Oxygen Isotope Fractionation Factors Research Articles

Oxygen and hydrogen isotope compositions of Permian pedogenic phyllosilicates: Development of modern surface domain arrays and implications for paleotemperature reconstructions

Mineralogic, chemical, and oxygen and hydrogen isotope compositions of 15 different phyllosilicate samples from Permo–Pennsylvanian-age paleosols of the eastern shelf of the Midland basin of Texas and the southern Anadarko basin are presented. Mixtures of 2 : 1 phyllosilicates and kaolinite dominate most samples, although some samples consist of relatively pure 2 : 1 phyllosilicates. Chemical and mineralogic data are used in conjunction with published thermodynamic data to calculate hydrogen and oxygen isotope fractionation factors for each sample. In turn, application of measured oxygen and hydrogen isotope compositions of the phyllosilicates to temperature-dependent fractionation equations are used to calculate paleotemperatures of crystallization. The δD values of the phyllosilicates range from − 69‰ to − 55‰. The δ 18O values range from 19.5‰ to 22.7‰. If these samples preserve a record of equilibrium with paleo-meteoric waters, the isotopic compositions of the phyllosilicates correspond to paleotemperatures of phyllosilicate crystallization ranging from 22 ± 3 °C to 35 ± 3 °C. In particular, the stratigraphic trend of calculated temperatures from Midland basin samples suggests that Early Permian surface temperatures may have been up to 10 °C warmer than those of the latest Pennsylvanian.

Oxygen isotope fractionation factors involving cassiterite (SnO 2): I. calculation of reduced partition function ratios from heat capacity and X-ray resonant studies

Oxygen isotope equilibrium fractionation constants (β 18O-factors) of cassiterite were evaluated on the basis of heat capacity and X-ray resonant (Mössbauer spectroscopy and X-ray inelastic scattering) data. The low-temperature heat capacity of cassiterite was measured in the range from 13 to 340 K using an adiabatic calorimeter. Results of measurements of two samples agree very closely but deviate more than 5% from previous heat capacity data used for calculation of thermodynamic functions. The temperature dependence of heat capacity was treated using the modern version of the Thirring expansion, and the appropriate temperature dependence of the vibrational kinetic energy was found. Measurements of temperature-dependent Mössbauer parameters of cassiterite were conducted in the range from 300 to 900 K. The attempt to describe Mössbauer fraction and the second order Doppler (SOD) shift on the basis of the Debye model failed. The first term of the Thirring expansion of the Mössbauer SOD shift agrees with that calculated from the Sn sublattice vibration density of states (VDOS) obtained via synchrotron X-ray scattering. Based on this agreement we calculated the kinetic energy of the cassiterite Sn sublattice from VDOS. From the kinetic energy of the total cassiterite crystalline lattice and its Sn sublattice, β 18O-factors of cassiterite were computed in the temperature range 300–1500 K by the method of Polyakov and Mineev (2000). Appropriate polynomials, which are valid at temperatures above 400 K, are the following: 1 0 3 ln ⁡ β SnO 2 = ( 7.176 ± 0.252 ) x − ( 0.07369 ± 0.00089 ) x 2 + ( 0.0008026 ± 0.0001022 ) x 3 , x = 10 6 / T 2 ; 10 3 ln ⁡ α CaCO 3 − S nO 2 = 4.607 x − 0.3463 x 2 + 0.01500 x 3 , x = 10 6 / T 2 ; 10 3 ln ⁡ α SiO 2 − S nO 2 = 4.942 x − 0.2963 x 2 + 0.01150 x 3 , x = 10 6 / T 2 . The evaluated cassiterite isotope fractionation factors are significantly different from those obtained by synthesis, increment and empirical methods. To resolve the differences, laboratory direct exchange experiments are needed.

Oxygen isotope fractionation factors involving cassiterite (SnO 2): II. determination by direct isotope exchange between cassiterite and calcite

Direct oxygen isotope fractionation between cassiterite and calcite has been investigated experimentally at 15 kbar with temperature ranging from 800 to 1000°C. Combined with the quartz-calcite fractionation measured with the same technique ( Clayton et al., 1989), the calcite-cassiterite and quartz-cassiterite oxygen isotope fractionations can be expressed as: Δ CaCO 3 − SnO 2 = 4.14 × 10 6 T − 2 Δ SiO 2 − SnO 2 = 4.52 × 10 6 T − 2 at temperatures higher than 600°C. These calibrations are in good agreement with those obtained from heat capacity and X-ray resonant data. The theory of using X-ray resonant data to calculate the reduced partition function ratios of transition elements that have at least one Mössbauer-sensitive isotope is evaluated with the current experimental result.

Calibration of the δ18O paleothermometer for dolomite precipitated in microbial cultures and natural environments

Decades of various and numerous isotopic studies to interpret the environmental conditions of dolomite formation proved to be inconclusive because the temperature-dependent oxygen isotope fractionation factor between dolomite and the solution from which it precipitated could not be determined experimentally at low temperatures. With the discovery of bacteria that mediate the precipitation of dolomite, it is now possible to overcome kinetic barriers and precipitate dolomite under controlled temperature conditions in culture experiments. Herein we report on the results of microbial experiments that have enabled us to calibrate the dolomite-water oxygen isotope fractionation factor and provide a paleothermometer to evaluate conditions of ancient dolomite formation. The temperature (T) dependence of the fractionation is defined by the equation: 1000 In α d o l o m i t e - w a t e r = 2.73 X 10 6 T - 2 + 0.26.

Mineralogical and isotopic properties of inorganic nanocrystalline magnetites

Inorganic magnetite nanocrystals were synthesized in an aqueous medium at 25°C, atmospheric pressure, ionic strength of 0.1 M, oxygen fugacity close to 0, and under controlled chemical affinity, which was maintained constant during an experiment and varied between different experiments. The total concentration of iron in the initial solutions, with Fe(III)/Fe(II) ratios of 2, was varied in order to measure the role of this parameter on the reaction rate, particle morphology, and oxygen isotopic composition. The reaction rates were followed by a pHstat apparatus. The nature and morphology of particles were studied by transmission electron microscopy and electron energy loss spectroscopy. Fractionation factors of oxygen isotopes were determined by mass spectrometry after oxygen extraction from the solid on BrF 5 lines. At low total iron concentrations, goethite and poorly crystalline iron oxides were observed coexisting with magnetite. At higher concentrations, euhedral single crystals of pure magnetite with an average characteristic size of 10 nm were formed, based on a first-order rate law with respect to total iron concentration. These results confirm that, under high supersaturation conditions, low-temperature inorganic processes can lead to the formation of well-crystallized nanometric magnetite crystals with narrow size distribution. The observed oxygen isotope fractionation factor between magnetite crystals and water was of 0–1‰, similar to the fractionation factor associated with bacterially produced magnetite. We suggest that the solution chemistry used in this study for inorganic precipitation is relevant to better understanding of magnetite precipitation in bacterial magnetosomes, which might thus be characterized by high saturation states and pH.

Oxygen isotope fractionation in synthetic magnesian calcite

Mg-bearing calcite was precipitated at 25°C in closed system free-drift experiments from solutions containing NaHCO 3, CaCl 2 and MgCl 2. The chemical and isotope composition of the solution and precipitate were investigated during time course experiments of 24-h duration. Monohydrocalcite and calcite precipitated early in the experiments (<8 h), while Mg-calcite was the predominant precipitate (>95%) thereafter. Solid collected at the end of the experiments displayed compositional zoning from pure calcite in crystal cores to up to 23 mol% MgCO 3 in the rims. Smaller excursions in Mg were superimposed on this chemical record, which is characteristic of oscillatory zoning observed in synthetic and natural solid-solution carbonates of differing solubility. Magnesium also altered the predominant morphology of crystals over time from the {104} to {100} and {110} growth forms. The oxygen isotope fractionation factor for the magnesian-calcite-water system (as 10 3lnα Mg-cl-H 2O ) displayed a strong dependence on the mol% MgCO 3 in the solid phase, but quantification of the relationship was difficult due to the heterogeneous nature of the precipitate. Considering only the Mg-content and δ 18O values for the bulk solid, 10 3lnα Mg-cl-H 2O increased at a rate of 0.17 ± 0.02 per mol% MgCO 3; this value is a factor of three higher than the single previous estimate (Tarutani T., Clayton R.N., and Mayeda T. K. (1969) The effect of polymorphims and magnesium substitution on oxygen isotope fractionation between calcium carbonate and water. Geochim. Cosmochim. Acta 33, 987–996). Nevertheless, extrapolation of our relationship to the pure calcite end member yielded a value of 27.9 ± 0.02, which is similar in magnitude to published values for the calcite-water system. Although no kinetic effect was observed on 10 3lnα Mg-cl-H 2O for precipitation rates that ranged from 10 3.21 to 10 4.60 μmol · m −2 · h −1, it was impossible to disentangle the potential effect(s) of precipitation rate and Mg-content on 10 3lnα Mg-cl-H 2O due to the heterogeneous nature of the solid. The results of this study suggest that paleotemperatures inferred from the δ 18O values of high magnesian calcite (>10 mol% MgCO 3) may be significantly underestimated. Also, the results underscore the need for additional experiments to accurately characterize the effect of Mg coprecipitation on the isotope systematics of calcite from a chemically homogeneous precipitate or a heterogeneous material that is analyzed at the scale of chemical and isotopic zonation.

04/01107 Numerical modelling of pipeline puncture under excavator loading. Part II: parametric study: Brooker, D. C. International Journal of Pressure Vessels and Piping, 2003, 80, (10), 727–735

New experimental results are reported on oxygen isotope fractionation factors, αT, between the δ18O compositions of carbon dioxide liberated by phosphoric acid in the temperature interval of 323 to 373K and that of total oxygen from a natural magnesite (MgCO3). These results are distinctly different from some previously published mutually inconsistent data, and can be expressed as a linear relationship: 103 lnαT = [{(6.845 ± 0.475)∗105}/T2] + (4.22 ± 0.08), where 103 lnαT refers to fractionation at different temperatures T in Kelvin.Fractionation factors have also been determined at 323 and 368K on a natural calcite. The results on calcite are in excellent agreement with previously published data and can be written as:103 lnαT = [{(5.608 ± 0.151)∗105}/T2] + (3.89 ± 0.08).The combined results on magnesite and calcite yield a computed value of α = 1.01117 for dolomite at 298K, assuming equal proportion of 0.5 mole of magnesium and calcium in dolomite, the previously reported experimental values being 1.01109 and 1.01110.

A model for 18O/ 16O variations in CO 2 evolved from goethite during the solid-state α-FeOOH to α-Fe 2O 3 phase transition

“Plateau” δ 18O values of CO 2 that evolved from the Fe(CO 3)OH component during isothermal vacuum dehydrations (200–230 °C) of 18 natural goethites range from 8.2 to 28.1‰. In contrast, the measured δ 18O values of the goethite structural oxygen range from −11.3 to 1.7‰. The results of this study indicate that the apparent oxygen isotope fractionation factor ( 18α app) between plateau CO 2 and initial goethite is systematically related to the rate of isothermal vacuum dehydration. The nonlinear correlation and the magnitudes of the 18α app values are predicted by a relatively simple mass balance model with the following assumptions: (1) the rate of isothermal vacuum dehydration of goethite (for the interval from 0 to ∼60 to 80% loss of structural hydroxyl hydrogen) can be reasonably well represented by first-order kinetics and (2) isotopic exchange between evolving H 2O vapor and solid occurs only in successive, local transition states. The generally good correspondence between the model predictions and the experimental data seems to validate these assumptions. Thus, the 18O/ 16O ratios of the evolved CO 2 can act as probes into the transient processes operating at the molecular level during the solid-state goethite-to-hematite phase transition. For example, the activation energy for the rate constant associated with the transition state, oxygen isotopic exchange between solid and H 2O vapor, is tentatively estimated as 28 ± 11 KJ/mol. Such knowledge may be of consequence in understanding the significance of 18O/ 16O ratios in hematites from some natural environments (e.g., Mars?). Kinetic data and δ 18O values of CO 2 are routinely obtained in the course of measurements of the abundance and δ 13C values of the Fe(CO 3)OH in goethite. The observed correlation between 18α app and dehydration rates suggests that plateau δ 18O values of evolved CO 2 may provide complementary estimates of the δ 18O values of total goethite structural oxygen (O, OH, CO 2) with an overall precision of about ±1‰. However, because of isotopic exchange during the dehydration process, δ 18O values of the evolved CO 2 do not reflect the original δ 18O values of the CO 2 that was occluded as Fe(CO 3)OH in goethite.

Calculation of oxygen isotope fractionation in magmatic rocks

The increment method is applied to calculation of oxygen isotope fractionation factors for common magmatic rocks. The 18O-enrichment degree of the different compositions of magmatic rocks is evaluated by the oxygen isotope indices of both CIPW normative minerals and normalized chemical composition. The consistent results are obtained from the two approaches, pointing to negligible oxygen isotope fractionation between rock and melt of the same compositions. The present calculations verify the following sequence of 18O-enrichment in the magmatic rocks: felsic rocks>intermediate rocks>mafic rocks>ultramafic rocks. Two sets of internally consistent fractionation factors are acquired for phenocryst–lava systems at the temperatures above 1000 K and rock–water systems in the temperatures range of 0–1200 °C, respectively. The present calculations are consistent with existing data from experiments and/or empirical calibrations. The obtained results can be used to quantitatively determine the history of water–rock interaction and to serve geological thermometry for various types of magmatic rocks (especially extrusive rocks).

Microbial respiration and diffusive transport of O 2, 16O 2, and 18O 16O in unsaturated soils: a mesocosm experiment

Although the flux of molecular O 2 between the atmosphere and the subsurface is intrinsically linked to the net soil production of greenhouse gasses, few studies have focused on the controls affecting the isotopic composition of O 2 in the subsurface. Here, we developed and tested a stable oxygen isotope tracer technique and gas transport modeling approach to evaluate O 2 cycling and fluxes from the subsurface that used an environmentally controlled soil mescosm. We measured the O 2 and δ 18O 2 profiles in a model unsaturated soil zone and quantified the O 2 consumption rates and the O 2 isotope fractionation factors resulting from the combined processes of subsurface microbial (including bacteria, fungi, and protozoa) consumption of O 2 and diffusive influx of O 2 from the atmosphere. We found that at high respiration rates in the mesocosm, there appeared to be very little isotope fractionation of O 2 by soil microorganisms. Although the mesocosm respiration rates are not typical of natural soils in northern temperate climes, they may be more representative of soils in warm and moist tropical environments. Our findings caution against the indiscriminate application of laboratory-determined oxygen isotope fractionation factors to field settings. The oxygen isotope tracer and modeling approach demonstrated here may be applied to gain a better understanding of biogenic gas production and O 2 cycling in subsurface systems and soils.

Temperature dependence of oxygen isotope fractionation of CO2 from magnesite-phosphoric acid reaction

New experimental results are reported on oxygen isotope fractionation factors, αT, between the δ18O compositions of carbon dioxide liberated by phosphoric acid in the temperature interval of 323 to 373K and that of total oxygen from a natural magnesite (MgCO3). These results are distinctly different from some previously published mutually inconsistent data, and can be expressed as a linear relationship: 103 lnαT = [{(6.845 ± 0.475)∗105}/T2] + (4.22 ± 0.08), where 103 lnαT refers to fractionation at different temperatures T in Kelvin.Fractionation factors have also been determined at 323 and 368K on a natural calcite. The results on calcite are in excellent agreement with previously published data and can be written as:103 lnαT = [{(5.608 ± 0.151)∗105}/T2] + (3.89 ± 0.08).The combined results on magnesite and calcite yield a computed value of α = 1.01117 for dolomite at 298K, assuming equal proportion of 0.5 mole of magnesium and calcium in dolomite, the previously reported experimental values being 1.01109 and 1.01110.

Chemical, mineralogical and isotope behavior, and phase transformation during the precipitation of calcium carbonate minerals from intermediate ionic solution at 25°C

A set of time-series experiments was performed in which CaCO3 was precipitated (25°C, 1 atm) from aliquots of an NaHCO3 (0.330 M)/CaCl2 (0.0023 M) solution in a closed system to evaluate the effect of phase transformation on calcium carbonate precipitation kinetics and isotope behavior.Monohydrocalcite precipitated initially at a solution saturation state (Ωmhc) slightly greater than one, whereas calcite crystallized later at a significant higher saturation state (Ωcl ≅ 14.4). The precipitation of calcite promoted the dissolution of monohydrocalcite at a rate that exceeded calcite precipitation, producing anomalous behavior in the chemical and isotope composition of the system.The carbon isotope fractionation factors (103 lnα) for monohydrocalcite-HCO3(aq)- and monohydrocalcite-CO2(g) were 0.36 ± 0.01‰ and 8.35 ± 0.01‰, respectively. The oxygen isotope fractionation factor for monohydrocalcite-H2O was 27.8 ± 0.1‰.The carbon isotope fractionation factors for calcite-HCO3 (aq)− and calcite-CO2(g) were 0.94 ± 0.06‰ and 8.93 ± 0.06‰, respectively, whereas the oxygen isotope fractionation factor for calcite-H2O(l) was 28.0 ± 0.2‰. A carbon isotope fractionation factor of 0.58 ± 0.07‰ was determined for the mineral pair calcite-monohydrocalcite, but no fractionation was observed for oxygen isotopes over time steps when both minerals co-precipitated.Fractionation factors for calcite were independent of precipitation rate over the range in rates of 103.96 to 105.63 μmol/m2h. These results extend the upper limit of characterization for the relationship between precipitation rate and isotope partitioning of carbon between calcite, HCO3 (aq)− and CO2(g), and quantitatively document for the first time the independence between precipitation rate and oxygen isotope partitioning in the calcite-H2O(l) system.

Empirical determination of oxygen isotope fractionation factors for titanite with respect to zircon and quartz

Measurements of oxygen isotope fractionations between coexisting igneous titanite (Ttn) and zircon (Zrc) have been used to formulate a self-consistent, empirical calibration of equilibrium oxygen isotope partitioning: 1000 lnα(Zrc−Ttn)=1.02×106/(T2) This calibration is based on the average measured Δ(Zrc-Igneous Ttn) = 1.2 ± 0.3‰ (n = 27 rocks) and a closure temperature of titanite to oxygen diffusion of approximately 650°C. The average measured fractionation between zircon and metamorphic titanite is 2.1 ± 0.4‰ (n = 5 rocks). These results show that Δ(Zrc-Ttn) can be used to distinguish igneous vs. metamorphic (or hydrothermal) titanite.The new zircon-titanite fractionation has been combined with published experimental and empirical data to provide a new quartz-titanite fractionation curve: 1000 lnα(Qtz−Ttn)=3.57×106/(T2) This new calibration is consistent with Δ(Qtz-Igneous Ttn) data from coexisting mineral pairs in granitic rocks. Modeling diffusional exchange of oxygen during cooling in a typical granitic rock, using the new calibration, results in the same Δ(Qtz-Igneous Ttn) as measured (4.5 ± 0.43‰, n = 25 rocks). Both measured Δ(Zrc-Igneous Ttn) and Δ(Qtz-Igneous Ttn) differ from fractionations calculated using the semiempirical increment method. The consistency of empirical Δ(Zrc-Igneous Ttn) in a range of geologic environments allows the calculation of a more accurate fractionation factor for titanite.

A low temperature oxygen isotope thermometer for cerussite, with applications at Broken Hill, New South Wales, Australia

Slow precipitation experiments have been used to determine the oxygen isotope fractionation between cerussite (PbCO3) and water over the temperature range 20 to 65°C. The temperature dependent fractionation for oxygen can be expressed as: 1000 ln α(cerussiteCO3-water)=2.63 (106/T2)−3.58 Independent determinations of the oxygen isotope fractionation factor for recent historical cerussite samples from the Bwlch-Glas Mine, Wales, UK, and the Pinnacles Mine, Broken Hill, NSW, Australia, plot close to our experimentally determined fractionation curve. Similarly, earlier high temperature cerussite experiments are in agreement with our thermometer. This new oxygen isotope thermometer was used to examine data for 7 cerussite samples from the Block 14 and 3 from the Kintore open cuts on the main Broken Hill lode, NSW, Australia. Broken Hill cerussite formed at temperatures significantly greater than modern ambient air temperatures. This is consistent with malachite thermometry results at Broken Hill. It is suggested that these higher apparent temperatures result from the exothermic oxidation of concentrated sulphides. Results indicate that oxidation of sulphides in the Block 14 Mine produced shallow subsurface temperatures up to at least 50°C. Carbon isotope data suggest incorporation of significant amounts of organic carbon, possibly from a subsurface bacterial carbon source associated with sulphide oxidation.

In situ Stable Isotopic Evidence for Protracted and Complex Carbonate Cementation in a Petroleum Reservoir, North Coles Levee, San Joaquin Basin, California, U.S.A

ABSTRACT Knowledge of the evolution of carbonate cementation in hydrocarbon reservoirs is key to understanding the history of fluid flow during petroleum accumulation. The Stevens sands is a sequence of marine shales and deep-sea fan sands that was deposited within the Miocene Monterey Formation in the south-central part of the San Joaquin basin, California, during the upper Miocene (10-6 Ma). Rapid, high-precision in situ oxygen and carbon isotopic analyses of carbonate phases using the ion microprobe operated in multi-collection mode, in conjunction with electron microprobe analyses, indicate that carbonate cement zones within the Stevens sands at North Coles Levee (NCL) have had a complex and protracted fluid history. Three main generations of carbonate cement were identified. The relative timing of carbonate cement precipitation within the Stevens sands at NCL was estimated using the thermal and burial history of the San Joaquin basin, in situ oxygen isotope data, and cementation temperatures derived from equilibrium oxygen isotope fractionation factors for calcite-water and dolomite-water. Precipitation of these cement zones began soon after sediment deposition ( 7 Ma) and is ongoing. Early dolomite was precipitated at a temperature of 10°C, near the sediment-water interface, and soon after sediment deposition. Calcite cements, which are the most abundant variety, precipitated semicontinuously between 4 Ma and 5 Ma, at temperatures between 50°C and 65°C, and depths of 800 m to 1300 m. Fe-dolomite, which is paragenetically late, appears to have precipitated at temperatures near 100°C in response to pore-pressure reduction, which accompanied exploitation of the gas cap within the last 35 years. Carbon in these cements was likely derived from several sources including marine, maturing hydrocarbons, and a zone of methanogenesis.

Oxygen isotopes of marine diatoms and relations to opal-A maturation

In order to develop the potential tool of diatom oxygen isotopes for paleoenvironmental studies we compared oxygen isotopes of natural marine diatoms sampled from ocean surface water, sediment traps and surface sediments with oxygen isotopic fractionations determined for laboratory diatom cultures. Freshly grown natural diatoms (phytoplankton samples and sediment trap material) and cultured diatoms reveal similar oxygen isotope fractionation factors. The fresh diatoms have 3 to 10‰ lower isotope fractionation factors than fossil (sedimentary) diatoms. A temperature-related oxygen isotope fractionation could not be established for the laboratory cultures (and the natural phytoplankton samples), and there is evidence that diatom growth rate until reaching the stationary growth state also controls the measured silica-water oxygen isotope fractionation factor. It is possible, however, that slow diatom growth in sea surface water may well lead to a temperature-dependent silica-water oxygen isotope fractionation which is the prerequisite for a use of diatom oxygen isotopes in paleo-surface water studies. FTIR-spectroscopic analyses of various diatomaceous materials revealed that the ratio of integrated peak intensities for Si-O-Si/Si-OH correlates with the 3 to 10‰ δ 18O silica increase from fresh to fossil diatoms. Open-system (flow-through) silica dissolution experiments suggest that the diatom frustules are isotopically homogenous and that the increase in 18O is therefore not due to dissolution of isotopically light surficial Si-OH groups. It is concluded that slow internal condensation reactions during silica maturation in surface sediments cause both an increase in the intensity ratio of Si-O-Si/Si-OH and the 18O content of framework oxygen. These findings also indicate that the oxygen isotope compositions of marine sediment diatoms do not indicate sea surface water temperature but rather reflect variable 18O contents of surface sediments.

Equilibrium Oxygen, Hydrogen and Carbon Isotope Fractionation Factors Applicable to Geologic Systems

As demonstrated by the chapters in this short course, stable isotope techniques are an important tool in almost every branch of the earth sciences. Central to many of these applications is a quantitative understanding of equilibrium isotope partitioning between substances. Indeed, it was Harold Urey’s (1947) thermodynamically based estimate of the temperature-dependence of 18O/16O fractionation between calcium carbonate and water, and a recognition of how this information might be used to determine the temperatures of ancient oceans, that launched the science of stable isotope geochemistry. The approach pioneered by Urey has since been used to estimate temperatures for a wide range of geological processes (e.g. Emiliani 1955; Anderson et al. 1971; Clayton 1986; Valley, this volume). In addition to their geothermometric applications, equilibrium fractionation data are also important in the study of fluid-rock interactions, including those associated with diagenetic, hydrothermal, and metamorphic processes (Baumgartner and Valley, this volume; Shanks, this volume). Finally, a knowledge of equilibrium fractionation is a necessary first step in evaluating isotopic disequilibrium, a widespread phenomenon that is increasingly being used to study temporal relationships in geological systems (Cole and Chakraborty, this volume). In the fifty-four years since the publication of Urey’s paper, equilibrium fractionation data have been reported for many minerals and fluids of geological interest. These data were derived from: (1) theoretical calculations following the methods developed by Urey (1947) and Bigeleisen and Mayer (1947); (2) direct laboratory experiments; (3) semi-empirical bond-strength models; and (4) measurement of fractionations in natural samples. Each of these methods has its advantages and disadvantages. However, the availability of a variety of methods for calibrating fractionation factors has led to a plethora of calibrations, not all of which are in agreement. In this chapter, we evaluate the major methods for determining fractionation factors. …

Stable Isotope Thermometry at High Temperatures

Determination of accurate temperatures for geological events has been the grail of stable isotope geochemistry since the seminal 1947 paper by Urey. In theory, this should be simple. Calibrated mineral pairs are common, analysis is rapid, and there is no significant pressure correction. However, in spite of widespread application, the promise of reliable thermometry has been elusive. Stable isotope temperatures in metamorphic and igneous rocks are often controversial; is the fractionation between two phases a thermometer, speedometer (rate dependant), hygrometer (P(H2O)-dependant), or chimera? A number of factors have contributed to this uncertainty including: incomplete and sometimes conflicting calibration of isotope fractionation factors; limited understanding of the kinetics of mineral exchange; and the lack of microanalytical techniques. This situation has improved dramatically over the past decade, permitting reliable and detailed thermal histories to be inferred. In a landmark paper, Bottinga and Javoy (1975) empirically calibrated the oxygen isotope fractionation factors among ten rock-forming minerals at T > 500°C. Their compilation shows that 84% of analyzed samples contain at least three minerals meeting a concordance test for equilibrium. They concluded “the great majority of igneous and metamorphic rocks …. has conserved a state of oxygen isotope exchange equilibrium.” If correct, this bodes well for thermometry. However, this conclusion was contested by Deines (1977) who critically examined the existing data and concluded that “less than half of the rocks analyzed to date would yield concordant 18O-derived temperatures.” While seeming irreconcilable, both conclusions are based on fact. All rocks contain some degree of disequilibrium, but many also preserve a thermometric record. It is the purpose of this review to discuss the new strategies for reliable isotope thermometry and to explore disequilibrium processes so they may be recognized and properly interpreted. The organization of this paper will be to discuss the assumptions …

Low temperature oxygen isotopic fractionation in the uraninite–UO 3–CO 2–H 2O system

Oxygen-isotope fractionation factors for uraninite–water and UO 3–water were determined from a series of uraninite–CO 2 and UO 3–CO 2 exchange experiments between 100 and 300°C. The measured fractionation factors are similar to the theoretical fractionation factors of Hattori and Halas (1982) at high temperatures and Zheng (1991) at low temperatures and are nonlinear over the temperature range examined. Regression through these curve gives 1000 ln α UO 2–H 2 O = 16.58 (10 6/ T 2) − 77.52 (10 3/ T) + 77.48, 1000 ln α UO 3–H 2 O = −2.21 (10 6/ T 2) + 25.06 (10 3/ T) − 45.90 with 2σ errors of ±2.5‰. Diffusion coefficients of oxygen with temperature for uraninite and UO 3 can be represented by D (uraninite) = 3.1 × 10 −10 exp (−63981/ RT) and D (UO 3) = 1.3 × 10 −15 exp (−21,197/ RT), where the activation energies for uraninite and UO 3 are 64.0 ± 0.2 kcal/mole and 21.2 ± 0.2 kcal/mol, respectively. Exceptionally low δ 18O values of natural uraninites (i.e., −32‰ to −19.5‰) from unconformity-type uranium deposits in Saskatchewan, in conjunction with theoretical and experimental uraninite–water and UO 3–water fractionation factors, suggest that primary uranium mineralization is not in oxygen isotopic equilibrium with coeval clay and silicate minerals. The low δ 18O values have been interpreted as having resulted from the low temperature recrystallization of primary uranium mineralization in the presence of relatively modern meteoric fluids having δ 18O values of ca. −18‰, despite petrographic and U–Pb isotope data that indicate limited alteration. Therefore, it appears that uranium minerals can exchange oxygen isotopes with later fluids, with only limited actinide remobilization and minor disturbances to their original chemical compositions and textures.

Oxygen-Isotope Fractionation between Aluminum-Hydroxide Phases and Water at &lt;60°C: Results of Decade-Long Synthesis Experiments

AbstractOxygen-isotope data were obtained for synthetic aluminum-hydroxide phases precipitated over 65–125 mo and have been compared to results from similar experiments conducted for 3–56 mo. The Al(OH)3 polymorphs, gibbsite, nordstrandite, and bayerite, were synthesized, but gibbsite was dominant in most samples, and commonly the only phase present. Using pure gibbsite samples, the following oxygen-isotope fractionation factors, αgibbsite−H2O, were obtained: 1.0167 ± 0.0003 (9 ± 1°C), 1.0147 ± 0.0007 (24 ± 2°C), 1.0120 ± 0.0003 (51 ± 2°C). These values, and the associated equation for an oxygen-isotope geothermometer for the interval 0–60°C 103ln αgibbsite−H2O=2.04×106/T2−3.61×103/T+3.65 (T in K), are not significantly different from those obtained from experiments of much shorter duration. These results, and the good agreement with αgibbsite−H2O values obtained for well-constrained natural systems, suggest that the experimentally determined fractionation factors describe equilibrium conditions for gibbsite that has precipitated directly from solution.As also proposed by others using a modified-increment calculation, our synthesis experiments suggest that αAl(OH)3−H2O is polymorph-dependent at low temperatures and that a significant temperature-dependent trend exists in the values of αAl(OH)3−H2O. However, previously calculated fractionation factors obtained using the modified-increment method are higher than those obtained from the experiments, with this discrepancy becoming larger as temperature decreases.