Oxygen Isotope Fractionation Factors Research Articles

Calculation of oxygen isotope fractionation in hydroxyl-bearing silicates

The modified increment method has been applied to the calculation of oxygen isotope fractionation factors for hydroxyl-bearing silicate minerals. The order of 18O enrichment obtained in common rock-forming minerals is: pyrophyllite > kaolinite > tourmaline ⩾ talc > prehnite ⩾ topaz > illite > phengite > lepidolite ⩾ muscovite ⩾ staurolite > epidote > glaucophane > serpentine ⩾ chlorite > tremolite > hornblende > phlogopite ⩾ biotite > humite > norbergite > ilvaite. Hydroxyl-bearing silicates are enriched in 18O relative to hydroxyl groups but depleted in 18O relative to anhydrous counterparts. Three sets of self-consistent fractionation factors: between quartz and the hydroxyl-bearing silicate minerals, between calcite and the silicate minerals, and between the silicate minerals and water, have been calculated for a temperature range of 0–1200°C. The fractionation factors calculated for mineral pairs are applicable to isotopic geothermometry in igneous, metamorphic and sedimentary petrology. They can be used as a test of isotopic equilibrium or disequilibrium in natural mineral assemblages over all temperature ranges of geological interest. The difference in oxygen isotope composition between the hydroxyl-bearing mineral and the OH group is quantitatively demonstrated to be temperature dependent and, therefore, can be used as a single-mineral geothermometer.

Oxygen isotope fractionation in SiO2 and Al2SiO5 polymorphs: effect of crystal structure

Oxygen isotope fractionation factors for SiO 2 and Al 2 SiO 5 polymorphs have been calculated by means of the modified increment method. The results suggest that the 18 O is depleted in the high pressure forms (i.e. the more dense phases). The effect of crystal structure on isotopic partitioning in polymorphic phases is important, because the cation-oxygen interatomic distances and coordination numbers can change due to compression or thermal expansion in geological processes

Oxygen-isotope fractionation in titanium-oxide minerals at low temperature

Oxygen-isotope results for natural and synthetic Ti-oxide minerals are presented. Rutile-water oxygen-isotope fractionation factors (α) of 1.0061 ± 0.0006 (1σ) at 22 °C and 1.0030 ± 0.001 ( 1σ) at 50°C have been obtained for synthesis experiments under a variety of conditions. These results are in close agreement with recent, semiempirical predictions and with values determined using naturally occurring rutile. Nevertheless, the true fractionation factors may lie toward the upper bounds of the error limits, given the possibilities that (1) isotopic equilibrium may not have been achieved in some experiments and that (2) hydration water may not have been removed completely from some samples prior to analysis. Anatase-water oxygen-isotope fractionation factors of 1.0087 ± 0.0015 (1σ) at 22°C and 1.0049 ± 0.0005 (1σ) at 50°C were obtained from synthesis experiments in which sulphate was not added to the starting solutions. Higher values of α for anatase-water vs. rutile-water are in accord with predictions based on theoretical and empirical considerations. However, fractionation factors increased dramatically to maxima of 1.0143 at 22°C and 1.0081 at 50°C for anatase-water syntheses in which sulphate ions were added to the starting solutions. The large increase in the δ 18O values of anatase produced in such solutions may result from donation of, or exchange with, high- 18O sulphate oxygen during formation of Ti-O-OH-SO 4 complexes. The range of fractionation factors obtained from the syntheses is mirrored by the range obtained for natural samples. This similarity suggests that the isotopic composition of anatase is controlled not only by its temperature of formation and the isotopic composition of associated water, but also by the mechanism of formation. Moreover, in cases where rutile formed by inversion of less stable anatase, rutile appears to at least partially inherit its isotopic composition from the precursor polymorph. However, coprecipitated quartz and rutile (formed directly, that is, in the absence of an anatase inversion) has potential as an oxygen-isotope geothermometer for low-temperature environments.

An experimental study of oxygen isotope fractionation in the quartz-wolframite-water system

Oxygen isotope fractionation was experimentally studied in the quartz-wolframite-water system from 200 to 420 °C. The starting wolframite was synthesized in aqueous solutions of Na2WO4 · 2H2O + FeCl2 · 4H2O or MnCl2 · 4H2O. The starting solutions range in salinity from 0 to 10 equivalent wt.% NaCl. Experiments were conducted in a gold-lined stainless steel autoclave, with filling degrees of about 50%. The results showed no significant difference in equilibrium isotope fractionation between water and wolframite, ferberite and huebnerite at the same temperature (310 °C ). The equilibrium oxygen isotope fractionation factors of wolframite and water tend to be equal with increasing temperature above 370 °C, but to increase significantly with decreasing temperature below 370 °C: 1000 ln αwf-H2o= 1.03×106T−2-4.91 (370 °C ±200 °C ) 1000 ln αwf-H2o = 0.21×106T −2-2.91 (420 °C -370 °C ±)

Room temperature oxygen isotope exchange between liquid CO 2 and H 2O

The equilibrium oxygen isotope fractionation factor for liquid carbon dioxide and water at 25.3°C is 41.94 ° 0.27%. as determined by direct exchange. Oxygen isotope exchange between the two phases is initially very rapid with over 50% of the exchange occurring in the first half hour. Complete exchange is only achieved after ~3 days. Both the speed of reaction between liquid carbon dioxide and water and the magnitude of the fractionation at room temperature have important implications for oxygen isotope exchange experiments involving carbon dioxide where trace quantities of water may be present. Reaction between the carbon dioxide and water will enrich the carbon dioxide in 18O, modifying its initial composition in a predictable way.

Oxygen isotope fractionation in wolframite

Thermodynamic oxygen isotope factors for wolframite are calculated using the modified increment method and assuming a predominant contribution of the 18 O-increment in complex anion to the I- 18 O index of complex mineral. The oxygen isotope fractionation factors obtained between quartz, wolframite and water are : 10 3 ln α quartz-wolframite =0.38×10 6 /T 2 +4.76×10 3 /T-2.48, 10 3 ln α wolfr a mite-water =3.86×10 6 /T 2 -8.54×10 3 /T + 1.44 for the temperature range from 0 to 1200°C. No difference in fractionation between ferberite and huebnerite is observed, indicating that the substitution of Fe for Mn does not significantly affect the 18 O/ 16 O partitioning in the wolframite series. When the present calibration for the quartz-wolframite system is applied to isotopic geothermometry in hydrothermal tungsten mineralizations, the temperatures calculated are well comparable to results from other geothermometric methods

Oxygen isotope fractionation in hematite and magnetite: A theoretical calculation and application to geothermometry of metamorphic iron-formations

Thermodynamic oxygen isotope factors for hematite and magnetite are calculated using the modified increment method and incorporating the theoretical and experimental reduced partition function ratios determined for quartz and water in the temperature range 0 to 1200 o C. The obtained oxygen isotope fractionation factors between hematite, magnetite and water are expressed as 10 3 ln α Fe2O3-H2O =2.69×10 6 /T 2 -12.82×10 3 /T+3.78, 10 3 ln α Fe3 O 4- H 2 O=3.02×10 6 /T 2 -12.00×10 3 /T+3.31, and those between quartz and hematite or magnetite are expressed as 10 3 ln α Si O 2-FeO3 =1.55×10 6 /T 2 +9.05×10 3 /T-4.82, 10 3 ln α SiO2-Fe3O4 =1.22×10 6 /T 2 +8.22×10 3 /T-4.35. The present results are well in agreement with existing theoretical, experimental and empirical calibrations involving magnetite and hematite. A significant oxygen isotope fractionation is obtained between hematite and magnetite, whith hematite being depleted in 18 O relative to magnetite. This matches with the observation from Precambrian iron-formation in the Hamersley Range, Australia. The rule of effect of oxidation state on the oxygen isotope fractionation seems not to work for the two iron oxides. Previous experimental calibration of magnetite-water system by reduction of hematite to magnetite is concordant with the present result for hematite-water system. This is attributed to the quantitative redox of iron in the experiments, that can in turn explain the nearly identical isotope composition of oxygen in coexisting hematite and magnetite pairs from the banded iron-formation in the Iron Quadrangle, Brazil

Oxygen and carbon isotope fractionations between CO 2 and calcite

Oxygen and carbon isotopic fractionation factors have been measured by direct exchange between carbon dioxide and calcite over the temperature range 400–950°C at pressures of 1–13 kilobars. Equilibrium constants for oxygen exchange are well determined from 400 to 800°C, and differ substantially from previous experimental and theoretical estimates. Carbon isotopic fractionations have larger uncertainties at all temperatures. Both oxygen and carbon isotopic fractionations are well reproduced by statistical mechanical calculations with simple treatment of the calcite lattice vibrations. A recalculated calcite-graphite fractionation curve based on new calculations of partition functions ratios for calcite suggests that this system is well suited for high-temperature isotopic thermometry of graphitic marbles as first proposed by Valley and O'Neil (1981). The rate of exchange of both oxygen and carbon between carbon dioxide and calcite is rapid, and proceeds by a combination of diffusion and minor recrystallization.

Recent advances in the oxygen and hydrogen isotope geochemistry of silicates and oxides in low temperature systems

The analytical technique for determining oxygen isotopic ratios of wairakite has been carefully examined. Channel water in wairakite was separated from its aluminosilicate framework by dehydration in vacuum at various temperatures from 250 to 950°C. Dehydration at temperatures greater than 400°C resulted in erroneous isotopic ratios of the framework oxygen because of oxygen isotopic exchange between the framework oxygen and dehydrating water. The optimum conditions are dehydration of the channel water at 300°C for 24 h, followed by stepwise heating at 350°C for 12 h, 400°C for 5 h, and finally at 700† for 15 min. With this stepwise heating, isotopic measurements of wairakite framework oxygen can be satisfactorily achieved within an accuracy of ±0.1‰.The oxygen isotope fractionation factor for the wairakite-water system has been experimentally determined at temperatures between 250 and 400°C and pressures between 0.5 to 1.5 kbar. The equilibrium fractionation factor between framework oxygen and external water is expressed by the following equation: 1031nαframeworkoxygen-externalwater=2.46(106/T2)−1.76 The fractionation factor for the channel water-external water is given by 1031nαchannelwater-externalwater=0.79(106/T2)−3.07 In the temperature range studied, the framework-water fractionation of wairakite is very similar to that of analcime (Karlsson and Clayton, 1990a) and stilbite (Feng and Savin, 1993b). The negative fractionation for the channel water-external water of wairakite (−3 to 0‰) is consistent with that for analcime and stilbite, but the magnitude is less than that of analcime and stilbite.Almost complete oxygen isotopic exchange during the hydrothermal runs indicates that the exchange between wairakite framework oxygen and water is rapid compared to that of the other silicate-water systems. This implies that the δ 18O value of natural wairakite is controlled by retrograde re-equilibration under hydrothermal conditions. The grain size did not change during the hydrothermal runs, suggesting that the exchange mechanism was dominated by sorption-exchange-desorption processes as proposed by Karlsson and Clayton (1990a). However, scanning electron microscopy images of the run products showed that the grain surfaces of the exchanged wairakite were covered by newly deposited crystals, suggesting that dissolution-crystallization was also involved to a small extent.

Oxygen isotope effects associated with the solid-state α-FeOOH to α-Fe2O3 phase transformation

Samples of synthetic and natural goethites were subjected to isothermal dehydration under both closed- and open-systein conditions at various teinperatures ranging from 160 to 300°C. The oxygen isotope ratios of the dehydration product (hematite) were systematically different for open- and closedsystein dehydration. In every instance, open-systein dehydration resulted in 18 O enrichments of the product hematite relative to the starting goethite. For all but a few cases, closed-system dehydration produced an 18 O depletion in the residual mineral relative to the starting goethite. These oxygen isotope effects indicate that significant mineral-vapor isotope exchange occurred during the solid-state, closed-system transformation of goethite to ∗∗ hematite. Calculated values of the apparent closed-system oxygen isotope fractionation factor between hematite and the corresponding liquid water suggest that equilibrium isotope exchange may have been approached in those samples for which the extent of closed-system dehydration was at least 95%. The rapidity of the mineral-vapor isotope exchange associated with the solid-state goethite to hematite transformation is indicated by the fact that the dehydration times (for those samples apparently approaching equilibrium) ranged from only 24 to 68 hours. These results suggest that the oxygen isotope ratios of many hematites in low temperature geological systems reflect the environmental conditions associated with their formation from some hydrous precursor. Information about the 18 O 16 O ratio in the hydrous precursor itself is not likely to be directly preserved in the hematite

Stable isotope fractionation factors of water in hydrated saline mineral-brine systems

Hydrogen and oxygen isotope fractionation factors of water between water of crystallization of carnallite, bischofite and tachyhydrite and the mother solutions were determined from the isotopic activity ratios of the mother solutions (α a ) in the temperature range of 10–40°C. Deuterium is depleted and 18 O enriched in the water of crystallization of all these minerals. No clear temperature dependence can be observed. The published fractionation factors in the literature, which were all determined from the isotopic concentration ratios of the mother solutions (α c ) were corrected to α a using the “salt effect” coefficients. It is proposed that only the α a values can be applied to the natural settings.

Electrostatic characterization of oxygen sites in minerals

The electrostatic potentials of approximately 500 anion sites in 165 rock-forming mineral end-members have been computed and presented along with coordination numbers, coordinating cations, and mean cation-anion distances. The electrostatic potentials computed for the individual oxygen sites range from a low of 19.56 v for the O2 site of (β-Mg 2SiO 4 (wadsleyite) to a high of 36.4 v for soda niter, with an unweighted average of 28.58 v and standard deviation of 2.51 v. The site potentials tend to vary inversely with coordination number, so that in the silicates the bridging oxygen sites have the highest potentials, oxygens bonded to one Si have lower potentials, and oxygens not bonded to silicon have the lowest. The site potentials of hydroxyl oxygens (with well-determined H-positions) average 28.3 v with a standard deviation of 2.0 v. The potentials of fluorine sites and hydroxyl sites modeled without the H, as a single charge of −1, range from 10.3 v to 16.8 v, with an average value of 10.87 and a standard deviation of 1.85 v. The electrostatic site potentials may thus be used along with coordinating cation types to identify possible sites for hydroxyl substitution in minerals. The average oxygen site potentials for various minerals weighted on the number of sites per unit cell range from a low of 20.9 v for lime (CaO) to a high of 36.4 v for soda niter and are strongly correlated with observed oxygen isotope fractionation trends in minerals. This correlation is best for minerals containing light cations ( Z < 20), so that minerals with larger oxygen site potentials tend to concentrate the heavier oxygen isotope. This correlation is sufficiently good to permit qualitative estimation of oxygen isotope fractionation factors of minerals such as high-pressure silicates that have not been measured experimentally. With further work on experimental measurement of oxygen isotope fractionation between mineral pairs, quantitative prediction of isotope fractionation factors may become possible based on electrostatic site potentials of anion sites.

Oxygen isotope fractionation factors among rock-forming minerals at high temperatures

Oxygen isotope fractionation factors among rock-forming minerals at high temperatures

Carbon dioxide-water oxygen isotope fractionation factor using chlorine trifluoride and guanidine hydrochloride techniques [Erratum to document cited in CA105(22):196865q

Carbon dioxide-water oxygen isotope fractionation factor using chlorine trifluoride and guanidine hydrochloride techniques [Erratum to document cited in CA105(22):196865q

Geochemical and isotopic studies of bauxitization in the Tatun volcanic area, northern Taiwan

The Tatun volcanic area in the northern tip of Taiwan is composed of Pleistocene andesite and andesitic tuff. Its high relief causes extraordinarily heavy rainfall (> 3000 mm yr −1), especially during the winter monsoon season. Bauxite deposits were formed on the windward northern slopes. Soil samples from two profiles in the bauxite deposits were studied by chemical analysis, X-ray diffraction, differential thermal analysis and hydrogen and oxygen isotopic analyses. The soil profiles can be divided into two zones by chemical characteristics. In the upper zone (zone A), the chemical composition of the soil samples undergoes little change except depletion of Na and Ca due to leaching. The major minerals are quartz, chlorite and illite. In the lower zone (zone B), all major elements suffer rapid dissolution except Al 2O 3. Gibbsite grows drastically at the expense of clays, while quartz diminishes sharply. Fe 2O 3 and TiO 2 are enriched at the boundary between the two zones due to coprecipitation of Fe and Ti as hematite and goethite, probably as a result of increasing oxygen fugacity. Mobilization of Fe in zone B is attributed to its low Eh. Thus zone A is inferred as a permeable layer and zone B as an aquifer. The weathering rate in this area is estimated to be 12 cm ka −1. The most thoroughly leached soil samples contain up to 80% gibbsite. The high gibbsite contents and the analyses of coexisting waters allow hydrogen and oxygen isotopic fractionation factors between gibbsite and water to be calculated as 0.992 ± 0.002 and 1.016 ± 0.001, respectively.

Isotopic exchange in mineral-fluid systems. II. Oxygen and hydrogen isotopic investigation of the experimental basalt-seawater system

Oxygen and hydrogen isotopic exchange reactions between basalt and seawater at T = 300° to 500°C were investigated, using oceanic tholeiitic basalt ( δ 18O = ∼5.7%.; δD = ∼−70%. ), natural seawater ( δ 18O = 1.2%.; δD = + 5%. ) and artificial seawater ( δ 18O = −5.3%. ) as starting materials. The starting basalts varied in the crystallinity (from holocrystalline to glass) but were ground to approximately the same grain size (−100 mesh). The water/rock mass ratios ranged from 1 to 3 and the duration of the experiments ranged from 167 to 576 days. In general, depletion of 18O and enrichment of D in basalts occur at all temperatures, with the magnitudes of change being greater as temperature, time and to a lesser degree, glass content increases. The trends in isotopic shifts are directly related to changes in the style and intensity of mineralogic alterations in the basalt ( e.g., smectite at 300°C, talc-actinoiite at 400°–500°C). The changes in the δ 18O values of basalts and seawater in the experimental systems were observed to follow closely with those expected from a first-order rate law. Rate constants for the oxygen isotopic exchange between rock and water range from 10 −9.5 to 10 −8.0 moles oxygen/m 2 of solid surface/sec for temperatures of 300° to 500°C. The activation energy for the isotopic exchange reaction was calculated to be 11.5 Kcal/ mol. An application of our experimental rate data to natural systems suggests that the oxygen isotope equilibrium between basalt and seawater in the mid-oceanic ridge may take place within approximately 1000 years at 350°C. Our experimental data also suggest the equilibrium oxygen isotopic fractionation factors between the altered basalt and seawater to be 3.5 ± 0.5%. at 300°C, 2.0 ± 0.4%. at 400°C and 0.5 ± 0.25%. at 500°C. The observed hydrogen isotopic fractionation factors between the altered basalts and seawater in our experimental systems were about −74%. at 300°C. about −62%. at 400°C and about −48%. at 500°C. These fractionation factors are probably attributable to the equilibrium fractionation between Fe-rich secondary phases and seawater.

Carbon dioxide-water oxygen isotope fractionation factor using chlorine trifluoride and guanidine hydrochloride techniques

ADVERTISEMENT RETURN TO ISSUEPREVArticleNEXTCarbon dioxide-water oxygen isotope fractionation factor using chlorine trifluoride and guanidine hydrochloride techniquesJoseph P. Dugan and James. BorthwickCite this: Anal. Chem. 1986, 58, 14, 3052–3054Publication Date (Print):December 1, 1986Publication History Published online1 May 2002Published inissue 1 December 1986https://doi.org/10.1021/ac00127a033RIGHTS & PERMISSIONSArticle Views120Altmetric-Citations6LEARN ABOUT THESE METRICSArticle Views are the COUNTER-compliant sum of full text article downloads since November 2008 (both PDF and HTML) across all institutions and individuals. These metrics are regularly updated to reflect usage leading up to the last few days.Citations are the number of other articles citing this article, calculated by Crossref and updated daily. Find more information about Crossref citation counts.The Altmetric Attention Score is a quantitative measure of the attention that a research article has received online. Clicking on the donut icon will load a page at altmetric.com with additional details about the score and the social media presence for the given article. Find more information on the Altmetric Attention Score and how the score is calculated. Share Add toView InAdd Full Text with ReferenceAdd Description ExportRISCitationCitation and abstractCitation and referencesMore Options Share onFacebookTwitterWechatLinked InReddit PDF (427 KB) Get e-Alerts Get e-Alerts

Calculated oxygen isotope fractionation factors between water and the minerals scheelite and powellite

Calculated oxygen isotope fractionation factors between water and the minerals scheelite and powellite

Oxygen isotope fractionation between amorphous silica and water at 34–93°C

The fractionation factor of oxygen isotopes between quartz and water has been extensively studied from experimental1,2, theoretical3,4,5 and empirical viewpoints6 and has been widely used to estimate the isotopic composition or the temperatures of water relating to various geological processes. At lower temperatures, where hydrothermal processes, diagenesis and other processes take place, however, the fractionation factor for quartz as well as other silica minerals has not been studied thoroughly. We report here experimentally determined oxygen isotope fractionation factors between amorphous silica and water at 34–93°C, obtained by measuring the oxygen isotope ratios of amorphous silica precipitated from geothermal waters of power plants. The relationship between the fractionation factor and temperature is similar to the extrapolated relationship of quartz observed at higher temperatures. This result is useful for the oxygen isotope study of water–rock interactions at temperatures below 200°C.

Thermodynamics of the oxygen isotope fractionation involving plagioclase

The thermodynamic relationship between the oxygen isotope fractionation properties of plagioclase and its composition has been derived by treating plagioclase as a “reciprocal solution” consisting of independent cationic and anionic substitutions, namely (NaAl) 5+⇌(CaSi) 5+and 18O⇌ 16O. The results show that the logarithm of the oxygen isotope fractionation factor, α, between plagioclase and a coexisting phase varies linearly with the anorthite content of plagioclase. The proportionality constant is given by the oxygen isotope fractionation factor between albite and anorthite, and has been derived from the experimental data of two groups of workers, O'Neil and Taylor [2] and Matsuhisa et al. [3], on the isotopic fractionation between each plagioclase end-member and aqueous solutions. It is found that O'Neil and Taylor's data on isotopic exchange of plagioclase end-members with only 2–3 M chloride solution, rather than with both pure water and the chloride solution, lead essentially to zero intercept of the ln α(Ab-An) vs. 1/T 2 relation, in accord with Bottinga and Javoy's [10] conclusion about the oxygen isotope fractionation between two anhydrous silicates at T<500°C.