic transition from aragonite without main to be clarified. The oxygen isotope compo isotopic re-equilibration during diagenesis and sition of cogenetic dolomite and calcite could metamorphism. This implies that the greenschist present a means to study the chemical mechanism facies marbles would be once all aragonite which of dolomite formation. has retained its isotopic signature even undergo Oxygen isotope studies of some metamorphic ing metemorphism at over 400°C, and that the assemblages containing dolomite and calcite quartz-calcite veins which were precipitated at pointed to an '8O-enrichment, by an extrapolation about 300°C could be originally precipitated as to sedimentary conditions, of 5 to 9%o in dolo metastable quartz-aragonite. The similar interpre mite relative to coexisting calcite (e.g., Clayton tation can be applied to oxygen isotope data for and Epstein, 1958; Engel et al., 1958; Clayton et calcite marbles from the Precambrian Wyman al., 1968; Sheppard and Schwarcz, 1970). Because Formation of Lone Mountain in Nevada, where dolomite has been not synthesized at 25°C, it is measured isotopic fractionations between quartz currently impossible to directly measure the equi and calcite exhibit a well behaved array that is librium oxygen isotope fractionation between do parallel an isotherm on a 8-6 diagram (Richards et lomite and water at sedimentary temperatures. al., 1996). In this regard, the temperature depen Simple extrapolation to 25°C of isotopic data from dence of oxygen isotope fractionations between laboratory experiments at high temperatures indi quartz and aragonite could have been conveyed cated an 180-enrichment of 4 to 7%o in dolomite 0 in carbonate and sulfate minerals 121 relative to calcite (e.g., Epstein et al., 1964; O'Neil water fractionation close to that of calcite-water. and Epstein, 1966; Northrop and Clayton, 1966; In other words, dolomite and calcite which were Hazma and Broecker, 1974). slowly coprecipitated in the same environment On the contrary, coexisting dolomite-calcite should have nearly identical oxygen isotope com pairs from sedimentary rocks usually show very positions. small fractionations (e.g., Friedman and Hall, The theoretically predicted fractionation be 1963; Degens and Epstein, 1964; Fritz, 1967; tween dolomite and aragonite is 5.1%o at 25°C Rumble et al., 1991). The present calculations and 2.3%o at 250°C. If there would be the oxygen suggest that dolomite behaves isotopically like isotope inheritance in calcite formation by poly calcite; the equilibrium fractionation between do morphic transition from aragonite, the theoretical lomite and calcite is only 0.56 permit at 25°C (Fig. predictions could be well used to account for the 1). As depicted in Fig. 5, the theoretically pre greater dolomite-calcite fractionations observed in dicted fractionations for the dolomite-water sys some natural assemblages. In this regard, the cal tem agree well with the experimental data of Fritz cite depleted in 180 relative to the coexisting do and Smith (1970) and Matthews and Katz (1977). lomite could be the product of polymorphic Epstein et al. (1964) experimentally found that the transformation from aragonite without isotopic isotopic composition of a dolomite which was resetting, and that both dolomite and aragonite synthesized from a parent calcite is equal to the could be coprecipitated in isotopic equilibrium expected value for calcite under the experimental with seawater. In order to gain an insight into the temperatures of 400 to 600°C. Fritz and Smith origin of dolomites by means of oxygen isotope (1970) precipitated protodolomite chemically in studies, the microstructural and petrographic fea laboratory and found by XRD and microscopic tures of the dolomite in question should be further examination that aragonite was the only crystalline examined by the electron microscope, carbonate co-precipitate of the protodolomite in cathodoluminesence and other techniques. their experiments at 78 and 59°C; the precipitates In nature, dolomite has been found to have the at 40 and 25°C indicate the presence of small 5180 values either greater than or similar to co amounts of magnesium calcite without aragonite. existing calcite. However, it is critical to examine If dolomite really has oxygen isotope composi whether the samples studied are cogenetic phases. tion identical to coexisting calcite, these observa In terms of oxygen isotope compositon alone, the tions can be explained by assuming oxygen iso greater dolomite-calcite fractionations could result tope inheritance in dolomitization by magnesium from one of the following two processes: (1) dia substitution for calcium in calcite, which could genetic or metamorphic alteration of the calcite, proceed through an essentially intact oxygen in and both calcite and dolomite were precipitated the carbonate anion without isotopic re-equilibra in isotopic equilibrium with seawater; or (2) tion. The oxygen isotope inheritance in mineral polymorphic transition from aragonite to calcite formation by cation substitution has been supposed without isotopic re-equilibration with ambient for the replacement of wolframite to scheelite in materials, and both aragonite and dolomite were nature (Zheng, 1992) and apatite to BiPO4 in precipitated in isotopic equilibrium with seawa laboratory (Shemesh et al., 1988; Zheng, 1996). ter. In both cases, the dolomite has remained iso The accuracy of the experimental data of topically unaffected. Northrop and Clayton (1966) from the partial ex If dolomite is found to have the similar 5180 change technique is questionable as discussed be values to coexisting calcite, it has an important fore. Therefore, the present calculations reason bearing on the origin of the dolomite: (1) the do ably predict oxygen isotope fractionations between lomite would be a primary chemical precipitate, dolomite and calcite. In this regard, dolomite can and both dolomite and calcite were precipitated be a primary chemical precipitate with a dolomite in isotopic equilibrium with seawater; or (2) the 122 Y.-F. Zheng

O Ln N 8088$ S 8 N 9 40 8 1-Thiscalculation 2-Chiba et al. (1981) 7 Barite 35 4-Uoyd3-Kusakabe(1968) &Robinson (1977)/ i 30 4/ : N 6 my m may. v i ~ 5 25 :

N Jib

20 i c0 4 C~ 2 C1 c 15 Ile U 3 B 0 T 10 T 2

/ 5 •% 1 %3 i 0 0 0 2 4 6 8 10 12 0 2 4 6 8 10 12 106/T2 106/T2 Fig. 8. Comparison of oxygen isotope fractionations Fig. 9. Calculated oxygen isotope fractionation fac between sulfate and water derived from the present tors between quartz and the sulfate minerals indicated, calculations with experimental calibrations. Solid respectively. curves denote the anhydrite-water system, dashed curves denote the barite-water system.

Sulfates The present calculations suggest an 180-en richment in anhydrite relative to barite at equilib dolomite was formed by replacing preexisting rium (Table 4). This is consistent with the hydro calcite via a solid-state diffusion mechanism thermal experiments of Lloyd (1968), Kusakabe without any change in the oxygen isotope compo and Robinson (1977) and Chiba et al. (1981). As sition, as previously suggested by Epstein et al. depicted in Fig. 8, the experimental calibrations (1964). However, the solid-state diffusion of of the sulfate-water systems are fairly comparable magnesium is too slow at sedimentary tempera with the theoretical ones. Anhydrite is slightly tures to accomplish the replacement (Anderson, depleted in 180 relative to quartz (Fig. 9) but en 1969). Alternatively, solution-recrystallization is riched in 180 relative to calcite (Table 4). Since an effective mechanism for the dolomitization temperature dependence of the equilibrium frac (Epstein et al., 1964; Katz and Matthews, 1977). tionations between quartz and anhydrite and be In this case, an internal fluid of the same 5180 tween anhydrite and calcite is very weak at tem value is appealed, not only from which the pre peratures above 200°C, it is difficult to derive cursor calcite precipitated, but also by which the isotopic temperatures from either of the two min dolomitization took place. If involved would be eral pairs on the basis of the observed fraction an external fluid that is depleted in 180 relative to ations in hydrothermal and metamorphic assem seawater (e.g., intrastratal fluid or formation wa blages. ter), the remaining calcite could become enriched in 180 relative to the newly formed dolomite. Temperature changes during the replacement CONCLUSIONS would result in irregular fractionations between Theoretical calculations suggest that oxygen coexisting dolomite and calcite. isotope fractionation between dolomite and cal

i O in carbonate and sulfate minerals 123 cite should be very similar at equilibrium, whereas 49453003) and the Chinese Academy of Science within aragonite would be significantly depleted in 180 the framework of the project "Stable Isotope Geo chemistry of the Earth's Crust and Mantle". Thanks relative to calcite at sedimentary temperatures. are due to Drs. M. E. Bottcher, R. N. Clayton, G. Coherent sets of isotopic temperature scales com Cortecci, J. Horita, A. Longinelli, F. J. Longstaffe and posed of the systems. carbonate-water and sulfate Z. D. Sharp as well as two anonymous reviewers for water have been made available. This enables their critical but helpful comments on the manuscript. quantitative analysis of sedimentary or metamor phic conditions recorded in the oxygen isotope REFERENCES composition of different carbonate and sulfate minerals. In particular, the present calculations Anderson, T. F. (1969) Self-diffusion of carbon and oxygen in calcite by isotope exchange with carbon provide an insight into the longstanding contro dioxide. J. Geophys. Res. 74, 3918-3932. versies involving the behavior of oxygen isotope Anderson, T. F. and Chai, B. H. (1974) Oxygen iso partitioning in the system aragonite-calcite-dolo tope exchange between calcite and water under hy mite-water. This has an important bearing on the drothermal conditions. Geochemical Transport and origin of dolomite in sedimentary rocks. Kinetics (Hofmann, A. W., Giletti, B. J., Yoder, H. The interpretation to the calculated results and S., Jr. and Yund, R. A., eds.), 219-227, Carnegie Institution of Washington Publication. the existing experimental and/or empirical data Becker, R. H. and Clayton, R. N. (1976) Oxygen iso suggests that dolomite could be a primary chemi tope study of a Precambrian banded iron formation, cal precipitate with the similar fractionation be Hamersley Range, western Australia. Geochim. havior to calcite. Metastable aragonite could pre Cosmochim. Acta 40, 1153-1166. fer to form in the first stage of CaCO3 crystalli Behrens, E. W. and Land, L. S. (1972) Subtidal Holo zation and subsequently convert to calcite with cene dolomite, Baffin Bay, Texas. J. Sed. Petrol. 42, 155-161. out isotopic resetting at low temperatures. It is Berry, L. G., Mason, B. and Dietrich R. V. (1983) possible that the polymorphic transition from ara Mineralogy: Concepts, Descriptions and Determina gonite to calcite in nature and laboratory could tions. WH Freeman and Co, San Francisco, 561 pp. proceed through an essentially intact oxygen Bigeleisen, J. and Mayer, M. G. (1947) Calculation of structure without isotopic resetting. In other words, equilibrium constants for isotopic exchange reactions. J. Chem. Phys. 15, 261-267. it could only involve the breaking and rebinding Bird, M., Longstaffe, F. J. and Fyfe, W. S. (1993) of the bond between Ca2+ and carbonate complex Oxygen-isotope fractionation in titantium-oxide [C03]2 rather than the bond between carbon and minerals at low temperatures. Geochim. Cosmochim. oxygen within the carbonate complex. However, Acta 57, 3083-3091. it remains to be tested whether the temperature Bottinga, Y. (1968) Calculation of fractionation fac dependence of oxygen isotope fractionations in tors for carbon and oxygen isotopic exchange in the system calcite-carbon dioxide-water. J. Phys. Chem. volving aragonite can be conveyed to calcite dur 72, 800-808. ing diagenesis and metamorphism and in synthe Bottinga, Y. (1969) Carbon isotope fractionation be sis experiments. It is possible that the phase change tween graphite, diamond and carbon dioxide. Earth of aragonite to calcite is a critical process for Planet. Sci. Lett. 5, 301-307. causing the difference in the oxygen isotopic Carothers, W. W., Adami, L. H. and Rosenbauer, R. J. compositions between some coexisting dolomite (1988) Experimental oxygen isotope fractionation between siderite-water and phosphoric acid liberated and calcite pairs. Once the controversial issues are CO2-siderite. Geochim. Cosmochim. Acta 52, 2445 clarified, the stable isotope study of coexisting 2450. dolomite and calcite can shed light on the mecha Chacko, T., Mayeda, T. K., Clayton, R. N. and Gold nism of mineralogical reactions in carbonate for smith, J. R. (1991) Oxygen and carbon isotope frac tionation between CO2 and calcite. Geochim. mation. Cosmochim. Acta 55, 2867-2882. Chiba, H., Kusakabe, M., Hirano, S.-I., Matsuo S. and Acknowledgments-This study was supported by funds Somiya, S. (1981) Oxygen isotope fractionation fac from the Natural Science Foundation of China (No. 124 Y.-F. Zheng

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