Fayek & Kyser 2000 uraninite fractionations.pdf - Queen's University

PergamonPII S0016-7037(99)00393-2Geochimica et Cosmochimica Acta, Vol. 64, No. 13, pp. 2185–2197, 2000Copyright © 2000 Elsevier Science LtdPrinted in the USA. All rights reserved0016-7037/00 $20.00 .00Low temperature oxygen isotopic fractionation in the uraninite–UO 3 –CO 2 –H 2 O systemMOSTAFA FAYEK †, * and T. KURTIS KYSERDepartment of Geological Sciences, Queens University, Kingston, Ontario, Canada K7L 3N6(Received July 7, 1998; accepted in revised form May 3, 1999)Abstract—Oxygen-isotope fractionation factors for uraninite–water and UO 3 –water were determined from aseries of uraninite–CO 2 and UO 3 –CO 2 exchange experiments between 100 and 300°C. The measuredfractionation factors are similar to the theoretical fractionation factors of Hattori and Halas (1982) at hightemperatures and Zheng (1991) at low temperatures and are nonlinear over the temperature range examined.Regression through these curve gives 1000 ln UO2 –H 2 O 16.58 (10 6 /T 2 ) 77.52 (10 3 /T) 77.48, 1000ln UO3 –H 2 O 2.21 (10 6 /T 2 ) 25.06 (10 3 /T) 45.90 with 2 errors of 2.5‰. Diffusion coefficientsof oxygen with temperature for uraninite and UO 3 can be represented by D (uraninite) 3.1 10 10 exp(63981/RT) and D (UO3 ) 1.3 10 15 exp (21,197/RT), where the activation energies for uraninite andUO 3 are 64.0 0.2 kcal/mole and 21.2 0.2 kcal/mol, respectively. Exceptionally low 18 O values of naturaluraninites (i.e., 32‰ to 19.5‰) from unconformity-type uranium deposits in Saskatchewan, in conjunctionwith theoretical and experimental uraninite–water and UO 3 –water fractionation factors, suggest thatprimary uranium mineralization is not in oxygen isotopic equilibrium with coeval clay and silicate minerals.The low 18 O values have been interpreted as having resulted from the low temperature recrystallization ofprimary uranium mineralization in the presence of relatively modern meteoric fluids having 18 O values of ca.18‰, despite petrographic and U–Pb isotope data that indicate limited alteration. Therefore, it appears thaturanium minerals can exchange oxygen isotopes with later fluids, with only limited actinide remobilizationand minor disturbances to their original chemical compositions and textures. Copyright © 2000 ElsevierScience Ltd1. INTRODUCTIONIsotope fractionation in solids depends primarily on thenature of the bonds between the isotopic species of an elementand the nearest atom in the structure. Bonds with small forceconstants, involving atoms with large masses, produce lowvibrational frequencies and tend preferentially to incorporatelight isotopes (Hattori and Halas, 1982; O’Neil, 1986). Forexample, heavy-element sulfides such as galena (PbS) are enrichedin the light isotope of sulfur ( 32 S) (Bachinski, 1969).O’Neil et al. (1969) and O’Neil (1977; 1986) have also shownthat carbonates with heavier metal cations have lower 18 Ovalues, and Taylor and Epstein (1962) and Becker, (1971) haveshown that iron oxides (i.e., magnetite) are depleted in 18 Orelative to quartz.Uraninite, is the most common reduced uranium-bearingmineral. It is found in most uranium deposits, except whereoxidizing fluids have eliminated the reduced mineral phases(Smith, 1984). Uraninite, consists of oxygen and uranium,which is the heaviest element in nature and predictably incorporatesthe light isotope of oxygen ( 16 O). However, uraninitevaries widely in composition from UO 2.07 to UO 2.25 , and issusceptible to alteration and oxidation by oxidizing meteoricfluids (Grandstaff, 1976; Smith, 1984; Burns et al., 1996),which make its isotopic properties particularly complex.Exceptionally low 18 O values of natural uraninites (i.e.,28‰) from unconformity-type uranium deposits inSaskatchewan and high 18 O values of coexisting silicate and* Author to whom correspondence should be addressed (mfa@ornl.gov).† Present address: Department of Geological Sciences, University ofTennessee, Knoxville, TN 37919.2185clay minerals (i.e., 10‰ for quartz) are compatible with thismass effect (Hattori et al., 1978; Kotzer and Kyser, 1993;Fayek and Kyser, 1997). However, theoretical UO 2 –water andUO 3 –water fractionation factors proposed by Hattori and Halas(1982) and Zheng (1991; 1995) suggest that primary uraninitefrom unconformity-type uranium deposits is not in oxygenisotopic equilibrium with the fluids that precipitated coexistingclay and silicate minerals. In addition, a large difference existsbetween the theoretical fractionation factors proposed byZheng (1991; 1995) and those of Hattori and Halas (1982).The purpose of this study is to determine experimentalCO 2 –uraninite and CO 2 –UO 3 fractionation factors and calculateuraninite–water and UO 3 –water fractionation factors todetermine the effects of uraninite–water interaction on thechemical and oxygen isotopic composition of uraninite. Exchangebetween CO 2 and uranium minerals is used becausewater reacts so readily with these minerals (Fayek et al., 1996).These fractionation factors will be compared to the theoreticalUO 2 –water and UO 3 –water fractionation factors and applied tonatural uraninite samples from the Athabasca Basin in northernSaskatchewan.2.1. Starting Material2. EXPERIMENTAL TECHNIQUESA drill-core sample of uraninite was collected from a zone ofhigh-grade uranium mineralization from the McArthur River depositlocated along the eastern margin of the Athabasca Basin in northernSaskatchewan. The sample was examined for homogeneity and alterationusing reflected light microscopy, backscattered electron imaging(BSE), quantitative electron microprobe analyses, and oxygen isotopeanalyses. The 200–230 mesh (60–74 m) size fraction was separatedusing conventional crushing and sieving techniques, and purity of thesample was further verified by X-ray diffraction and scanning electron

Table 1. Starting material. aUraninite–CO 2Run #Temp. (°C)CO 2 mol 18 OCO 2 (‰) 13 C PDB CO 2 (‰) Ur (mg) 18 OUr(‰)Yield (UO 2 ) b(mol/mg) (Ur–CO 2 ) c1 100 73.2 110.6 11.6 12.68 25.1 4.7 91.82 100 58.5 110.6 11.6 10.23 25.1 4.7 91.83 100 71.3 110.6 11.6 9.54 25.1 4.7 91.84 100 76.3 110.6 11.6 11.65 25.1 4.7 91.86 100 74.1 31.6 11.3 11.55 25.1 4.7 56.57 100 72.3 31.6 11.3 11.64 25.1 4.7 56.58 100 74.4 31.6 11.3 10.33 25.1 4.7 56.59 100 67.7 31.6 11.3 11.35 25.1 4.7 56.510 200 73.7 108.0 11.6 14.99 25.1 4.7 88.911 200 55.1 108.0 11.6 12.20 25.1 4.7 88.912 200 65.5 108.0 11.6 10.59 25.1 4.7 88.913 200 59.0 108.0 11.6 10.65 25.1 4.7 88.914 200 69.7 108.0 11.6 11.22 25.1 4.7 88.915 200 69.1 108.0 11.6 13.10 25.1 4.7 88.916 200 70.3 31.6 11.3 13.44 25.1 4.7 56.517 200 71.9 31.6 11.3 13.24 25.1 4.7 56.518 200 71.9 31.6 11.3 13.16 25.1 4.7 56.519 200 72.3 31.6 11.3 12.14 25.1 4.7 56.520 200 67.7 31.6 11.3 14.39 25.1 4.7 56.521 200 67.8 31.6 11.3 11.83 25.1 4.7 56.522 300 73.3 108.0 11.6 9.73 25.1 4.7 88.923 300 62.3 108.0 11.6 10.89 25.1 4.7 88.924 300 77.0 108.0 11.6 10.84 25.1 4.7 88.925 300 70.3 108.0 11.6 11.87 25.1 4.7 88.926 300 67.9 108.0 11.6 12.13 25.1 4.7 88.927 300 80.7 108.0 11.6 14.05 25.1 4.7 88.928 300 71.5 31.6 11.3 14.94 25.1 4.7 56.529 300 71.3 31.6 11.3 9.50 25.1 4.7 56.530 300 73.3 31.6 11.3 14.47 25.1 4.7 56.531 300 69.8 31.6 11.3 12.14 25.1 4.7 56.532 300 70.1 31.6 11.3 9.65 25.1 4.7 56.533 300 66.3 31.6 11.3 13.55 25.1 4.7 56.534 100 62.3 75.0 11.5 20.68 25.1 4.7 52.535 200 70.6 75.0 11.5 17.03 25.1 4.7 52.536 300 76.9 75.0 11.5 15.18 25.1 4.7 52.537 300 81.0 75.0 11.5 12.02 25.1 4.7 52.5Uraninite-CO 2Yield (UO 3 ) bRun # Temp. (°C) CO 2 mol 18 OCO 2 (‰) UO 3 –CO 2 13 C PDB CO 2 (‰) UO 3 (mg) 18 OUO 3 (‰) (mol/mg) (UO 3 –CO 2 ) c1a 100 70.7 88.0 11.4 14.37 5.0 5.4 87.12a 100 75.2 88.0 11.4 14.25 5.0 5.4 87.13a 100 73.6 88.0 11.4 15.74 5.0 5.4 87.14a 100 73.2 88.0 11.4 12.40 5.0 5.4 87.15a 100 72.6 88.0 11.4 15.63 5.0 5.4 87.16a 100 72.5 88.0 11.4 12.89 5.0 5.4 87.17a 100 78.4 29.0 11.3 14.53 5.0 5.4 33.68a 100 80.0 29.0 11.3 18.22 5.0 5.4 33.69a 100 81.3 29.0 11.3 18.79 5.0 5.4 33.610a 100 80.5 29.0 11.3 16.48 5.0 5.4 33.611a 100 81.8 29.0 11.3 16.49 5.0 5.4 33.612a 100 82.3 29.0 11.3 16.03 5.0 5.4 33.613a 200 70.4 101.0 11.4 16.20 5.0 5.4 101.514a 200 71.4 101.0 11.4 17.12 5.0 5.4 101.515a 200 75.3 101.0 11.4 12.21 5.0 5.4 101.516a 200 74.4 101.0 11.4 12.28 5.0 5.4 101.517a 200 70.3 101.0 11.4 13.05 5.0 5.4 101.518a 200 76.0 26.0 11.3 11.42 5.0 5.4 30.719a 200 75.7 26.0 11.3 11.99 5.0 5.4 30.720a 200 77.3 26.0 11.3 13.76 5.0 5.4 30.721a 200 76.8 26.0 11.3 12.84 5.0 5.4 30.722a 200 77.4 26.0 11.3 12.12 5.0 5.4 30.723a 300 57.1 88.0 11.4 16.33 5.0 5.4 87.124a 300 58.1 88.0 11.4 11.93 5.0 5.4 87.125a 300 56.6 88.0 11.4 12.59 5.0 5.4 87.126a 300 56.8 88.0 11.4 14.15 5.0 5.4 87.127a 300 58.0 88.0 11.4 12.10 5.0 5.4 87.128a 300 58.5 88.0 11.4 10.52 5.0 5.4 87.129a 300 70.9 88.0 11.4 12.68 5.0 5.4 87.130a 300 76.9 27.0 11.3 15.56 5.0 5.4 31.731a 300 75.6 27.0 11.3 12.95 5.0 5.4 31.732a 300 78.8 27.0 11.3 11.26 5.0 5.4 31.733a 300 80.6 27.0 11.3 11.83 5.0 5.4 31.734a 300 97.2 27.0 11.3 11.90 5.0 5.4 31.735a 300 78.8 27.0 11.3 11.66 5.0 5.4 31.736a 300 80.7 27.0 11.3 11.71 5.0 5.4 31.737a 100 49.5 66.0 11.3 15.36 5.0 5.4 63.338a 200 62.9 66.0 11.3 13.02 5.0 5.4 63.339a 300 62.2 66.0 11.3 12.02 5.0 5.4 63.3a Abbreviations: Temp. temperature; Ur uraninite.b Error on yield is 0.4 mol/mg (2).c ( A–B) 10 3 ln (A–B) .

Low temperature oxygen isotopic fractionation2187Table 2. Results and conditions of oxygen isotope exchange experiments between silica tubes and CO 2 . aBatchTime(h)Temp.(°C)mol CO 2(initial)mol CO 2(final) 18 OCO 2b(initial) 18 OCO 2b(final) 13 C PDB CO 2(initial) 13 C PDB CO 2(final)35b 525 100 55.0 55.6 113.7 112.5 11.7 11.7311b 7200 100 54.3 54.8 110.6 111.6 11.6 11.638b 525 100 77.8 78.1 31.2 31.6 11.2 11.2312b 7200 100 69.6 69.3 31.6 31.5 11.2 11.236b 525 200 57.0 58.0 113.7 112.5 11.7 11.6313b 4500 200 55.6 56.2 109.0 108.0 11.7 11.839b 525 200 77.0 77.4 31.2 31.7 11.2 11.2314b 4500 200 74.5 75.4 31.0 31.0 11.2 11.51b 36 300 62.1 62.5 114.6 115.0 11.7 11.72b 108 300 80.0 80.3 114.6 114.6 11.7 11.737b 346 300 58.3 59.0 113.7 111.7 11.7 11.73b 732 300 73.2 73.6 114.6 113.1 11.7 11.74b 36 300 63.5 64.7 31.6 31.2 11.2 11.35b 108 300 68.4 70.4 31.6 31.5 11.2 11.3400b 346 300 75.3 76.0 31.2 31.2 11.2 11.36b 732 300 67.5 69.5 31.6 31.3 11.2 11.3a Note: all isotopic values are in per mil.b Values are in per mil relative to VSMOW. Temp. temperature.microscopy (SEM). X-ray diffraction data showed that the uraninitestructure was consistent with the U 4 O 9 structure. Chemical compositionsof the uraninite (U 3 O 8 83.5–93.2; SiO 2 0.1–0.5; CaO 0–1.2; PbO13.4–16.5; FeO 0.1–1.3) was determined by wavelength dispersivespectroscopy (WDS) using a fully automated Joel JXA-8600 X-raymicro analyzer at an operating voltage 20 keV, a beam diameter of 2m, and counting times of 40 s per element. Detection limits of theelements were on the order 0.1 wt%. Data reduction for the variouselements was done by taking into account the matrix correctionsbetween standards and samples and the analytical parameters (i.e., takeoff angle and operating voltage) of the electron microprobe.Fine grained (5–15 m) synthetic UO 3 powder was purchased fromK & K laboratories, Hollywood, CA. Purity of the sample was verifiedby X-ray diffraction and scanning electron microscopy (SEM). Theinitial source of CO 2 gas was a tank of liquid CO 2 that has a 18 O valueof 31.6‰. CO 2 gas with several different 18 O values was preparedby equilibrating the “tank” CO 2 with 18 O depleted water, using conventionalCO 2 –H 2 O equilibration techniques of Epstein and Mayeda(1953). Repeat oxygen-isotope analyses of the starting materials over aperiod of six months were reproducible to better than 0.5‰ foruraninite and UO 3 , and 0.1‰ for the CO 2 gas (2). The oxygenisotopic composition of the starting material for each experiment islisted in Table 1.2.2. ExperimentsThe design of our experiments was similar in principal to thatemployed by O’Neil and Epstein (1966), Liu and Epstein (1984), andStolper and Epstein (1991). Each experiment consisted of a series ofcompanion exchange runs. The companion exchange runs were two orthree experimental runs that were similar in every respect except for thestarting oxygen isotopic composition of the CO 2 gas (Table 1).Each experimental run consisted of a silica tube (General Electric#214; 5 ppm OH ) that was 16 cm long, had an OD of 6 mm, and anID of 4 mm. The silica tubes were placed in an oven at a temperatureof 110°C for 1–2 weeks to remove water vapor. A sample of uraniniteor UO 3 powder (9.4–17.1 mg) was loaded in each silica tube, whichwere then attached to a vacuum line and flamed with a torch to removeexcess water vapor. CO 2 gas (49.5–83.2 mols) was frozen into eachtube along with the uraninite or UO 3 sample, after which the tubes weresealed with a torch. Tubes were loaded into the hot-spots of home built,wire wound, vertical resistance furnaces and held at 100°C, 200°C, and300°C. Temperatures were monitored with type K thermocouples,controlled by a Powerstat L 116°C variable autotransformer, and typicallyvaried by less than 1°C. Temperature gradients were such that thetemperature varied by at most 5°C over the length of each sample tube,and typically less than 2°C over the segment of the tube containing theuraninite or UO 3 sample. Several experiments were conducted bysealing 55–77 mols of CO 2 gas in empty silica tubes and holdingthese tubes at 100–300°C for up to 6 months to examine the kinetics ofoxygen isotope exchange between CO 2 gas and the silica tubes (Table2). In addition, several more experiments were conducted by sealing9.4–17.1 mg of uraninite or UO 3 in silica tubes without any CO 2 gasand holding these tubes at 100–300°C for up to 6 months (Table 2).The synthetic UO 3 powder produced a small amount of gas (2–9mols) during initial heating which contained CO 2 with a 18 O valueof 35‰. Therefore, the initial oxygen isotopic composition of theCO 2 in the experiments that contained UO 3 was adjusted accordingly.The magnitude of the correction ranged between 0.1‰ and 15‰ forCO 2 with 31.7‰ and 108‰, and for experiments at 100°C and300°C, respectively.2.3. Analytical TechniquesCompanion exchange runs were periodically removed from thefurnaces over a period of 6 months and allowed to cool in air. The CO 2gas from each tube was collected and measured manometrically, and its 18 O value and 13 C value analyzed on a Finnigan Mat 252 gas sourcemass spectrometer at Queen’s University. Oxygen was liberated fromthe uranium oxide minerals using the BrF 5 technique of Clayton andMayeda (1963). Oxygen and carbon isotopic compositions are reportedin units of ‰ relative to Vienna Standard Mean Ocean Water (V-SMOW) and Pee Dee belemnite (VPDB), respectively. Replicate analysesfor 18 O and 13 C values are reproducible to / 0.1‰, and the 18 O value of NIST-28 quartz is 9.6‰. The oxygen yields for uraninite(U 4 O 9 ) and UO 3 are 4.7 0.4 and 5.4 0.4, respectively.As the purpose of this work was to determine oxygen-isotope fractionationbetween CO 2 and uraninite or UO 3 , it was important toestablish that these phases were stable at run conditions. Startingmaterial and several run products were examined optically, and bytransmission electron microscopy (TEM, SEM, and XRD). No changesor growth of new phases were detected. Uraninite and UO 3 in thepresence of oxidizing water readily alter to uranyl oxide hydrateminerals (Smith, 1984; Fayek et al., 1996). Therefore, direct measurementof oxygen-isotope exchange between uraninite or UO 3 and H 2 Ois not possible because these phases would not be stable under runconditions. The oxygen-isotope fractionation factors determined experimentallyfor CO 2 –uraninite and CO 3 –UO 3 in conjunction with theCO 2 –H 2 O fractionation factor of Truesdell (1974) (1.029 24, 1.018 23,and 1.012 47 at 100°C, 200°C, and 300°C, respectively) were used tocalculate uraninite– and UO 3 –H 2 O fractionation factors.

2188 M. Fayek and T. K. KyserTable 3. Uraninite–CO 2 equilibration results. aRun #Temp.(°C)Time(h) 18 OCO 2b(‰)mol CO 2(final) 18 OUr b(‰)Yield (Ur)(mol/mg)M Bal c(‰) (int.) d(‰) (fin.) d(‰)1 100 212 107.8 72.4 26.7 4.6 0.76 91.8 87.02 100 708 104.2 57.4 32.0 4.6 0.20 91.8 77.53 100 4800 101.5 71.6 34.5 5.0 1.68 91.8 71.94 100 7200 103.0 77.0 32.5 4.7 1.11 91.8 75.76 100 212 31.2 75.2 24.7 4.7 0.05 56.5 55.77 100 708 31.2 73.3 25.1 4.7 0.22 56.5 56.18 100 4800 30.1 75.7 22.0 5.4 0.39 56.5 51.99 100 7200 31.7 67.7 23.0 4.6 1.01 56.5 54.510 200 20 102.8 74.2 30.3 4.6 0.04 88.9 77.711 200 68 97.2 56.4 32.8 4.7 1.08 88.9 68.912 200 212 95.6 66.3 36.5 4.9 1.76 88.9 63.313 200 379 94.2 60.1 39.4 4.7 0.47 88.9 58.714 200 1512 83.5 69.7 53.6 4.9 0.87 88.9 32.115 200 4500 74.2 69.1 61.5 4.6 0.36 88.9 13.616 200 20 31.0 70.8 24.2 4.7 0.13 56.5 55.017 200 68 29.3 74.2 22.4 4.8 0.10 56.5 51.518 200 212 28.4 73.9 20.9 5.0 0.34 56.5 49.119 200 379 29.0 73.5 21.2 4.8 0.37 56.5 50.020 200 1512 23.3 70.2 18.9 4.6 0.83 56.5 42.121 200 4500 23.4 70.3 16.9 4.7 0.56 56.5 40.222 300 15 94.5 75.9 44.2 4.9 0.50 88.9 54.123 300 40 86.7 65.2 49.6 4.6 0.06 88.9 39.824 300 86 88.7 79.8 51.0 4.7 0.63 88.9 40.525 300 240 77.1 72.2 64.2 4.5 1.14 88.9 13.926 300 398 77.4 69.3 58.2 4.5 0.55 88.9 20.627 300 732 74.3 83.2 68.5 4.5 2.18 88.9 6.228 300 15 25.0 73.7 19.0 4.7 0.11 56.5 43.929 300 40 23.8 73.9 17.2 4.7 1.52 56.5 40.930 300 86 23.7 76.3 15.8 4.7 0.64 56.5 39.331 300 240 19.0 70.7 12.2 4.6 0.74 56.5 31.132 300 398 20.7 72.2 13.4 4.7 1.69 56.5 34.033 300 732 19.9 67.5 13.7 4.5 0.03 56.5 33.534 100 7200 61.9 61.9 33.5 5.0 0.32 52.2 29.835 200 4500 47.2 71.2 35.6 6.0 6.86 52.2 12.136 300 732 65.2 78.2 36.4 4.8 0.68 52.2 30.337 300 15 60.9 81.0 42.1 4.8 0.85 52.2 19.8a Abbreviations: Temp. temperature; h hours; Ur uraninite; int. initial; fin. final.b 18 O of the products.c Calculated mass balance: (X 18 O) (reactants) .d 10 3 ln Ur–CO2 .3. RESULTSThe results of the exchange experiments are shown in Tables3 and 4. Oxygen-isotope mass balance for each run was calculatedby summing the product of the oxygen mass fraction andthe 18 O/ 16 O ratio for each reactant and subtracting the resultfrom a similar sum from the product. A random distribution ofpositive and negative mass balance values were calculated asexpected. Propagation of analytical errors using the method ofRees (1984) suggests that mass balance should be 0.0 1.1‰.Most runs fall within this bound, but some do not. Therandom, trace impurities in the uraninite and UO 3 startingmaterial may account for the largest discrepancies. Naturaluraninite used in these experiments contains minor amounts ofCA, Si, REEs, and Th (Pagel and Ruhlmann, 1985; Kotzer andKyser, 1993; Fayek and Kyser, 1997). High concentration (1–2wt%) of these elements may affect the rate of oxygen isotopeexchange as well as form stable oxide phases which are difficultto detect by XRD, electron microprobe or SEM becausethey are fine grained. The synthetically prepared UO 3 containedsome residual gas impurities. Heating the UO 3 powder at200°C and 300°C for 15 hours caused the sample to changecolor from bright yellow to mustard yellow. This was likely dueto the release of the residual gas. These impurities also resultedin rather variable yields (0.5 mol O 2 /mg) (Tables 3 and 4).Experiments where CO 2 gas was sealed in empty silica tubesand held at 100–300°C for up to 6 months showed that reactionbetween the CO 2 gas and silica tubes was minimal. Only smallshifts in the 18 O values of the CO 2 gas (0.8‰) occurred insamples where the 18 O values of the CO 2 gas greatly differed(100‰) from that of the silica tubes (Table 2). Moreover,diffusion of oxygen through silica glass is considerably slower(Stolper and Epstein, 1991) than diffusion of oxygen throughuraninite or UO 2 (Auskern and Belle, 1961). Therefore, CO 2loss did not occur as confirmed by our manometric measurements(Table 2).Some of these potential sources of error were difficult toquantify and increased the overall analytical error of the experiments.However, after careful consideration of thesesources of experimental error and their effect on the measureduraninite–CO 2 and UO 3 –CO 2 fractionations, only runs with a

Low temperature oxygen isotopic fractionation2189Table 4. UO 3 -CO 2 equilibration results. aRun #Temp.(°C)Time(h) 18 OCO 2b(‰)mol CO 2(final) 18 OUO 3b(‰)Yield (UO 3 )(mol/mg)M Bal c(‰) (int.) d(‰) (fin.) d(‰)1a 100 20 71.0 69.2 19.4 5.2 0.87 87.1 54.12a 100 39 83.1 72.4 15.8 5.2 2.89 87.1 70.83a 100 136 68.8 68.5 19.6 5.2 1.40 87.1 51.54a 100 176 76.6 69.2 17.9 5.2 0.02 87.1 61.65a 100 349 69.3 69.9 20.6 5.3 0.42 87.1 51.06a 100 525 73.4 70.5 23.1 5.2 1.11 87.1 52.97a 100 20 33.6 76.5 1.8 5.3 3.91 33.6 34.88a 100 39 30.2 78.3 5.8 5.2 0.12 33.6 35.69a 100 136 29.5 79.2 6.1 5.2 0.37 33.6 35.210a 100 176 29.5 78.7 5.0 5.7 0.23 33.6 34.111a 100 349 29.7 80.4 4.7 5.7 0.49 33.6 34.012a 100 525 29.5 80.4 5.1 5.4 0.19 33.6 34.213a 200 3.5 49.3 70.6 43.7 5.5 1.20 101.5 5.914a 200 7.5 46.2 71.7 42.3 5.2 3.69 101.5 4.115a 200 25 57.2 74.9 57.4 5.3 0.66 101.5 0.216a 200 384 53.2 78.4 60.6 5.2 0.04 101.5 7.817a 200 525 50.4 72.1 54.6 5.4 0.44 101.5 4.418a 200 3.5 24.2 76.7 4.0 5.5 0.53 30.7 27.919a 200 7.5 23.3 76.4 2.0 5.4 0.07 30.7 25.020a 200 25 22.8 78.2 2.8 5.3 0.58 30.7 25.321a 200 384 22.5 77.2 1.5 5.2 0.24 30.7 23.822a 200 525 22.8 79.0 1.9 5.7 0.23 30.7 24.423a 300 3.5 37.8 58.3 42.0 5.4 2.73 87.1 4.424a 300 7.5 48.2 60.6 49.6 5.2 3.78 87.1 1.525a 300 15 41.1 62.2 40.9 5.9 0.09 87.1 0.226a 300 36 37.5 61.6 42.7 5.3 0.31 87.1 5.427a 300 133 41.0 64.6 47.5 5.4 0.41 87.1 6.828a 300 173 46.7 63.2 43.9 5.8 0.36 87.1 2.929a 300 346 46.6 76.7 51.7 5.8 3.46 87.1 5.430a 300 3.5 22.9 79.2 0.5 5.5 0.96 31.7 22.131a 300 7.5 27.7 78.4 5.5 5.2 5.32 31.7 21.832a 300 15 23.4 84.4 0.3 5.6 0.09 31.7 23.433a 300 36 23.4 85.8 1.3 5.2 0.44 31.7 24.434a 300 133 24.3 103.5 0.6 5.8 0.75 31.7 23.435a 300 173 23.2 85.5 0.4 5.6 0.37 31.7 22.536a 300 346 23.1 86.3 0.3 5.5 0.08 31.7 23.137a 100 525 35.3 50.0 24.5 5.5 0.95 63.3 11.138a 200 525 28.1 64.9 42.1 5.2 0.98 63.3 14.539a 300 346 27.1 66.1 39.5 5.6 0.75 63.3 12.8a Abbreviations: Temp. temperature; h hours; int. initial; fin. final.b 18 O of the products.c Calculated mass balance: (X 18 O) (products) –(X 18 O) (reactants) .d 10 3 ln (UO3 –CO 2 ).mass balance within 1.5‰ of 0‰, the majority of our data,were included.At each temperature there are at least one set of pairedexperiments of similar duration where the mineral–CO 2 gasfractionation of the starting material was opposite in sign.Using the equations of Criss et al. (1987) and Northorp andClayton (1966), equilibrium uraninite–CO 2 and UO 3 –CO 2fractionation and percent exchange for each set of runs at agiven temperature were calculated (Table 5). The oxygen isotopefractionation factors determined experimentally for uraninite–CO2 and UO 3 –CO 2 in conjunction with the CO 2 –H 2 Ofractionation factor of Truesdell (1974) were used to calculateuraninite– and UO 3 –H 2 O fractionation factors (Table 5).The oxygen isotopic composition of uraninite and CO 2changes with time and are shown to approached equilibriumfrom two opposite directions at relatively equal rates (Figs.1a,b,c). The UO 3 –CO 2 experiments show similar trends (Figs.2a,b,c), although the UO 3 –CO 2 experiments show a higherdegree of exchange relative to uraninite–CO 2 experiments (Table5). This likely occurred because the grain size of thesynthetic UO 3 powder was much smaller (5–15 m) relative tothe uraninite (60–74 m), and decreasing the grain size generallydecreases the time required to achieve a given degree ofexchange (Cole and Ohmoto, 1986).3.1. Fractionation FactorTheoretical fractionation factors for UO 3 –H 2 O and UO 2 –H 2 O of Hattori and Halas (1982) (1) and (2), and Zheng (1995)(3) and (4) are nonlinear over the temperature range of 100°Cto 300°C (Fig. 3). Regression of these curves gives1000 ln UO2–H 2O 3.6310 6 /T 2 13.2910 3 /T 4.42,(1)1000 ln UO3–H 2O 3.2910 6 /T 2 13.6310 3 /T 4.65,(2)

2190 M. Fayek and T. K. KyserTable 5. Experimentally determined fractionation factors. aTime (h)Uraninite–CO 2Uraninite–H 2 O 18 O (Ur–H2 O) (‰)Temp. (°C) Eq (‰) n % exchange7200 100 39.7 2.4 (2) 12 10.9 2.54500 200 30.1 2.4 (2) 63 12.1 2.5398 300 19.7 2.4 (2) 63 7.3 2.5Time (h)UO 3 –CO 2Temp. (°C) Eq (‰) n % exchangeUO 3 –H 2 O 18 O (UO3 –H 2 O) (‰)525 100 23.4 2.4 (3) 51 5.4 2.5525 200 29.8 2.4 (3) 86 2.8 2.5173 300 21.3 2.4 (3) 79 8.9 2.5a Determined from experimental data in Tables 3 and 4 using the least squares algorithm of Mahon (1996) to fit the equation of Northrop andClayton (1966) and propogate 2 errors. Equation (6d) in Criss et al. (1987) was also used to calculate equilibrium fractionation factors (Eq). n number of runs used in calculating Eq. 18 O (Ur–H2 O) and 18 O (UO3 –H 2 O) fractionations were calculated from values in this table and the fractionationfactors between CO 2 and H 2 O (Truesdell, 1974). Ur uraninite.1000 ln UO2–H 2O 2.7310 6 /T 2 13.5310 3 /T 4.54,(3)1000 ln UO3–H 2O 4.1110 6 /T 2 10.2210 3 /T 3.065.(4)Oxygen isotope fractionation factors determined experimentallyfor our CO 2 –uraninite and CO 3 –UO 3 experiments, togetherwith the CO 2 –H 2 O fractionation factor of Truesdell(1974), were used to calculate uraninite–H 2 O and UO 3 –H 2 Ofractionation factors. Regression through these curves gives1000 ln UO2–H 2O 16.5810 6 /T 2 77.5210 3 /T 77.48,(5)1000 ln UO3–H 2O 2.2110 6 /T 2 25.0610 3 /T 45.90(6)with 2 errors of 2.5‰. Over the temperature range of 100°Cto 300°C the calculated fractionation factors have concaveparabolic shapes with minima at 200°C (uraninite–H 2 O) and300°C (UO 3 –H 2 O) (Fig. 3). Fractionation factors from thisstudy (Fig. 3) are similar to those of Zheng (1995) at lowtemperatures (100–200°C) and those of Hattori and Halas(1982) at higher temperatures (300°C).The nonlinearity is not well understood because isotopefractionation factors at high temperatures (400°C) shouldvary linearly as a function of 1/T 2 whereas they should varylinearly as a function of 1/T at low temperatures (O’Neil, 1986;Criss, 1991). This theoretical prediction is based on the implicitassumption that the vibrational modes of the phases involvedremain constant over the temperature range examined. Thismay not be applicable to uranium oxides because oxygen isfound in interstitial and lattice sites. The fluorite-type structurerepresenting uraninite has holes approximately the size of theoxygen ion at (1/2, 0, 0) and 1/2, 1/2, 1/2) positions which arefilled with oxygen atoms in the end-member UO 3 . This is rarelythe case for natural uraninites, which generally have a structurethat is consistent with the U 4 O 9 (UO 2.25 ) structure, where thesesites are only partially filled (Berman, 1957). The distributionand amount of oxygen in these sites is dependent on severalfactors including the degree of oxidation of the mineral and theREE/ThO 2 ratio (Berman, 1957). These different domainslikely have different activation energies. Therefore, at low verytemperatures (100°C) the interstitial sites may be active relativeto the lattice sites because of lower activation energies. Atincreasingly higher temperatures both domains may becomeactive (see “Diffusion” section below). Therefore, vibrationalmodes of the oxygen molecules from these different domainsmay not remain constant with temperature.3.2. DiffusionThere are three main processes by which oxygen isotopesexchange between minerals and a fluid or gas: (1) solutionprecipitation, wherein the mineral partially dissolves, then recrystallizesas other portions of the mineral go into solution; (2)chemical reaction, where the activity of a component in thefluid is sufficiently different from the activity of the samecomponent so that a finite portion of the mineral breaks downand a new crystal forms which is in isotopic equilibrium withthe fluid; this process continues, thereby advancing the reactionfront into the crystal because the newly formed phase is morepermeable than the original; and (3) diffusion, wherein themineral remains as a single, solid reservoir with constant dimensionsand shape, and isotopes exchange by volume diffusionaltransport (Giletti, 1985). The principal difference amongthe three processes is whether or not isotope exchange occursby the dissolution of the mineral.Oxygen exchange between uraninite/UO 3 and CO 2 in thisstudy most likely occurred by volume diffusion because XRD,SEM, and TEM analyses of run products showed no signs ofreaction (formation of a new phase), recrystallization (changein grain size or morphology) or oxidation in the case of uraninite.Oxygen diffusion likely occurred when an interstitialand lattice oxygen moved in combination, where a latticeoxygen atom moves into a neighboring vacant interstitial sitefollowed by an interstitial oxygen moving into the vacatedlattice site. Thus equilibrium and diffusion occur in the samestep (Secco, 1958). The mobility of lattice oxygen atoms isconsidered slower than the mobility of interstitial oxygen.Diffusion coefficients (Table 6) were calculated using the

Low temperature oxygen isotopic fractionation2191Fig. 1. Representative fractionation as a function of time for (a) uraninite–CO 2 at 100°C, (b) uraninite–CO 2 at 200°C, and(c) uraninite–CO 2 at 300°C showing initial (open symbols) and final (solid symbols) isotopic differences for each run. Linesindicate equilibrium uraninite–CO 2 fractionations. Arrows indicate direction of isotopic exchange from initial to finalfractionation. Data are from Table 3. Symbol size is indicative of 2 error.

2192 M. Fayek and T. K. KyserFig. 2. Representative fractionation as a function of time for (a) UO 3 –CO 2 at 100°C, (b) UO 3 –CO 2 at 200°C, and (c)UO 3 –CO 2 at 300°C showing initial (open symbols) and final (solid symbols) isotopic differences for each run. Linesindicate equilibrium UO 3 –CO 2 fractionations. Arrows indicate direction of isotopic exchange from initial to final fractionation.Data are from Table 4. Symbol size is indicative of 2 error.

Low temperature oxygen isotopic fractionation2193Fig. 3. Oxygen isotopic fractionation as a function of temperature between uraninite and water, and UO 3 and water. Dataare from Table 5. Errors for curves from this study are 2.5‰. Abbreviations: H&H Hattori and Halas (1982).

2194 M. Fayek and T. K. KyserTable 6. Diffusion coefficients of oxygen in uranium oxides.Sample Temp. (°C) D (cm 2 /s)Uraninite 100 3.10E-19200 3.34E-17300 3.28E-16UO 3 100 1.31E-18200 7.52E-18300 1.34E-17aUO 2.002 550 1.90E-16591 4.90E-15625 1.30E-14647 5.30E-14666 1.80E-13712 3.00E-13748 1.50E-12780 3.10E-12a Data from Auskern and Belle (1961).Temp. temperature.data in Tables 1, 3, 4, 5, and the equations of Crank (1975) fordiffusion in a sphere. The coefficients are listed in Table 6. Thelogarithms of the diffusion coefficients from this study andfrom Auskern and Belle (1961) for stoichiometric UO 2 areplotted versus the reciprocal temperature in Fig. 4. The diffusioncoefficients between 100°C and 300°C for uraninite andUO 3 , and between 550°C and 780°C for stoichiometric UO 2 ,can be described by three linear trends. Linear regressionthrough these curves and 2 errors were calculated using theprogram Origin 4.1. The principal source of error in the diffusioncoefficient measurements is likely variation in particlesize. The smaller particles may reach equilibrium early in theexchange process whereas at higher degrees of exchange, theexchange rate slows down because diffusion kinetics are governedby the larger particles.The variation of diffusion coefficient with temperature foruraninite (7) and UO 3 (8) can be represented by the equationsFig. 4. An Arrhenius plot of D versus 10 3 /T of oxygen diffusion coefficients for stoichiometric UO 2 , uraninite and UO 3 .Data for UO 2.002 is from Auskern and Belle (1961). Data for uraninite and UO 3 are from Table 6.

Low temperature oxygen isotopic fractionation2195D 3.1 10 10 exp 63,981/RT, (7)D 1.3 10 15 exp 21,197/RT. (8)The activation energies for uraninite and UO 3 are 64.0 0.2kcal/mole and 21.2 0.2 kcal/mol, respectively. The activationenergy for oxygen in uraninite is very similar to the activationenergy reported by Auskern and Belle (1961) for stoichiometricUO 2 (65.3 5 kcal/mol) and is higher than the activationenergy for UO 3 which suggests that U 6 –O bond is weakerthan the U 4 –O bond and the U 6 –O 3 molecule has differentinternal energies, heat capacities and entropies from the U 4–O 2molecule. The slightly lower activation energy for oxygen inuraninite (U 4 O 9 ) relative to that in stoichiometric UO 2 (i.e.,change in slope between high and low temperature results forstoichiometric UO 2 and uraninite, respectively) suggest that theweaker U 6 –O bonds are present in the more oxidized uraninitestructure (U 4 O 9 ), which would effectively lower theactivation energy.4. DISCUSSIONThe fractionation factor curves between uranium oxides andwater (Fig. 3) have a broad minima at 200°C, which isuncommon for most common rock-forming minerals curves.This inflection is likely caused by the high vibrational frequenciesof water which are slightly shifted by isotopic substitutionand very the low vibrational frequencies of uranium oxides,which are significantly shifted by isotopic substitution (Stern etal., 1968).Fractionation factor curves between other oxide minerals(e.g., magnetite, hematite, quartz, rutile, and cassiterite) andwater also show a broad minima around 200°C. Among theseminerals, the equilibrium oxygen isotopic fractionation betweenquartz and water is largest because Si is the lightestelement of those combined with oxygen (see Hattori and Halas,1982; O’Neil, 1986; Kyser, 1987). The rule of “mass effect” onthe isotope fractionation (i.e., the heavy isotopes are enriched inspecies containing the lighter atoms) seems generally to be thecase. However, some of the oxides such as magnetite, hematite,UO 3 and uraninite, and rutile, cassiterite, have similar fractionationfactors (see Hattori and Halas, 1982; O’Neil, 1986; Kyser,1987; Zheng, 1991), although the masses of the metals are verydifferent. Unlike carbonates which have the same crystal structureand where the rule of the “mass effect” works for oxygenisotope exchange, these oxides have different crystal structures.This implies that the “mass effect” on the isotope fractionationcannot be applied to mineral groups of completely differentstructures in which the nature of the bonds and their nearestatom are quite different. Similarly, sulfide minerals of differentcrystal structure do not follow the “mass effect” rule (Huberten,1980).Theoretical fractionation factors of Hattori and Halas (1982)and Zheng (1995) indicate that the fractionation between UO 2and water at 200°C is approximately 11‰. The dominantfluids that equilibrated with silicate and clay minerals in unconformity-typeuranium deposits, which are associated withprimary uranium mineralization, are 200°C, saline fluids having 18 O values of 4 2‰ (Kotzer and Kyser, 1990; 1993;1995). Therefore, using these theoretical fractionation factorsin conjunction with the 18 O value of 4 2‰ for the fluidassociated with uranium mineralization, primary unaltered uraninitefrom the unconformity-type uranium deposits shouldhave 18 O values near 11 2‰. However, recent studies ofuranium minerals from unconformity-type uranium deposits inSaskatchewan, Canada, indicate that high reflectance unalteredprimary uranium mineralization have 18 O values from 32‰to 19.5‰ (Hoekstra and Katz, 1956; Hattori et al., 1978;Kotzer and Kyser, 1993; 1995; Fayek and Kyser, 1997). Theanomalously low 18 O values of primary uranium mineralizationare interpreted to have resulted from the interaction betweenuranium minerals and relatively recent, low temperaturemeteoric fluids with 18 O values of ca. 18‰ (Hoekstra andKatz, 1956; Hattori et al., 1978; Kotzer and Kyser, 1990; 1993).If true, the high reflectivities and lack of chemical alterationsuch as elevated Si and Ca contents of primary uranium minerals,are interpreted to indicate that uranium minerals canexchange oxygen isotopes with fluids resulting in only minordisturbances to their chemical composition and texture.Alternatively, oxygen isotope fractionation between uraniniteand water at 200°C may be greater than 11‰. Thetheoretical fractionation factors may not be applicable to naturalsystems because they use end member UO 2 and UO 3 intheir calculations and do not take into account the complexchemistry of natural uraninite (Janezcek and Ewing, 1992).However, the experimental uraninite–water and UO 3 –waterfractionation factors from this study are similar to those ofHattori and Halas (1982) and Zheng (1995). Therefore, thesefractionation factors in conjunction with the low 18 O values ofca. 28‰ for natural uraninite samples from unconformitytypeuranium deposits in Saskatchewan, indicate that the oxygenisotopic composition of the uraninite was completely overprintedby low-temperature ground water and that the uraninitecan exchange oxygen isotopes with fluids resulting in onlyminor disturbances to their chemical composition and texture.However, these fluids must be reducing because uranium mineralchemistry and solubility is largely a function of fO 2 anduraninite is stable only under very reducing conditions (fO 2 10 25 , Romberger, 1985; Cramer, 1995).Groundwater near the ore zone at the Cigar Lake unconformity-typeuranium deposit in Saskatchewan is characterized bylow fO 2 values (10 65 ) and low uranium concentrations(30 ppm or 10 6.5 mol/l; Cramer and Nesbitt, 1994; Cramer,1995). Therefore, under the reducing conditions that exist nearthese unconformity-type uranium deposits, interaction betweenuraninite and ground waters likely occurred, overprinting theoxygen isotopic composition of the uraninite, but with limitedactinide remobilization and only minor changes to the chemistryand texture of the uranium ore.Spent nuclear fuel which consists largely of UO 2 and uraninitehave many similarities which include topologically identicalstructures, high resistance to radiation damage and similarbehavior during oxidation (Janeczek et al., 1996). Therefore,underground repositories for the disposal of high level nuclearwaste should be located in areas where the groundwater isreducing so that spent nuclear fuel is most stable.

2196 M. Fayek and T. K. Kyser5. CONCLUSIONSThe experimental oxygen-isotope fractionation factors foruraninite–water and UO 3 –water are similar to the theoreticalfractionation factors of Hattori and Halas (1982) at high temperaturesand Zheng (1991; 1995) at low temperatures and arenonlinear over the temperature range 100°C to 300°C. Thesefractionation factors, in conjunction with the exceptionally low 18 O values of natural uraninites (i.e., 32‰ to 19.5‰)from unconformity-type uranium deposits in Saskatchewansuggest that primary uranium mineralization is not in oxygenisotopic equilibrium with coeval clay and silicate minerals andthat the low 18 O values are a result of low temperaturerecrystallization of primary uranium mineralization in the presenceof relatively modern meteoric fluids. Therefore, it appearsthat uranium minerals can exchange oxygen isotopes with laterfluids, with only limited actinide remobilization and minordisturbances to their original chemical compositions and textures.Diffusion coefficients of oxygen with temperature foruraninite and UO 3 were calculated and indicate that the activationenergy for oxygen in uraninite is higher than the activationenergy for UO 3 which suggests that U 6 –O bond isweaker than the U 4 –O bond and the U 6 –O 3 molecule hasdifferent internal energies, heat capacities and entropies fromthe U 4 –O 2 molecule.Acknowledgments—The Natural Sciences and Engineering Council ofCanada (NSERC) supported this work via an operating grant to T. K.Kyser and a Post-Doctoral Fellowship to M. Fayek. Technical assistancewas expertly provided by Kerry Klassen and Aliona Valyashko.This paper benefited greatly for the reviews of Dr. R. E. Criss and Dr.F. J. Longstaffe.REFERENCESAuskern A. B. and Belle J. (1961) Oxygen self-diffusion in uraniumdioxide. J. Nucl. Mater. 3, 267–276.Bachinski D. J. (1969) Bond strength and sulfur isotopic fractionationin coexisting sulfides. Econ. Geol. 64, 56–65.Becker R. H. (1971) Carbon and oxygen isotope ratios in iron formationsand associated rocks from the Hamersley range of WesternAustralia and their implications. Ph.D. thesis, University of Chicago.Berman R. M. (1957) The role of lead and excess oxygen in uraninite.The American Mineralogist. J. Mineral. Soc. Amer. 42, 705–731.Burns P. C., Miller M. L., and Ewing R. C. (1996) U 6 minerals andinorganic phases: A comparison and hierarchy of crystal structures.Can. Min. 34, 845–880.Clayton R. and Mayeda T. K. (1963) The use of bromine pentafluoridein the extraction of oxygen from oxides and silicates for isotopicanalysis. Geochim. Cosmochim. Acta. 27, 43–52.Cole D. R. and Ohmoto H. (1986) Kinetics of isotopic exchange atelevated temperatures and pressures. In Stable Isotopes in HighTemperature Geological Processes (eds. J. W. Valley, H. P. Taylor,and J. R. O’Neil), Rev. Mineral. 16, 41–90, Mineralogical Society ofAmerica.Cramer J. J. (1995) The Cigar Lake uranium deposit: Analogue informationfor Canada’s nuclear fuel waste disposal concept. AECL Rep.11204, 32.Cramer J. and Nesbitt W. (1994) Groundwater evolution and redoxgeochemistry. In Final Report of the AECL/SKB Cigar Lake AnalogStudy (eds. J. Cramer and J. A. T. Smellie), AECL 10851, 191–207.Crank J. (1975) The Mathematics of Diffusion. 2nd ed. Oxford UniversityPress.Criss R. E. (1991) Temperature dependence of isotopic fractionationfactors. In Stable Isotope Geochemistry: A tribute to Samuel Epstein(eds. H. P. Taylor, J. R. O’Neil, and J. R. Kaplan), Geological Soc.Spec. Pub. 3, 11–16.Criss R. E., Gregory T. E., and Taylor H.P., Jr. (1987) Kinetic theoryof oxygen isotopic exchange between minerals and water. Geochim.Cosmochim. Acta 51, 1099–1108.Epstein S. and Mayeda T. (1953) Variation of the 18 O content of watersfrom natural sources. Geochim. Cosmochim. Acta 4, 213–224.Fayek M., Kyser T. K., Ewing R. C., and Miller M. L. (1996) Uraninite-waterinteractions in an oxidizing environment. Mat. Res.Symp. Porc. 465, 1201–1207.Fayek M. and Kyser T. K. (1997) Characterization of multiple fluidevents and rare-earth-element mobility associated with formation ofunconformity-type uranium deposits in the Athabasca Basin,Saskatchewan. Can. Min. 35, 627–658.Giletti B. J. (1985) The nature of oxygen transport with minerals in thepresence of hydrothermal water and the role of diffusion. Chem.Geol. 53, 197–206.Grandstaff D. E. (1976) A kinetic study of the dissolution of uraninite.Econ. Geol. 71, 1493–1506.Hattori K. and Halas S. (1982) Calculation of oxygen isotope fractionationbetween uranium dioxide, uranium trioxide and water.Geochim. Cosmochim. Acta 46, 1863–1868.Hattori K., Muehlenbachs K., and Morton D. (1978) Oxygen isotopegeochemistry of uraninites. Geol. Soc. Amer.–Geol. Soc. Can., Pgm.Abst. 10, A417.Hoekstra H. R. and Katz J. J. (1956) Isotope geology of some uraniumminerals. Geological Survey Professional Paper 300. Contributionsto the geology of uranium and thorium by the United States GeologicalSurvey and Atomic Energy Commission for the UnitedNations International Conference on Peaceful Uses of Atomic Energy,Geneva, Switzerland, pp. 543–547, IAEA.Huberten H. W. (1980) Sulfur isotope fractionations in the Pb–S, Cu–S,and Ag–S systems. Geochem. J. 14, 177–184.Janeczek J. and Ewing R. C. (1992) Structural formula of uraninite.J. Nucl. Mater. 190, 128–132.Janeczek J., Ewing R. C., Oversby V. M., and Werme L. O. (1996)Uraninite and UO 2 in spent nuclear fuel: A comparison. J. Nucl.Mater. 238, 121–130.Kotzer T. G. and Kyser T. K. (1990) Fluid history of the AthabascaBasin and its relation to uranium deposits. In Summary of investigations1990. Sask. Energy and Mines, Sask. Geol. Surv., Misc. Rep.90-4, 153–157.Kotzer T. G. and Kyser T. K. (1993) O, U and Pb isotopic and chemicalvariations in uraninite: Implications for determining the temporaland fluid history of ancient terrains. Amer. Mineral. 78, 1262–1274.Kotzer T. G. and Kyser T. K. (1995) Petrogenesis of the ProterozoicAthabasca Basin, northern Saskatchewan, Canada, and its relation todiagenesis, hydrothermal uranium mineralization and paleohydrogeology.Chem. Geol. 120, 45–89.Kyser T. K. (1987) Equilibrium fractionation factors for stable isotopes.In Stable Isotope Geochemistry of Low Temperature Fluids(ed. T. K. Kyser). Mineral. Assoc. Can. Short Course 13, 1–84,Mineral. Assoc. Can.Liu K. K. and Epstein S. (1984) The hydrogen isotope fractionationbetween kaolinite and water. Isotope Geosci. 2, 335–350.Mahon K. I. (1996) The new “York” Regression: Application of animproved statistical method to geochemistry. International Geol.Rev. 38, 293–303.Northrop D. A. and Clayton R. N. (1966) Oxygen isotope fractionationsin systems containing dolomite. J. Geol. 74, 174–196.O’Neil J. R. (1977) Stable isotopes in mineralogy. Phys. Chem. Mineral.2, 105–123.O’Neil J. R. (1986) Theoretical and experimental aspects of isotopefractionation. In Stable Isotopes in High Temperature GeologicalProcesses (eds. J. W. Valley, H. P. Taylor, and J. R. O’Neil) Rev.Mineral. 16, 1–37, Mineral Soc. Amer.O’Neil J. R. and Epstein S. (1966) Oxygen isotope fractionation in thesystem dolomite–calcite–carbon dioxide. Science 152, 198–201.O’Neil J. R., Clayton R. N. and Mayeda T. K. (1969) Oxygen isotopefractionation in divalent metal carbonates. J. Chem. Phys. 51, 5547–5558.

Low temperature oxygen isotopic fractionation2197Pagel M. and Ruhlmann F. (1985) Chemistry of uranium minerals indeposits and showings of the Carswell structure (Saskatchewan-Canada). In The Carswell Structure Uranium Deposits, Saskatchewan(ed. R. Laine, D. Alonso, and M. Svab). Geol. Assoc. Can. Spec.Paper 29, 153–164.Rees C. E. (1984) Error propagation calculations. Geochim. Cosmochim.Acta 48, 2309–2311.Romberger S. B. (1985) Transport and deposition of uranium in hydrothermalsystems of temperatures up to 300°C: Geological implications.In Uranium Geochemistry, Resources: The Institution ofMining and Metallurgy, 12–17.Secco E. A. (1958) Diffusion and exchange of zinc in crystalline zincsulfide. J. Chem. Phys. 29, 406.Smith D. K. (1984) Uranium mineralogy. In Uranium Geochemistry,Mineralogy, Geology, Exploration and Resources. Inst. Mining Metall.43–88.Stern M. J., Spindel W., and Monse E. U. (1968) Temperature dependenciesof isotope effects. J. Chem. Phys. 48, 2908–2919.Stolper E. and Epstein S. (1991) An experimental study of oxygenisotope partitioning between silica glass and CO 2 vapor. In StableIsotope Geochemistry: A Tribute to Samuel Epstein (ed. H. P. Taylor,Jr. et al). Geochem. Soc. Spec. Pub. 3, 35–51.Taylor H. P., Jr. and Epstein S. (1962) Relationship between 18 O/ 16 Oratios in coexisting minerals in igneous and metamorphic rocks, Part1. Bull. Geol. Soc. Amer. 73, 461–480.Truesdell A. H. (1974) Oxygen isotope activities and concentration inaqueous salt solutions at elevated temperatures: Consequences forisotope geochemistry. Earth Plan. Sci. Lett. 23, 387–396.Zheng Y. (1991) Calculation of oxygen isotope fractionation in metaloxides. Geochim. Cosmochim. Acta. 55, 2299–2307.Zheng Y. (1995) Isotope fractionation in uranium oxides. Nucl. Sci.Tech. 6, 193–197.