Biology:Δ13C

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Short description: Measure of relative carbon-13 concentration in a sample


Foraminifera samples

In geochemistry, paleoclimatology, and paleoceanography δ13C (pronounced "delta c thirteen") is an isotopic signature, a measure of the ratio of the two stable isotopes of carbon—13C and 12C—reported in parts per thousand (per mil, ‰).[1] The measure is also widely used in archaeology for the reconstruction of past diets, particularly to see if marine foods or certain types of plants were consumed.[2]

The definition is, in per mille:

[math]\displaystyle{ \delta \ce{^{13}C} = \left( \frac{(\ce{^{13}C}/\ce{^{12}C})_\mathrm{sample}}{(\ce{^{13}C}/\ce{^{12}C})_\mathrm{standard}} - 1 \right) \times 1000 }[/math]

where the standard is an established reference material.

δ13C varies in time as a function of productivity, the signature of the inorganic source, organic carbon burial, and vegetation type. Biological processes preferentially take up the lower mass isotope through kinetic fractionation. However some abiotic processes do the same. For example, methane from hydrothermal vents can be depleted by up to 50%.[3]

Reference standard

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 Peedee Formation in South Carolina. This material had an anomalously high 13C:12C ratio (0.0112372[4]), and was established as δ13C value of zero. Since the original PDB specimen is no longer available, its 13C:12C ratio can be back-calculated from a widely measured carbonate standard NBS-19, which has a δ13C value of +1.95‰.[5] The 13C:12C ratio of NBS-19 was reported as [math]\displaystyle{ 0.011078/0.988922=0.011202 }[/math].[6] Therefore, one could calculate the 13C:12C ratio of PDB derived from NBS-19 as [math]\displaystyle{ 0.011202 / (1.95/1000 +1)= 0.011202/1.00195=0.01118 }[/math]. Note that this value differs from the widely used PDB 13C:12C ratio of 0.0112372 used in isotope forensics[7] and environmental scientists;[8] this discrepancy was previously attributed by a wikipedia author to a sign error in the interconversion between standards, but no citation was provided. Use of the PDB standard gives most natural material a negative δ13C.[9] A material with a ratio of 0.010743 for example would have a δ13C value of −44‰ from [math]\displaystyle{ (0.010743 \div 0.01124 - 1) \times 1000 }[/math]. 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 calibrated to the same ratio, including one known as VPDB (for "Vienna PDB"), have replaced the original.[10] The 13C:12C ratio for VPDB, which the International Atomic Energy Agency (IAEA) defines as δ13C value of zero is 0.01123720.[11]

Causes of δ13C variations

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 Paleocene–Eocene Thermal Maximum.[12]

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. Therefore, an increase in δ13C in marine fossils is indicative of an increase in the abundance of vegetation.[citation needed]

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.

Geologic significance of δ13C excursions

C3 and C4 plants have different signatures, allowing the abundance of C4 grasses to be detected through time in the δ13C record.[13] Whereas C4 plants have a δ13C of −16 to −10‰, C3 plants have a δ13C of −33 to −24‰.[14]

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

Positive δ13C excursions are interpreted as an increase in burial of organic carbon in sedimentary rocks following either a spike in primary productivity, a drop in decomposition under anoxic ocean conditions or both.[15]

The evolution of large land plants in the late Devonian led to increased organic carbon burial and consequently a rise in δ13C.[16]

Other important applications of δ13C involves understanding its signatures from soft sediments especially in lacustrine environments. This depends on the system from which it is extracted (open system, closed system, etc.). Temporal variations in δ13C in organic matter are influenced by diverse internal and external processes:[17]

  1. Changes in the Dominant Source of Dissolved Inorganic Carbon: In stratified lakes, the accumulation of 13C-depleted carbon in deep water is common as sinking and degrading phytoplankton cells contribute to this pool. Recirculating this water to the surface can lead to a significant decrease in δ13C. Prolonged stratification enriches the dissolved inorganic carbon (DIC) pool in the epilimnion with 13C. Long-term variations in factors affecting upwelling intensity or depth, such as windiness, water temperature, or salinity-related stratification, manifest as shifts between more negative and positive δ13C values.
  2. Changes in Productivity/Eutrophication: Increased productivity accelerates the transfer of organic matter with negative δ13C values to the hypolimnion, affecting the δ13C of residual epilimnetic DIC. This impact, combined with mixing effects, results in variations in the δ13C signal.
  3. Changes in Metabolic Pathways for Carbon Fixation: Major changes in lake alkalinity influence benthic and planktonic primary production. Shifts in the dominant source of DIC for photosynthesis, driven by pH changes, can lead to trends toward more positive δ13C, particularly in lakes dominated by autochthonous organic matter and exhibiting evidence of high alkalinity.
  4. Changes in Availability of Dissolved CO2: Cool water can dissolve higher concentrations of CO2 than warmer water, affecting δ13C in organic matter during cooling events. Changes in atmospheric CO2 concentrations also influence δ13C, with lower pCO2 during glacial periods causing isotopic discrimination in plants using dissolved CO2.
  5. Changes in Dominant Vegetation Within the Watershed: Shifts in watershed vegetation, especially transitions between C3 and C4 photosynthetic pathways, significantly alter the carbon isotopic composition in lake sediments. These changes can be indicative of broader paleoclimatic shifts.
  6. Diagenetic Trends: Diagenetic processes, such as the loss of reactive components like amino acids, result in sustained shifts in δ13C in organic matter. Marsh sediments, rich in carbon, exhibit shifts towards more negative bulk organic matter. These diagenetic trends should be considered when interpreting isotopic changes accompanying major Total Organic Carbon (TOC) changes or methanogenesis.

Understanding these processes is crucial for interpreting δ13C variations in lake sediments and reconstructing paleoenvironmental conditions.

Major excursion events

  • Lomagundi-Jatuli event (2,300–2,080 Ma) Paleoproterozoic - Positive excursion
  • Steptoean positive carbon isotope excursion (494.6-492 Ma) Paleozoic - Positive excursion

See also

References

  1. ↑ Libes, Susan M. (1992). Introduction to Marine Biogeochemistry, 1st edition.. New York: Wiley. 
  2. ↑ Schwarcz, Henry P.; Schoeninger, Margaret J. (1991). "Stable isotope analyses in human nutritional ecology". American Journal of Physical Anthropology 34 (S13): 283–321. doi:10.1002/ajpa.1330340613. 
  3. ↑ McDermott, J.M., Seewald, J.S., German, C.R. and Sylva, S.P., 2015. Pathways for abiotic organic synthesis at submarine hydrothermal fields. Proceedings of the National Academy of Sciences, 112(25), pp.7668–7672.
  4. ↑ Craig, Harmon (1957-01-01). "Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide" (in en). Geochimica et Cosmochimica Acta 12 (1): 133–149. doi:10.1016/0016-7037(57)90024-8. ISSN 0016-7037. Bibcode: 1957GeCoA..12..133C. https://dx.doi.org/10.1016/0016-7037%2857%2990024-8. 
  5. ↑ Brand, Willi A.; Coplen, Tyler B.; Vogl, Jochen; Rosner, Martin; Prohaska, Thomas (2014-03-20). "Assessment of international reference materials for isotope-ratio analysis (IUPAC Technical Report)" (in en). Pure and Applied Chemistry 86 (3): 425–467. doi:10.1515/pac-2013-1023. ISSN 1365-3075. 
  6. ↑ Meija, Juris; Coplen, Tyler B.; Berglund, Michael; Brand, Willi A.; De Bièvre, Paul; Gröning, Manfred; Holden, Norman E.; Irrgeher, Johanna et al. (2016-01-01). "Isotopic compositions of the elements 2013 (IUPAC Technical Report)" (in en). Pure and Applied Chemistry 88 (3): 293–306. doi:10.1515/pac-2015-0503. ISSN 1365-3075. 
  7. ↑ Meier-Augenstein, Wolfram (28 September 2017). Stable isotope forensics : methods and forensic applications of stable isotope analysis (Second ed.). Hoboken, NJ. ISBN 978-1-119-08022-0. OCLC 975998493. https://www.worldcat.org/oclc/975998493. 
  8. ↑ Michener, Robert, ed (2007-07-14) (in en). Stable Isotopes in Ecology and Environmental Science. Oxford, UK: Blackwell Publishing Ltd. doi:10.1002/9780470691854. ISBN 978-0-470-69185-4. 
  9. ↑ Overview of Stable Isotope Research. The Stable Isotope/Soil Biology Laboratory of the University of Georgia Institute of Ecology.
  10. ↑ Miller & Wheeler, Biological Oceanography, p. 186.
  11. ↑ "Reference and intercomparison materials for stable isotopes of light elements". International Atomic Energy Agency. 1995. https://www-pub.iaea.org/MTCD/publications/PDF/te_825_prn.pdf. 
  12. ↑ 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. Bibcode: 2008Geo....36..315P. 
  13. ↑ Retallack, G.J. (2001). "Cenozoic Expansion of Grasslands and Climatic Cooling". The Journal of Geology 109 (4): 407–426. doi:10.1086/320791. Bibcode: 2001JG....109..407R. 
  14. ↑ O'Leary, M. H. (1988). "Carbon Isotopes in Photosynthesis". BioScience 38 (5): 328–336. doi:10.2307/1310735. 
  15. ↑ Canfield, Donald E.; Ngombi-Pemba, Lauriss; Hammarlund, Emma U. (15 October 2013). "Oxygen dynamics in the aftermath of the Great Oxidation of Earth's atmosphere". Proceedings of the National Academy of Sciences of the United States of America 110 (42): 16736–16741. doi:10.1073/pnas.1315570110. PMID 24082125. Bibcode: 2013PNAS..11016736C. 
  16. ↑ Joachimsk, M.M.; Buggisch, W.. "THE LATE DEVONIAN MASS EXTINCTION – IMPACT OR EARTH-BOUND EVENT?". https://www.lpi.usra.edu/meetings/impact2000/pdf/3072.pdf. 
  17. ↑ Cohen, Andrew S. (2003-05-08), "The Geological Evolution of Lake Basins", Paleolimnology (Oxford University Press), http://dx.doi.org/10.1093/oso/9780195133530.003.0006, retrieved 2023-12-19 

Further reading

  • Miller, Charles B.; Patricia A. Miller (2012). Biological Oceanography (2nd ed.). Oxford: John Wiley & Sons. ISBN 978-1-4443-3301-5. 
  • Mook, W. G., & Tan, F. C. (1991). Stable carbon isotopes in rivers and estuaries. Biogeochemistry of major world rivers, 42, 245–264.



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