Stable Isotope Geochemistry of Clay Minerals

Abstract The equilibrium H- and O-isotope fractionations can be approximated by the following equations which are based on experimental, empirical and/or theoretical data: Hydrogen: 1000 ln αkaolinite-water = -2.2 × 106 × T-2 - 7.7 Oxygen: 1000 ln αkaolinite-water = 2.76 × 106 × T-2 - 6.75 1000 ln αsmectite-water = 2.55 × 106 × T-2 - 4.05 1000 ln αillite-water = 2.39 × 106 × T-2 - 3.76 The equilibrium H-isotope fractionation factors vs. 106 × T-2 for kaolinite and probably smectite and illite are monotonic curves between 350-0℃. More complex curves, with a minimum fractionation near 200℃, are probably influenced by surface effects and/or disequilibrium fractionations among the different hydrogen sites. The H-isotope fractionations between smectite-water increase by ~70‰ from Fe-poor montmorillonite to nontronite at low temperatures. The pore-interlayer water in smectite H-isotope fractionation at low temperatures is ~20 ± 10‰. The presence of organic matter can modify both the δD value of the clay analysis and its 'water' content. Clays - kaolinite, illite, smectite and probably halloysite - tend to retain their D/H and 18O/16O ratios unless subjected to more extreme diagenetic or metamorphic conditions or special local processes. Kinetic information is still only qualitative: for comparable grain sizes, hydrogen exchanges more rapidly than oxygen in the absence of recrystallization. Low-temperature diffusion coefficients cannot be calculated with sufficient precision from the higher temperature exchange data. The H- and O-isotope studies of clays can provide useful information about their conditions of formation.

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