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Lomagundi-Jatuli Carbon Isotope Excursion

The Lomagundi-Jatuli Carbon Isotope Excursion or Lomagundi-Jatuli Event (LJE) was a carbon isotope excursion that occurred in the Paleoproterozoic between 2.3-2.1 Ga, possessing the largest magnitude and longest duration of positive δ13C values found in marine carbonate rocks.[1][2] The  δ13C values range from +5 to + 30‰.[3][4] Carbon isotope compositions in marine carbonates typically fluctuate around zero per mil (‰) through time.[5] To coincide with the LJE global δ13Ccarb levels, the amount of buried organic carbon would have needed to double or triple, and over millions of years.

Measuring δ13Ccarb values within marine carbonate rocks provides scientists with a window into the history of fluxes in the global carbon cycle over the course of Earth history.

In the context of the global carbon cycle, "flux" refers to the movement or flow of carbon between different reservoirs or components of the Earth system. This includes the atmosphere, oceans, terrestrial biosphere (plants and soil), and geosphere (rocks and sediments).[1] These carbon fluxes are driven by various processes, including photosynthesis (which removes CO2 from the atmosphere and incorporates it into plant biomass), respiration and decay (which return CO2 to the atmosphere), weathering of rocks (which can transfer carbon from the geosphere to the hydrosphere and atmosphere), and the dissolution and precipitation of carbonate minerals in the ocean[1]

Understanding these fluxes, especially within the LJE, is crucial for studying the global carbon cycle. They determine the concentration of carbon dioxide in the Earth's atmosphere, which in turn influences the planet's climate. While the LJE's high δ13Ccarb values were first thought to show a substantial local increase in organic carbon (forg) in the localities in which the elevated values were found, marine carbonate outcrops with similarly elevated values have since been found around the world, shifting consideration that this event reflects a global increase.[3]

Changes in carbon fluxes can lead to significant variations in atmospheric CO2 levels and thus have been a major focus of research, and debate, especially in the context of anthropogenic climate change.

Locations and duration

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Assuming this excursion is globally synchronous in its commencement and termination, the duration has been dated to range from a maximum of 249 ± 9 Myr (2306 ± 9 Ma to 2057 ± 1 Ma) to a minimum of 128 ± 9.4 Myr (2221 ± 5 Ma to 2106 ± 8 Ma).[3]

Region Formation Age (Ma) Method Max or Min Reference
Wyoming, USA Fletcher Park Rhyolite 1780 +- 6 ID-TIMS Min [3]
Wyoming, USA Keystone Quartz Diorite 1781 +- 7 ID-TIMS Min [3]
Uruguay Uruguayan Dyke swarm 1790 +- 5 ID-TIMS Min [3]
Pilbara Craton, Australia June Hill Volcanics 1795 +- 7 SIMS Min [3]
Gabon, Africa Francevillian Basin 2050 +- 30 ID-TIMS Max [3]
South Africa, Africa Rooihoogte Formation 2316+-7 TIMS (Re-Os) Max [3]
Russia, Kola Craton Polisarka Sedimentary Formation 2434 +-1.2 ID-TIMS Max [3]

Table 1: Lomagundi-Jatuli Event localities presenting similarly elevated δ13C values, formations of occurrence, dated age of formation, and procedural method of δ13C value analysis.

The extremely positive carbon isotope values having occurred during the LJE  can be seen on all continents, with the notable exception of Antarctica, having stratigraphic thickness ranging from several to tens of meters[3] The highly elevated δ13C values were first found in the Lomagundi Group in Zimbabwe and the Jatuli group in Fennoscandia at a time when the LJE was first hypothesized to have been a local event.[3]

Geographic and geological location Carbonate lithology δ13Ccarb variation (‰) Stratigraphic Thickness References
Africa
Lomagundi Group, Zimbabwe Dolostones 4.0 to +13.4 300 m [2]
Francevillian Series, Gabon Dolostones 2.6 to +6.3 600–2040 m [6]

[7]

Gumbu Group, South Africa Limestones 4.6  to +7.0 [8]
Duitschland Formation, Pretoria Group Dolostones −2.0 to +10.1 1000m [9]
Sengoma formation, Botswana Dolostones and limestones 7.6 to 9.2 167 m [10]
Silverton formations Pretoria Group, South Africa Dolostones and limestones 8.3 to +11.2 500 to 700 m [10]
Lucknow Formation, Elim Group Dolostones 8.7 to +10.4 200m [11]
Australia
Bubble Well Member, Juderina Formation, Nabberu Basin Dolostones 5.7 to +8.8 160m [12]

[13]

Asia
Jhamarkotra Formation, Aravalli Supergroup, India Dolostones 5.4 to +11.1 1500m [14]
South America
Cercadinho Formation, Minas Supergroup, Brazil Dolostones 3.3 to +5.4 317m [15]
Fecho do Funil Formation, Minas Supergroup, Brazil Dolostones 5.6 to +7.4 38-50m [15]
Rio Itapicuru Greenstone Belt, Brazil Dolostones 5.5 to +9.0 20-30m [13]
Ipueira-Medrado, Itabuna-Salvador-Curaçá orogen, Brazil Dolostones 2.2 to +6.9 30-40m [13]
Paso Severino Fm., Río de la Plata craton, Uruguay Dolostones −5.6 to +11.6 2100-2700 [13]
North America
Gordon Lake Formation, Canada Dolostones -1 to +8.2 300-700 meters [16]
Kona Dolomite, United States Dolostones 1.9 to +9.5 870 meters [16]
Nash Fork Formation, United States Dolostones 0.2 to 28.2 1700m [4]
Slaughterhouse Formation, United States Dolostones 5.6 to +16.6 500m [4]

Table 2: Carbonate lithology within global formations, including associated δ13Ccarb variation (‰) values, and stratigraphic thicknesses of each.

Methods

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Scientists choose which geochronology method is best suited for the types of rocks and sediments they work with. Attempting to find the age of marine carbonate rocks is not without challenge, especially in the use of uranium (U) and lead (Pb). These rocks do not have an initially high composition of uranium, and contain too much lead, both of which can further undergo modification over geologic time (diagenetic overprinting).[17]

Isotope Dilution-Thermal Ionization Mass Spectrometry (ID-TIMS) has been utilized for the analysis of marine carbonate rock δ13 values due its precision <1‰ (single analysis or weighted mean dates) in evaluating 206Pb/238U dates.[17] This method involves a two-step process: isotope dilution, where a known amount of isotopically enriched tracer is mixed with the sample to quantify the concentration of elements, and thermal ionization mass spectrometry, where the sample is ionized at high temperatures to measure the isotopic ratios of elements.[17]

Secondary ion mass spectrometry (SIMS) is a form of desorption mass spectrometry, which is used to analyze the composition of solid surfaces and thin films by sputtering the surface of the specimen with a focused primary ion beam and collecting and analyzing ejected secondary ions.[18] The principle behind SIMS is straightforward but involves sophisticated instrumentation and techniques to achieve detailed surface compositional analysis at the micro to nano scale.[18]

Thermal Ionization Mass Spectrometry (TIMS) is a highly precise and sensitive analytical technique used primarily for isotopic analysis (Rhenium-Osmium) and the determination of elemental concentrations in samples.[19] TIMS operates on the principle of thermal ionization, where a sample is vaporized and ionized in a high-temperature environment, allowing for the separation and measurement of isotopes based on their mass-to-charge ratios.[19]

The Re-Os (Rhenium-Osmium) geochronology method is based on the decay of 187Re to 187Os, which occurs over a long half-life, making it suitable for dating geological samples that range in age from millions to billions of years. [19] This method is particularly useful for dating organic-rich rocks, such as black shales, and relies on the principle that the ratio of these isotopes in a sample will change over time due to the radioactive decay of 187Re to 187Os. By measuring the present-day isotope ratios and knowing the decay rate of 187Re, scientists can calculate the age of the sample. [19]

One of the challenges in Re-Os geochronology is dealing with the error correlations that arise from measuring these isotopes, particularly because 188Os, which is used in the denominator of the isotope ratio calculations, is associated with larger mass spectrometer uncertainties.[19] This issue can lead to strong error correlations and potentially obscure geologically significant trends.

Genesis of the LJE

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Synchronous, global-scale disturbance

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The global view concludes that during the Lomagundi carbon isotope excursion carbonates were deposited world-wide with large amounts of 13C enrichment.[20] The hypothesis that the LJE is related to the Great Oxidation Event (GOE), with the LJE causing a large deviation in the global carbon reservoir, which in turn led to disequilibrium of the carbon cycle and released oxygen.[21]

To explain this global 13C enrichment, the oxidation of siderite (FeCO3, with other Fe(2+) carbonate minerals), was proposed as a hypothesis because it produces 4 times the amount of CO2 than it consumes O2.[22] The oxidation of siderite was the driver for the carbon needed in burial and further oxidation, as well as the accumulation of O2, making the length of the LJE dependent on the size of the Archean siderite reservoir.[22]

Another hypothesis to explain the global nature of the LJE, is large  tectonic change leading to increased degassing of volcanic CO2, which could have increased deposition of carbonates and organic matter, due to higher weathering rates and nutrients to the ocean.[23] Similar to a tectonic change, the formation of subaerial continents or global glaciations could have also enhanced volcanic CO2 leading to the same outcome of CO2 and O2 in carbonates and atmosphere.[23] Supporting this, there is evidence for the first large continental plates around 2.2-2.1 Gyr experiencing rifting and a global orogeny.[24] During this time frame there was an increase in seawater 87Sr/ 86Sr, which indicates there was higher levels of continental erosion.[24] To reinforce the effect of high 87Sr/86Sr, the first known glaciations occurred during 2.2-2.1 Gyr as well which favours weathering rates by lowering sea level.[24]

Localized, Facies driven process

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This hypothesis acknowledges that there is a global change to the carbon cycle, and agrees that it was a globally synchronous event, but with the idea that different facies environments drive the high C-isotope values. This meaning that the values of d13C and changes in the values are because of processes in individual basins, depending on where the locality is along a carbonate platform/slope.[25] Using 13C carbonate data from locations worldwide and the stratigraphic descriptions, the values can be organized into open marine, nearshore marine-inner shelf and intertidal-coastal-sabka, with a noticed correlation between facies and 13C carbonate values.[25] For open marine the mean 13C carbonate value was +1.5 ± 2.4‰, +6.2 ± 2.0‰ for inner-shelf, and +8.1 ± 3.8‰ for intertidal settings.[25] Using this hypothesis, the extremely positive d13C values can be explained by changes in local dissolved inorganic carbon (DIC) pools, influenced in individual basins, not representative of word-wide change in ocean DIC.[25]

Localized, Diagenesis or Methanogenesis

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For a global carbon isotope excursion, sedimentary organic carbon (shales) tend to show a trend as well, as an excursion would affect the d13C value of the biosphere and therefore sedimentary organic matter.[20] Between 2.60 and 1.60 Ga there is no trend within organic carbon. Fluctuations in d13C can be linked to isotopic alteration from breakdown of organic matter due to diagenesis and metamorphism.[26]

A process in sediment columns that can contribute to carbonates with high levels of 13C, methanogenesis, could have caused carbonates enriched in 13C, creating an explanation for the d13C values reaching +28‰. To explain the LJE, a deeper methanic zone in oceans during the start of ocean oxygenation (the GOE) would push pore water DIC to higher d13C values. The carbonates forming at this time would record d13C enrichment.[1]

References

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  1. ^ a b c d Hodgskiss, Malcolm S.W.; Crockford, Peter W.; Turchyn, Alexandra V. (31 May 2023). "Deconstructing the Lomagundi-Jatuli Carbon Isotope Excursion". Annual Review of Earth and Planetary Sciences. 51 (1): 301–330. Bibcode:2023AREPS..51..301H. doi:10.1146/annurev-earth-031621-071250.
  2. ^ a b Schidlowski, Manfred; Eichmann, Rudolf; Junge, Christian E. (April 1976). "Carbon isotope geochemistry of the Precambrian Lomagundi carbonate province, Rhodesia". Geochimica et Cosmochimica Acta. 40 (4): 449–455. Bibcode:1976GeCoA..40..449S. doi:10.1016/0016-7037(76)90010-7.
  3. ^ a b c d e f g h i j k l Martin, Adam P.; Condon, Daniel J.; Prave, Anthony R.; Lepland, Aivo (December 2013). "A review of temporal constraints for the Palaeoproterozoic large, positive carbonate carbon isotope excursion (the Lomagundi–Jatuli Event)". Earth-Science Reviews. 127: 242–261. Bibcode:2013ESRv..127..242M. doi:10.1016/j.earscirev.2013.10.006.
  4. ^ a b c Bekker, A (10 February 2003). "Chemostratigraphy of Paleoproterozoic carbonate successions of the Wyoming Craton: tectonic forcing of biogeochemical change?". Precambrian Research. 120 (3–4): 279–325. Bibcode:2003PreR..120..279B. doi:10.1016/S0301-9268(02)00164-X.
  5. ^ Krissansen-Totton, J.; Buick, R.; Catling, D. C. (April 2015). "A statistical analysis of the carbon isotope record from the Archean to Phanerozoic and implications for the rise of oxygen". American Journal of Science. 315 (4): 275–316. Bibcode:2015AmJS..315..275K. doi:10.2475/04.2015.01.
  6. ^ Gauthier-Lafaye, F.; Weber, F. (December 1989). "The Francevillian (Lower Proterozoic) uranium ore deposits of Gabon". Economic Geology. 84 (8): 2267–2285. Bibcode:1989EcGeo..84.2267G. doi:10.2113/gsecongeo.84.8.2267.
  7. ^ Lecomte, Andreï; Michels, Raymond; Cathelineau, Michel; Morlot, Christophe; Brouand, Marc; Flotté, Nicolas (August 2020). "Uranium deposits of Franceville basin (Gabon): Role of organic matter and oil cracking on uranium mineralization". Ore Geology Reviews. 123: 103579. Bibcode:2020OGRv..12303579L. doi:10.1016/j.oregeorev.2020.103579.
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  12. ^ Lindsay, John F; Brasier, Martin D (February 2002). "Did global tectonics drive early biosphere evolution? Carbon isotope record from 2.6 to 1.9 Ga carbonates of Western Australian basins". Precambrian Research. 114 (1–2): 1–34. Bibcode:2002PreR..114....1L. doi:10.1016/S0301-9268(01)00219-4.
  13. ^ a b c d Maheshwari, Anil; Sial, Alcides N.; Gaucher, Claudio; Bossi, Jorge; Bekker, Andrey; Ferreira, Valderez P.; Romano, Antonio W. (October 2010). "Global nature of the Paleoproterozoic Lomagundi carbon isotope excursion: A review of occurrences in Brazil, India, and Uruguay". Precambrian Research. 182 (4): 274–299. Bibcode:2010PreR..182..274M. doi:10.1016/j.precamres.2010.06.017.
  14. ^ Sreenivas et al. (1999, 2001)[full citation needed]
  15. ^ a b Bekker, A. (December 2003). "Chemostratigraphy of Carbonates from the Minas Supergroup, Quadrilatero Ferryifero (Iron Quadrangle), Brazil: A Stratigraphic Record of Early Proterozoic Atmospheric, Biogeochemical and Climactic Change". American Journal of Science. 303 (10): 865–904. Bibcode:2003AmJS..303..865B. doi:10.2475/ajs.303.10.865.
  16. ^ a b Bekker, A.; Karhu, J.A.; Kaufman, A.J. (20 July 2006). "Carbon isotope record for the onset of the Lomagundi carbon isotope excursion in the Great Lakes area, North America". Precambrian Research. 148 (1–2): 145–180. Bibcode:2006PreR..148..145B. doi:10.1016/j.precamres.2006.03.008.
  17. ^ a b c Ovtcharova, Maria; Samankassou, Elias; Guillong, Marcel; Linnemann, Ulf; Hofmann, Mandy; Zieger, Johannes; Peyrotty, Giovan; Vennemann, Torsten; Messori, Fabio; Merino-Tomé, Oscar; Nuriel, Perach; Kouzmanov, Kalin; Müller, Inigo (2023). Can we date marine carbonates at high(er) precision with U-Pb ID-TIMS method?. Goldschmidt 2023. doi:10.7185/gold2023.18321.
  18. ^ a b Zhao, Wenhao; Li, Qiuli; Liu, Yu; Tang, Guoqiang; Ling, Xiaoxiao; Li, Jiao; Li, Xianhua (February 2022). "Long-Term Reproducibility of SIMS Zircon U-Pb Geochronology". Journal of Earth Science. 33 (1): 17–24. Bibcode:2022JEaSc..33...17Z. doi:10.1007/s12583-021-1549-1.
  19. ^ a b c d e 武鲁亚; 金之钧; 储著银; 刘可禹 (2021). "富有机质样品Re-Os同位素定年实验方法" [Re-Os isotopic dating procedures for organic-rich samples]. 石油实验地质 [Petroleum Geology & Experiment] (in Chinese). 43 (3): 513–523. doi:10.11781/sysydz202103513.
  20. ^ a b Karhu, Juha A.; Holland, Heinrich D. (1996). "Carbon isotopes and the rise of atmospheric oxygen". Geology. 24 (10): 867. Bibcode:1996Geo....24..867K. doi:10.1130/0091-7613(1996)024<0867:CIATRO>2.3.CO;2.
  21. ^ Melezhik, V. A. (1994). "A World-Wide 2.2-2.0 Ga-Old Positive δ13Ccarb Anomaly as a Phenomenon in Relation to the Earth's Major Palaeoenviromental Changes". Mineralogical Magazine. 58A (2): 593–594. Bibcode:1994MinM...58..593M. doi:10.1180/minmag.1994.58a.2.45.
  22. ^ a b Bachan, Aviv; Kump, Lee R. (26 May 2015). "The rise of oxygen and siderite oxidation during the Lomagundi Event". Proceedings of the National Academy of Sciences. 112 (21): 6562–6567. Bibcode:2015PNAS..112.6562B. doi:10.1073/pnas.1422319112. PMC 4450422. PMID 25964326.
  23. ^ a b Eguchi, James; Seales, Johnny; Dasgupta, Rajdeep (January 2020). "Great Oxidation and Lomagundi events linked by deep cycling and enhanced degassing of carbon". Nature Geoscience. 13 (1): 71–76. Bibcode:2020NatGe..13...71E. doi:10.1038/s41561-019-0492-6. PMC 6894402. PMID 31807138.
  24. ^ a b c Marais, David J. Des; Strauss, Harald; Summons, Roger E.; Hayes, J. M. (October 1992). "Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment". Nature. 359 (6396): 605–609. Bibcode:1992Natur.359..605M. doi:10.1038/359605a0. PMID 11536507.
  25. ^ a b c d Prave, A.R.; Kirsimäe, K.; Lepland, A.; Fallick, A.E.; Kreitsmann, T.; Deines, Yu.E.; Romashkin, A.E.; Rychanchik, D.V.; Medvedev, P.V.; Moussavou, M.; Bakakas, K.; Hodgskiss, M.S.W. (January 2022). "The grandest of them all: the Lomagundi–Jatuli Event and Earth's oxygenation". Journal of the Geological Society. 179 (1). Bibcode:2022JGSoc.179...36P. doi:10.1144/jgs2021-036. hdl:10023/23664.
  26. ^ Schopf, J. William (1983). Earth's Earliest Biosphere: Its Origin and Evolution. Princeton University Press. ISBN 978-0-691-02375-5.[page needed]
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