Oxygen Isotope Fractionation Factors Research Articles

First application of mass spectrometry and gas chromatography in investigation of α‐cellulose hydrolysates: the influence of climate changes on glucose molecules in pine tree‐rings

We present the first results of the quantitative and qualitative gas chromatographic and isotope ratio mass spectrometric analysis of monosaccharides derived from acid hydrolysis of α-cellulose extracted from annual pine tree-rings. The conifers investigated in this study grew in the Niepolomice Forest in Poland, and the annual rings covered the time span from 1940 to 2000 AD. The main components of the α-cellulose samples were two saccharides: glucose and mannose. The amount of glucose in the annual rings varied between 17 and 44%. The δ(13)C of glucose was found to be less negative than that of α-cellulose and the δ(18)O values in glucose were less positive than those in α-cellulose. The content of monosaccharides in the α-cellulose samples has an influence on the isotope fractionation factors. The values of the carbon isotope fractionation factor increase with an increase in the monosaccharides concentration in α-cellulose, while the values of the oxygen isotope fractionation factor decrease with an increase in monosaccharides concentration in α-cellulose. The challenge is to establish, with respect to climate changes and environmental conditions, the significance of the interannual variations in the observed monosaccharide concentration.

On the theoretical calculations of oxygen isotope fractionation factors for carbonate-water systems

On the theoretical calculations of oxygen isotope fractionation factors for carbonate-water systems

Hydrothermal recrystallization of the Lower Ordovician dolomite and its significance to reservoir in northern Tarim Basin

Discovered in S15 and some other wells, the Lower Ordovician in the northern Tarim Basin consists mainly of brown gray-dark gray very fine-fine crystalline dolomite, with a minor portion of locally light gray-white medium-coarse crystalline dolomite. Silicification can be observed in the medium-coarse dolomite, and some euhedral drusy quartz can also be found in pores and fractures of the dolomite. The homogenization temperature of the fluid inclusions in the medium-coarse dolomite is between 110 and 200°C with maximum between 140 and 190°C, and the salinity is between 10.7 and 18.5 wt.% NaCl Eq. The homogenization temperature and salinity of the fluid inclusions in the medium-coarse dolomite are similar to those in the drusy quartz. Compared with the very fine-fine dolomite, the medium-coarse phase contains relatively high Fe and Mn. The average concentration of FeO and MnO in the medium-coarse dolomite is 1.917% and 0.323%, respectively. The medium-coarse dolomite has a remarkable negative Eu anomaly, consistent with the REE pattern of the intermediate-felsic igneous rocks in the Tarim Basin. The oxygen isotopic composition of the medium-coarse dolomite is relatively lighter than that of the very fine-fine dolomite. The δ18OPDB values of the medium-coarse dolomite are between −10.35‰ and −7.31‰. The δ18OSMOW values of the fluid associated with the medium-coarse dolomite can be calculated according to homogenization temperature and oxygen isotope fractionation factor between dolomite and fluid, and the calculated values are between +4‰ and +10‰, consistent with those of the hydrothermal fluid. The medium-coarse dolomite has relatively high 87Sr/86Sr ratios as well, indicating an origin associated with intermediate-felsic igneous rock. The homogenization temperature, element composition, REE pattern, oxygen and strontium isotopes demonstrate that the medium-coarse dolomite is the result of recrystallization of very fine-fine dolomite under hydrothermal environment. The hydrothermal dolomite recrystallization is a special event but exists extensively in the Tarim Basin. The recrystallized dolomite becomes well reservoir bed for many intercrystalline pores and dissolution pores are produced in this event, so that more attention should be paid to the altered dolomite during the petroleum exploration in the lower Paleozoic in the Tarim Basin.

Oxygen isotope fractionation of dissolved oxygen during abiological reduction by aqueous sulfide

The stable oxygen isotope fractionation factor of dissolved oxygen gas has been measured during abiological reduction by aqueous sulfide at 10–40 °C under circum-neutral conditions in order to understand the stable isotope behavior of oxygen during interaction with sulfide-rich natural waters. The oxygen isotopic composition of O 2 becomes isotopically heavier as the reaction proceeds, with the oxygen isotope enrichment factor ε ranging from − 8.2 to − 9.8‰ at 10 °C, from − 8.8 to − 10.9‰ at 23 °C, and from − 6.0 to − 10.7‰ at 40 °C. Values of ε show little dependence on the initial sulfide concentration, reaction temperature, or salinity in the solutions. In addition to the direct reaction of sulfide with O 2, aqueous sulfide oxidation can also progress through complex and multiple reaction pathways where O 2 can further react with intermediate products of sulfide oxidation such as S 0, S x 2−, S 2O 3 2−, and SO 3 2−. Hence, the measured values of ε are a result of multiple isotope fractionations during O 2 consumption by these sulfur species, where the net measured isotope fractionation factor is dependent upon the relative magnitude of each individual O 2-consuming reaction, and the isotope fractionation factor associated with each of these reactions. Nevertheless, the results will have application to the identification and possible quantification of O 2 consumption by reaction with sulfide in natural environments.

Oxygen isotopic fractionation during inorganic calcite precipitation ― Effects of temperature, precipitation rate and pH

Stable oxygen isotopic fractionation during inorganic calcite precipitation was experimentally studied by spontaneous precipitation at various pH (8.3 < pH < 10.5), precipitation rates (1.8 < log R < 4.4 μmol m − 2 h − 1 ) and temperatures (5, 25, and 40 °C) using the CO 2 diffusion technique. The results show that the apparent stable oxygen isotopic fractionation factor between calcite and water ( α calcite–water) is affected by temperature, the pH of the solution, and the precipitation rate of calcite. Isotopic equilibrium is not maintained during spontaneous precipitation from the solution. Under isotopic non-equilibrium conditions, at a constant temperature and precipitation rate, apparent 1000ln α calcite–water decreases with increasing pH of the solution. If the temperature and pH are held constant, apparent 1000ln α calcite–water values decrease with elevated precipitation rates of calcite. At pH = 8.3, oxygen isotopic fractionation between inorganically precipitated calcite and water as a function of the precipitation rate ( R) can be described by the expressions 1000 ln α calcite – water = - 1.102 log R + 34.56 1000 ln α calcite – water = - 1.094 log R + 30.87 1000 ln α calcite – water = - 0.534 log R + 26.80 at 5, 25, and 40 °C, respectively. The impact of precipitation rate on 1000ln α calcite–water value in our experiments clearly indicates a kinetic effect on oxygen isotopic fractionation during calcite precipitation from aqueous solution, even if calcite precipitated slowly from aqueous solution at the given temperature range. Our results support Coplen's work [Coplen T. B. (2007) Calibration of the calcite–water oxygen isotope geothermometer at Devils Hole, Nevada, a natural laboratory. Geochim. Cosmochim. Acta 71, 3948–3957], which indicates that the equilibrium oxygen isotopic fractionation factor might be greater than the commonly accepted value.

Oxygen and hydrogen isotope fractionation in serpentine–water and talc–water systems from 250 to 450 °C, 50 MPa

Oxygen and hydrogen isotope fractionation factors in the talc–water and serpentine–water systems have been determined by laboratory experiment from 250 to 450 °C at 50 MPa using the partial exchange technique. Talc was synthesized from brucite + quartz, resulting in nearly 100% exchange during reaction at 350 and 450 °C. For serpentine, D–H exchange was much more rapid than 18O– 16O exchange when natural chrysotile fibers were employed in the initial charge. In experiments with lizardite as the starting charge, recrystallization to chrysotile enhanced the rate of 18O– 16O exchange with the coexisting aqueous phase. Oxygen isotope fractionation factors in both the talc–water and serpentine–water systems decrease with increasing temperature and can be described from 250 to 450 °C by the relationships: 1000 ln α talc – water 18 O – 16 O = 11.70 × 10 6/T 2 − 25.49 × 10 3/T + 12.48 and 1000 ln α serpentine – water 18 O – 16 O = 3.49 × 10 6/T 2 − 9.48 where T is temperature in Kelvin. Over the same temperature interval at 50 MPa, talc–water D–H fractionation is only weakly dependent on temperature, similar to brucite and chlorite, and can be described by the equation: 1000 ln α talc – water D / H = 10.88 × 10 6/T 2 − 41.52 × 10 3/T + 5.61 where T is temperature in Kelvin. Our D–H serpentine–water fractionation factors calibrated by experiment decrease with temperature and form a consistent trend with fractionation factors derived from lower temperature field calibrations. By regression of these data, we have refined and extended the D–H fractionation curve from 25 to 450 °C, 50 MPa as follows: 1000 ln α serpentine – water D – H = 3.436 × 10 6/T 2 − 34.736 × 10 3/T + 21.67 where T is temperature in Kelvin. These new data should improve the application of D–H and 18O– 16O isotopes to constrain the temperature and origin of hydrothermal fluids responsible for serpentine formation in a variety of geologic settings.

Oxygen isotope fractionation and equilibration kinetics between CO 2 and H 2O as a function of salinity of aqueous solutions

Oxygen isotope fractionation and equilibration kinetics between CO 2 and H 2O have been investigated at 313 K for salinities ( S) ranging from 0 to 250 g L − 1 . In this range of salinity, times needed to reach oxygen isotope equilibrium between CO 2 and H 2O increase from 4 h to 12 h. Isotopic exchanges are comparable for KCl and NaCl-like (sea salt) solutions and are described by first-order kinetic reactions with ln( k) = − 8.1485(± 0.0057) − 0.00474(± 3.87 × 10 − 5 ) S. The oxygen isotope fractionation factor between CO 2 and H 2O increases with salinity for both sea salt and KCl solutions with concentrations ranging from 0 to 250 g L − 1 according to the following equation: 1000 ln( α CO2–H2O)sea salt = 37.02(± 5 × 10 − 3 ) + 3.96 × 10 − 3 (± 1.1 × 10 − 4 ) S − 6.38 × 10 − 6 (± 4.5 × 10 − 7 ) S 2 ( R 2 = 0.998). The oxygen isotope analysis of seawater samples with a salinity of 35 g L − 1 requires minor corrections of − 0.15‰ (V–SMOW). However, oxygen isotope ratios are overestimated by 0.4‰ to 0.6‰ in the case of highly saline natural waters (100 < S < 250 g L − 1 ). Corrections of the oxygen isotope ratios due to changes in the salinity-dependent fractionation factors between CO 2 and H 2O must be taken into account during the study of waters sampled from salt marshes, hypersaline lakes and lagoons, or hydrothermal brines.

Oxygen isotope fractionation of dissolved oxygen during reduction by ferrous iron

The oxygen isotope fractionation factor of dissolved oxygen gas has been measured during inorganic reduction by aqueous FeSO 4 at 10−54 °C under neutral (pH 7) and acidic (pH 2) conditions, with Fe(II) concentrations ranging up to 0.67 mol L −1, in order to better understand the geochemical behavior of oxygen in ferrous iron-rich groundwater and acidic mine pit lakes. The rate of oxygen reduction increased with increasing temperature and increasing Fe(II) concentration, with the pseudo-first-order rate constant k ranging from 2.3 to 82.9 × 10 −6 s −1 under neutral conditions and 2.1 to 37.4 × 10 −7 s −1 under acidic conditions. The activation energy of oxygen reduction was 30.9 ± 6.6 kJ mol −1 and 49.7 ± 13.0 kJ mol −1 under neutral and acidic conditions, respectively. Oxygen isotope enrichment factors ( ε) become smaller with increasing temperature, increasing ferrous iron concentration, and increasing reaction rate under acidic conditions, with ε values ranging from −4.5‰ to −11.6‰. Under neutral conditions, ε does not show any systematic trends vs. temperature or ferrous iron concentration, with ε values ranging from −7.3 to −10.3‰. Characterization of the oxygen isotope fractionation factor associated with O 2 reduction by Fe(II) will have application to elucidating the process or processes responsible for oxygen consumption in environments such as groundwater and acidic mine pit lakes, where a number of possible processes (e.g. biological respiration, reduction by reduced species) may have taken place.

Theoretical calculation of oxygen isotope fractionation factors in carbonate systems

Using established methods of statistical mechanical calculation and a recent compilation of vibrational frequency data, we have computed oxygen isotope reduced partition function ratios ( β values) for a large number of carbonate minerals. The oxygen isotope β values of carbonates are inversely correlated to both the mass and radius of the cation bonded to the carbonate anion but neither correlation is good enough to be used as a precise and accurate predictor of β values. There is an approximately 0.6% relative increase in the β values of aragonite per 10 kbar increase in pressure. These estimates of the pressure effect on β values are broadly similar to those deduced previously for calcite using the methods of mineral physics. In comparing the β values of our study with those derived recently from first-principles lattice dynamics calculations, we find near-perfect agreement for calcite and witherite (<0.3% deviation), reasonable agreement for dolomite (<0.9% deviation) and somewhat poorer agreement for aragonite and magnesite (1.5–2% deviation). In the system for which we have the most robust constraints, CO 2–calcite, there is excellent agreement between our calculations and experimental data over a broad range of temperatures (0–900 °C). Similarly, there is good to excellent correspondence between calculation and experiment for most other low to moderate atomic mass carbonate minerals (aragonite to strontianite). The agreement is not as good for high atomic mass carbonates (witherite, cerussite, otavite). In the case of witherite and cerussite, the discrepancy may be due, in part, to our calculation methodology, which does not account for the effect of cation mass on the magnitude of vibrational frequency shifts associated with heavy isotope substitution. However, the calculations also reveal an incompatibility between the high- and low-temperature experimental datasets for witherite and cerussite. Specifically, the shapes of fractionation factor versus 1/ T 2 curves in the calcite–witherite and calcite–cerussite systems do not conform to the robust constraints on the basic shape of these curves provided by theory. This suggests that either the high- or low-temperature datasets for both minerals is in error. Dolomite–calcite fractionation factors derived from our calculations fall within the wide range of fractionations for this system given by previous experimental and natural sample studies. However, our compilation of available low-temperature (25–80 °C) experimental data reveal an unusual temperature dependence of fractionations in this system; namely, the data indicate an increase in the magnitude of fractionations between dolomite (or proto-dolomite) and calcite with increasing temperature. Such a trend is incompatible with theory, which stipulates that fractionations between carbonate minerals must decrease monotonically with increasing temperature. We propose that the anomalous temperature dependence seen in the low-temperature experimental data reflect changes in the crystallinity and degree of cation ordering of the dolomite phase over this temperature interval and the effect these changes have on the vibrational frequencies of dolomite. Similar effects may be present in natural systems at low-temperature and must be considered in applying experimental or theoretical fractionation data to these systems. In nearly all cases, carbonate mineral–calcite fractionation factors given by the present calculations are in as good or better agreement with experimental data than are fractionations derived from semi-empirical bond strength methods.

Intra-tooth study of modern rhinoceros enamel δ18O: Is the difference between phosphate and carbonate δ18O a sound diagenetic test?

The carbonate and phosphate oxygen isotopic difference (Δ 18O C–P) has been suggested as a test for diagenetic alteration of isotope values in fossil apatite samples. The generally accepted Δ 18O C–P value for well-preserved enamel is ~ 9‰. Only a few studies have focussed on the potential natural variability of this difference in modern and fossil mammal teeth. The aim of this study is to document the value of the carbonate-phosphate oxygen isotopic fractionation factor ( α CO3–PO4) at an intra-tooth level. We measured the enamel phosphate and carbonate oxygen isotopic compositions of two upper molars (M1 and M2) of Rhinoceros unicornis following a sequence from the apex to the cervix of the crown. The average α CO3–PO4 is 1.0082 ± 0.0007. The mean value of the Δ 18O C–P difference is 8.4 ± 0.7‰, close to the literature values given for other species. However, we also show that the values of the Δ 18O C–P difference for R. unicornis upper molars can vary on a ~ 2‰ range in the same tooth. We develop a physiological model to explain the origin of inter-species and intra-tooth Δ 18O C–P difference and to understand their implications on the status of Δ 18O C–P as a tool for identifying diagenesis. The results imply that, by increasing the dataset of phosphate and carbonate oxygen isotopic composition for modern mammal species, we will be able to define a maximal and precise envelope corresponding to the well-preserved samples and to check a possible species dependence on the δ 18O C/ δ 18O P relationship.

An expression for the overall oxygen isotope fractionation between the sum of dissolved inorganic carbon and water

Stable oxygen isotope fractionation factors (α's) between individual dissolved carbonate species (CO2(aq.), H2CO3, HCO3−, CO32−) and water, and their sum ( = [CO2(aq.)] + [H2CO3] + [HCO3−] + [CO32−]) and water are fundamental geochemical parameters. While values of , α, and have recently been examined, an overall mass balance for stable oxygen isotopes in the carbonate system still needs to be derived. The theoretical treatment of this mass balance is not trivial (in contrast to, e.g., stable carbon isotopes) because (1) oxygen is exchanged with H2O and (2) complications arise from the nature of the oxygen mole fractions. I use the recently reported individual α's to derive an expression for the overall oxygen isotope fractionation between and water. This allows the calculation of α(−H2O) at any given pH of the solution. The result differs from a relationship published earlier, and implications for previous studies are discussed. Numerical routines in MATLAB and Excel for the present calculation have been developed and are available as auxiliary material (Software S1). Applications of the expression include, for example, examination of biomineralization mechanisms and corrections to isotopic paleotemperature equations. The latter application can significantly alter reconstructions of Phanerozoic climate change.

A stable isotope study of microbial dolomite formation in the Coorong Region, South Australia

The Coorong Region of South Australia is a classic location for the study of modern dolomite formation. A number of ephemeral lakes here contain dolomite, whilst other lakes in close proximity, contain aragonite with minor magnesite. Our previous field and experimental work has revealed that sulphate-reducing bacteria play a key role in overcoming kinetic inhibitors to dolomite formation in this area. We here use a combination of oxygen, carbon and sulphur isotopes to help to further constrain the processes leading to dolomite precipitation in modern hypersaline environments. We report isotopic data from four dolomite and three aragonite containing lakes. δ 34S data for dissolved lake- and porewater sulphates reveal 34S enrichment of residual sulphate in all seven lakes, consistent with bulk closed system microbial fractionation by sulphate-reducing bacteria. Calculations using temperature-dependent oxygen isotope fractionation factors between dolomite and the solution from which it precipitated are consistent with a strong microbial component to carbonate precipitation in these lakes. A comparison of δ 18O data from dolomite and aragonite sediment and from their associated lakewaters indicates seasonal mineral precipitation occurs late in the evaporative cycle of the lakes within a yoghurt-like saturated mud horizon coincident with the zone of sulphate reduction. Surprisingly, δ 13C data from the carbonates do not show the characteristic wide range of values often associated with ‘organogenic’ carbonates, but instead show that these bacterially-mediated precipitates incorporate carbon primarily from the inorganic, lakewater reservoir, partially diluted by an organic component. A ‘Coorong type’ microbial model of dolomite formation may be particularly applicable to, for example, Precambrian sedimentary dolomites where benthic microbial communities, which are today restricted to extreme environments, dominated the shallow-marine ecosystem prior to the emergence of Metazoa.

Calibration of the calcite–water oxygen-isotope geothermometer at Devils Hole, Nevada, a natural laboratory

The δ 18O of ground water (−13.54 ± 0.05 ‰) and inorganically precipitated Holocene vein calcite (+14.56 ± 0.03 ‰) from Devils Hole cave #2 in southcentral Nevada yield an oxygen isotopic fractionation factor between calcite and water at 33.7 °C of 1.02849 ± 0.00013 (1000 ln α calcite–water = 28.09 ± 0.13). Using the commonly accepted value of ∂( α calcite–water)/∂ T of −0.00020 K −1, this corresponds to a 1000 ln α calcite–water value at 25 °C of 29.80, which differs substantially from the current accepted value of 28.3. Use of previously published oxygen isotopic fractionation factors would yield a calcite precipitation temperature in Devils Hole that is 8 °C lower than the measured ground water temperature. Alternatively, previously published fractionation factors would yield a δ 18O of water, from which the calcite precipitated, that is too negative by 1.5 ‰ using a temperature of 33.7 °C. Several lines of evidence indicate that the geochemical environment of Devils Hole has been remarkably constant for at least 10 ka. Accordingly, a re-evaluation of calcite–water oxygen isotopic fractionation factor may be in order. Assuming the Devils Hole oxygen isotopic value of α calcite–water represents thermodynamic equilibrium, many marine carbonates are precipitated with a δ 18O value that is too low, apparently due to a kinetic isotopic fractionation that preferentially enriches 16O in the solid carbonate over 18O, feigning oxygen isotopic equilibrium.

Response to the Comment by J. Horita and R.N. Clayton on “The studies of oxygen isotope fractionation between calcium carbonates and water at low temperatures”

The apparent inconsistency in calcite–water fractionation does occur between the arithmetic combination of Zhou and Zheng [Zhou G.-T., and Zheng Y.-F. (2003) An experimental study of oxygen isotope fractionation between inorganically precipitated aragonite and water at low temperatures. Geochim. Cosmochim. Acta 67, 387–399] and the experimental determination of Zhou and Zheng [Zhou G.-T., and Zheng Y.-F. (2005) Effect of polymorphic transition on oxygen isotope fractionation between aragonite, calcite and water: a low-temperature experimental study. Am. Mineral 90, 1121–1130]. To resolve this issue is to acknowledge whether or not the isotope salt effect of dissolved minerals would occur on oxygen isotope exchange between water and the minerals of interest. The question is whether or not a term of mineral–water interaction should be taken into account when calculating mineral–water 10 3ln α factors by an arithmetic combination between theoretical 10 3ln β factors for mineral and water, respectively. The hydrothermal experiments of Hu and Clayton [Hu G.-X., and Clayton R.N. (2003) Oxygen isotope salt effects at high pressure and high temperature, and the calibration of oxygen isotope geothermometers. Geochim. Cosmochim. Acta 67, 3227–3246] demonstrate the absence of isotope salt effect on the oxygen isotope fractionation between calcite and water, and this abnormal behavior reasonably explains the so-called inconsistency in the calcite–water fractionations of Zhou and Zheng (2003, 2005). We argue that the mineral–water correction is still necessary for calculation of fractionations in mineral–water systems. New experimental data for oxygen isotope fractionations involving dolomite and cerussite are consistent with the calculations of Zheng [Zheng Y.-F. (1999a) Oxygen isotope fractionation in carbonate and sulfate minerals. Geochem. J. 33, 109–126], but also shed light on the assumptions used in modifying the increment method. We argue that the modified increment method has developed into a theoretical mean of predictive power for calculation of oxygen isotope fractionation factors for crystalline minerals of geochemical interest.

Novel Corresponding-States Principle Approach for the Equation of State of Isotopologues: H218O as an Example

The corresponding-states principle (CSP) has been considered for the development of the equations of state (EOS) of minor isotopologues that are usually unknown. We demonstrate that, for isotopologues of a given molecular fluid, a general extended multi-parameter corresponding-states EOS can be reduced to the three-parameter EOS, utilizing the critical parameters (temperature and density) and Pitzer's acentric factor as correlation parameters. Appropriate general CSP mathematical formalism and equations for constructing the EOS of minor isotopologues are described in detail. The formalism and equations were applied to isotopologues of water and demonstrated that the isotopic effect on the critical parameters and the acentric factor of H(2)(18)O can be successfully calculated from the EOS of H2O and experimental data on the isotope effects (liquid-vapor isotope fractionation factor and molar volume isotope effect). We have also shown that the experimental data on the vapor pressure isotope effect (VPIE) for 18O-substituted water are inconsistent within the framework of thermodynamics with the liquid-vapor oxygen isotope fractionation factor. The novel approach of CSP to isotopologues developed in this study creates a new opportunity for constructing the EOS of minor isotopologues for many other molecular fluids.

Mechanisms of equilibrium and kinetic oxygen isotope effects in synthetic aragonite at 25 °C

Aragonite was precipitated in the laboratory at 25 °C in isotopic equilibrium with Na-Ca-Mg-Cl-CO 3 solutions at two different pH values (i.e., pH = ∼8.2 and ∼10.8) by the constant addition method. On the basis of the oxygen isotope composition of the aragonite precipitates, it was demonstrated that the equilibrium aragonite–water fractionation factor is independent of the pH of the parent solution and equal to: 1000 ln α ( aragonite – H 2 O ) = 29.12 ± 0.09 To elucidate the mechanism(s) of aragonite precipitation, the equilibrium oxygen isotope fractionations between HCO 3 - as well as CO 3 2 - and H 2O at 25 °C were also determined experimentally to estimate the oxygen isotope compositions of the HCO 3 - and CO 3 2 - ions in the parent solutions. The isotopic composition of BaCO 3 (witherite) precipitated quantitatively from solutions of various pH values and the speciation of carbonic acid species in these solutions were combined to yield the following oxygen isotope fractionation factors: 1000 ln α ( HCO 3 - – H 2 O ) = 30.53 ± 0.08 1000 ln α ( CO 3 2 - – H 2 O ) = 23.71 ± 0.08 The oxygen isotope composition of witherite obtained from two types of fractional precipitation experiments revealed that CO 3 2 - ions are preferentially incorporated into the growing crystal. In addition, preferential deprotonation of isotopically light HCO 3 - ions and the incorporation of the light CO 3 2 - isotopologues are proposed to account for the kinetic isotope effects observed in the course of aragonite and witherite precipitations.

Effect of massive gas hydrate formation on the water isotopic fractionation of the gas hydrate system at Hydrate Ridge, Cascadia margin, offshore Oregon

Because gas hydrate is preferentially enriched in the heavy water isotopes, the δ18O and δD values of pore waters collected from gas hydrate–bearing sediment can provide information on the abundance and mechanisms of gas hydrate formation. Pore waters sampled from deep‐seated (40 to 125 mbsf) gas hydrate deposits in Hydrate Ridge during ODP Leg 204 show depletion in dissolved Cl− and enrichments in 18O and D due to gas hydrate destabilization during core recovery. The oxygen and hydrogen isotopic fractionation factors (αO = 1.0025 and αH = 1.022) estimated from an extensive data set (n = 30 samples) correspond to experimentally determined values. In contrast, pore waters from shallow samples (&lt;25 mbsf) at the ridge summit (n = 32) are highly enriched in dissolved Cl− and depleted in 18O and D, consistent with formation of massive gas hydrate deposits at rates faster than those at which these anomalies would be removed by advection or diffusion. The water isotopic fractionation factors in the brine are significantly lower than those experimentally determined, with αO of 1.0010 (average value of 1.0012) and αH of 1.008 (average value of 1.008). We discuss several factors that may be causing this anomalous fractionation and suggest that low gas occupancy in hydrate lattice (high hydration number) may be responsible for the observed small fractionation. If this were the case, the oxygen and hydrogen fractionation may serve as an indicator of hydration number during formation of gas hydrate in natural systems.

On the Direction and Magnitude of Oxygen Isotope Fractionation Between Calcite and Aragonite at Thermodynamic Equilibrium

Oxygen isotope fractionation factors between calcium carbonates and water have been applied to ancient marine geochemistry principally for the purpose of geothermometry. The problem was encountered, however, with respect to the direction and magnitude of oxygen isotope fractionation between calcite and aragonite at thermodynamic equilibrium. This basically involves sound understanding of both thermodynamics and kinetics of oxygen isotope fractionation between inorganically precipitated carbonate and water at low temperatures. Thus the crucial issues are to acknowledge the processes of chemical reaction and isotopic exchange during precipitation of CaCO3 minerals in solution, the kinetic mechanism of isotope equilibrium or disequilibrium, the effect of polymorphic transition from metastable aragonite to stable calcite under hydrous or anhydrous conditions, and the presence or absence of isotope salt effect on oxygen isotope exchange between carbonate and water in response to the hydrous or anhydrous conditions at thermodynamic equilibrium. Because good agreements exist in carbonate–water oxygen isotope fractionation factors between theoretical calculations and experimental determinations, it is encouraging to applying the thermodynamic and kinetic data to isotopic paleothermometry and geochemical tracing.

Experimental studies of low-T oxygen isotope fractionation between carbonate minerals and water

Oxygen isotope fractionation factors between calcium carbonates and water have been applied to ancient marine geochemistry principally for the purpose of geothermometry. The problem was encountered, however, with respect to the direction and magnitude of O isotope fractionation between calcite and aragonite at thermodynamic equilibrium. This study presents a series of synthesis experiments at 0–70 C to resolve these intriguing issues. The experiments were designed with oneand two-step approaches, respectively, corresponding to the two complex anions of dissolved carbonate prior to aragonite precipitation. Witherite and metastable aragonite were directly synthesized by the direct precipitation and the overgrowth techniques, respectively. The twostep experiments delivered equilibrium fractionations between the precipitated carbonate and water, whereas one-step experiments did not. Significant difference was observed in experiment-derived calcite–aragonite fractionations at above 25 C; the two-step experiments at different periods from 20 to 60 days demonstrate that the dissolved [CO3] 2 and H2O have achieved O isotope equilibrium prior to the aragonite precipitation. Thus this difference cannot be ascribed to rapid (disequilibrium) precipitation of aragonite from either [CO3] 2 or [HCO3] dominant solutions. The experimental results show that the equilibrium fractionations derived from the two-step experiments are independent of the effects of high solution pH, the low O/O ratio of [CO3] 2 and rapid precipitation. Our experiments suggest negligible effect of Mg concentrations on the O isotope fractionations within the analytical error, indicating that no salt effect of Mg cation occurred under the experimental conditions. The O isotope composition of precipitated aragonite does not depend on the concentration of dissolved carbonate species in aqueous solutions provided that the O isotope equilibrium is really achieved thermodynamically between dissolved carbonate and water. Although the effects of the chemical composition of solutions may occur in synthesis experiments on the carbonate–water fractionations, enough attention must be paid to the kinetic mechanism of O isotope exchange and equilibrium in carbonate-water systems. It is of critical importance in judging the achievement of O isotope equilibrium between the dissolved carbonate complex anions and water prior to the precipitation of either aragonite or calcite. According to the currently available observations, it appears that the direction of O isotope fractionation between calcite and aragonite is positive rather than negative at thermodynamic equilibrium.

Preferential formation of 13C– 18O bonds in carbonate minerals, estimated using first-principles lattice dynamics

Equilibrium constants for internal isotopic exchange reactions of the type: Ca 12 C 18 O 16 O 2 + Ca 13 C 16 O 3 ↔ Ca 13 C 18 O 16 O 2 + Ca 12 C 16 O 3 for individual CO 3 2− groups in the carbonate minerals calcite (CaCO 3), aragonite (CaCO 3), dolomite (CaMg(CO 3) 2), magnesite (MgCO 3), witherite (BaCO 3), and nahcolite (NaHCO 3) are calculated using first-principles lattice dynamics. Calculations rely on density functional perturbation theory (DFPT) with norm-conserving planewave pseudopotentials to determine the vibrational frequencies of isotopically substituted crystals. Our results predict an ∼0.4‰ excess of 13 C 18 O 16 O 2 2 − groups in all studied carbonate minerals at room-temperature equilibrium, relative to what would be expected in a stochastic mixture of carbonate isotopologues with the same bulk 13C/ 12C, 18O/ 16O, and 17O/ 16O ratios. The amount of excess 13 C 18 O 16 O 2 2 − decreases with increasing temperature of equilibration, from 0.5‰ at 0 °C to <0.1‰ at 300 °C, suggesting that measurements of multiply substituted isotopologues of carbonate could be used to infer temperatures of ancient carbonate mineral precipitation and alteration events, even where the δ 18O of coexisting fluids is uncertain. The predicted temperature sensitivity of the equilibrium constant is ∼0.003‰/°C at 25 °C. Estimated equilibrium constants for the formation of 13 C 18 O 16 O 2 2 − are remarkably uniform for the variety of minerals studied, suggesting that temperature calibrations will also be applicable to carbonate minerals not studied here without greatly compromising accuracy. A related equilibrium constant for the reaction: Ca 12 C 18 O 16 O 2 + Ca 12 C 17 O 16 O 2 ↔ Ca 12 C 18 O 17 O 16 O + Ca 12 C 16 O 3 in calcite indicates formation of 0.1‰ excess 12C 18O 17O 16O 2− at 25 °C. In a conventional phosphoric acid reaction of carbonate to form CO 2 for mass-spectrometric analysis, molecules derived from 13 C 18 O 16 O 2 2 − dominate (∼96%) the mass 47 signal, and 12C 18O 17O 16O 2− contributes most of the remainder (3%). This suggests that carbonate internal equilibration temperatures can be recovered from acid-generated CO 2 if abundances of isotopologues with mass 44–47 can be measured to sufficient precision. We have also calculated 18O/ 16O and 13C/ 12C reduced partition function ratios for carbonate minerals, and find them to be in good agreement with experiments and empirical calibrations. Carbon and oxygen isotope fractionation factors in hypothetical 40Mg—magnesite and 40Ba—witherite indicate that M 2+-cation mass does not contribute significantly to equilibrium isotopic fractionations between carbonate minerals.