Δ18O
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In geochemistry, paleoclimatology and paleoceanography δ 18 O or delta-O-18 is a measure of the deviation in ratio of stable isotopes oxygen-18 ( 18 O) and oxygen-16 ( 16 O). It is commonly used as a measure of the temperature of precipitation, as a measure of groundwater/mineral interactions, and as an indicator of processes that show isotopic fractionation, like methanogenesis. In paleosciences, 18 O: 16 O data from corals, foraminifera and ice cores are used as a proxy for temperature. It is defined as the deviation in "per mil" (‰, parts per thousand) between a sample and a standard: : \delta \ce{^{18}O} = \left( \frac{\left( \frac\ce{^{18}O}\ce{^{16}O} \right)_\mathrm{sample}}{\left( \frac\ce{^{18}O}\ce{^{16}O} \right)_\mathrm{standard}} - 1 \right) \times 1000 ‰ where the standard has a known isotopic composition, such as Vienna Standard Mean Ocean Water (VSMOW). The fractionation can arise from kinetic, equilibrium, or mass-independent fractionation.
In geochemistry, paleoclimatology and paleoceanography δ 18 O or delta-O-18 is a measure of the deviation in ratio of stable isotopes oxygen-18 ( 18 O) and oxygen-16 ( 16 O). It is commonly used as a measure of the temperature of precipitation, as a measure of groundwater/mineral interactions, and as an indicator of processes that show isotopic fractionation, like methanogenesis. In paleosciences, 18 O: 16 O data from corals, foraminifera and ice cores are used as a proxy for temperature. It is defined as the deviation in "per mil" (‰, parts per thousand) between a sample and a standard: \delta \ce{^{18}O} = \left( \frac{\left( \frac\ce{^{18}O}\ce{^{16}O} \right)_\mathrm{sample}}{\left( \frac\ce{^{18}O}\ce{^{16}O} \right)_\mathrm{standard}} - 1 \right) \times 1000 ‰ where the standard has a known isotopic composition, such as Vienna Standard Mean Ocean Water (VSMOW). The fractionation can arise from kinetic, equilibrium, or mass-independent fractionation.
Extrapolation of temperature
Based on the simplifying assumption that the signal can be attributed to temperature change alone, with the effects of salinity and ice volume change ignored, Epstein et al. (1953) estimated that a increase of 0.22‰ is equivalent to a cooling of 1 °C (or 1.8 °F). More precisely, Epstein et al. (1953) give a quadratic extrapolation for the temperature, as T = 16.5 - 4.3\mathrm{\delta} + 0.14 \mathrm{\delta^2} where T is the temperature in °C (based on a least-squares fit for a range of temperature values between 9 °C and 29 °C, with a standard deviation of ±0.6 °C, and δ is δ 18 O for a calcium carbonate sample). Recent research proposes that early ocean temperatures have been overestimated due to powerful carbonatization and silicification processes in the oceanic crust which consumed oxygen-18.
Paleoclimatology
and Raymo (2005)]]
See also
Global meteoric water line (GMWL), relation between and (also known as ) Isotopic signature Isotope analysis Isotope geochemistry Paleothermometer Stable isotope ratio
References
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