δ13C

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Foraminifera samples.

In geochemistry, paleoclimatology and paleoceanography δ13C is an isotopic signature, a measure of the ratio of stable isotopes 13C : 12C, reported in parts per thousand (per mil, ‰).[1]

The definition is, in per mil:

\mathrm{\delta ^{13}C} = \Biggl( \mathrm{\frac{\bigl( \frac{^{13}C}{^{12}C} \bigr)_{sample}}{\bigl( \frac{^{13}C}{^{12}C} \bigr)_{standard}}} -1 \Biggr) \ * 1000\ ^{o}\!/\!_{oo}

where the standard is an established reference material.

δ13C varies in time as a function of productivity, organic carbon burial and vegetation type.


Reference standard[edit]

The standard established for carbon-13 work was the Pee Dee Belemnite (PDB) and was based on a Cretaceous marine fossil, Belemnitella americana, which was from the Pee Dee Formation in South Carolina. This material had an anomalously high 13C:12C ratio (0.0112372), and was established as δ13C value of zero. Use of this standard gives most natural material a negative δ13C.[2] The standards are used for verifying the accuracy of mass spectroscopy; as isotope studies became more common, the demand for the standard exhausted the supply. Other standards, including one known as VPDB (for "Vienna PDB") have replaced the original.[3]


What affects δ13C?[edit]

Methane has a very light δ13C signature: biogenic methane of −60‰ thermogenic methane −40‰. The release of large amounts of methane clathrate can impact on global δ13C values, as at the PETM.[4]

More commonly, the ratio is affected by variations in primary productivity and organic burial. Organisms preferentially take up light 12C, and have a δ13C signature of about −25‰, depending on their metabolic pathway.

An increase in primary productivity causes a corresponding rise in δ13C values as more 12C is locked up in plants. This signal is also a function of the amount of carbon burial; when organic carbon is buried, more 12C is locked out of the system in sediments than the background ratio (because organic carbon is lighter).

Geologically significant δ13C excursions[edit]

C3 and C4 plants have different signatures, allowing the importance of C4 grasses to be detected through time in the δ13C record.[5] Whereas C4 plants have a δ13C of −16 to −10 ‰, C3 plants have a δ13C of −33 to −24‰.[6]

Mass extinctions are often marked by a negative δ13C anomaly thought to represent a decrease in primary productivity and release of plant-based carbon

The evolution of large land plants in the late Devonian led to increased organic carbon burial and consequently a rise in δ13C.[7]

See also[edit]

Notes[edit]

  1. ^ Libes, Susan M. (1992). Introduction to Marine Biogeochemistry, 1st edition. New York: Wiley. 
  2. ^ http://www.uga.edu/sisbl/stable.html#calib Overview of Stable Isotope Research – The Stable Isotope/Soil Biology Laboratory of the University of Georgia Institute of Ecology
  3. ^ Miller & Wheeler, Biological Oceanography, p. 186.
  4. ^ Panchuk, K.; Ridgwell, A.; Kump, L.R. (2008). "Sedimentary response to Paleocene-Eocene Thermal Maximum carbon release: A model-data comparison". Geology 36 (4): 315–318. doi:10.1130/G24474A.1. 
  5. ^ Retallack, G.J. (2001). "Cenozoic Expansion of Grasslands and Climatic Cooling". The Journal of Geology 109 (4): 407–426. Bibcode:2001JG....109..407R. doi:10.1086/320791. 
  6. ^ O'Leary, M. H. (1988). "Carbon Isotopes in Photosynthesis". BioScience 38 (5): 328–336. doi:10.2307/1310735. JSTOR 1310735.  edit
  7. ^ http://www.lpi.usra.edu/meetings/impact2000/pdf/3072.pdf

References[edit]

  • Miller, Charles B.; Patricia A. Miller (2012) [2003]. Biological Oceanography (2nd ed.). Oxford: John Wiley & Sons. ISBN 978-1-4443-3301-5.