CO2 drawdown and cooling at the onset of the Great Oxidation Event recorded in 2.45 Ga paleoweathering crust
Introduction
Atmospheric carbon dioxide concentration is an important factor in control of the Earth's surface temperatures and chemical processes (Broecker, 2018; Dessert et al., 2003; Kasting, 1987). High atmospheric CO2 concentration lowers the pH of the meteoric precipitation and by greenhouse effect increases air temperature, both leading to intensified weathering (Kanzaki and Murakami, 2018a, Kanzaki and Murakami, 2018b). Chemical weathering itself has an important impact on the evolution of the Earth's surface environments by regulating atmospheric CO2 through a negative feedback mechanism (Dessert et al., 2003; Kump et al., 2000; Walker et al., 1981) and release of bio-limiting nutrients into the ocean (Hao et al., 2017; Filippelli, 2011). On a long term the level of atmospheric CO2 acts as a climate thermostat, keeping Earth environments in the range where water is mostly present in liquid form and surface temperatures are favorable for life (Kump et al., 2000).
Elevated CO2 levels (or other greenhouse gases such as methane) and the enhanced greenhouse effect have played an important role in early Earth history by compensating for the faint-young-Sun effect (Feulner, 2012; Haqq-Misra et al., 2008; Sagan and Mullen, 1972; Walker, 1982). Models of stellar evolution have led to interpretations that the Sun was 20–25% less luminous in the Archean as compared to modern (Wolf and Toon, 2014; Ozaki et al., 2018), and became 6% less luminous during the Proterozoic (Donnadieu et al., 2004) suggesting that Earth's surface would have been in a frozen state under present atmospheric composition (von Paris et al., 2008). However, there is ample evidence that the surface and ocean temperatures of the Archean were the same, or perhaps even warmer as compared to present day (Knauth and Lowe, 2003; Robert and Chaussidon, 2006; Hren et al., 2009; Blake et al., 2010). Although the need for high CO2 levels in the Archean has been debated (e.g., Rosing et al., 2010) to overcome the lower solar luminosity, it has been suggested that CO2 concentrations had to have been at least 100–1000 times that of the present level (Kasting, 1993) in order to maintain the temperatures above the water freezing point.
Direct estimates of the Archean-Proterozoic pCO2 levels are mainly limited to weathering crusts (paleosols) that are capable of recording the paleoatmospheric composition (e.g., Rye and Holland, 1998; Holland, 2006, Holland, 2009; Sheldon, 2006; Sheldon and Tabor, 2009; Kanzaki and Murakami, 2015, Kanzaki and Murakami, 2018a, Kanzaki and Murakami, 2018b). Paleoatmospheric pCO2 estimates in the Archean and Proterozoic (Kanzaki and Murakami, 2015, Kanzaki and Murakami, 2018b; Sheldon, 2006) generally agree with, or are lower than, modeled atmospheric composition (Kasting, 2010). However, Precambrian paleosols have a relatively low preservation potential and are commonly preserved as denudation surfaces rather than as complete weathering profiles (Peters and Gaines, 2012). As a result, the pCO2 estimates obtained from paleosols can vary by more than an order of magnitude (Sheldon, 2006).
The focus of this paper is on pCO2 modeling at the beginning of Proterozoic using a recently described c. 2.50–2.44 Ga paleoweathering crust formed on basaltic parent rock of the Kuksha Volcanic Formation in the Imandra-Varzuga Greenstone Belt of the Kola Peninsula (Soomer et al., 2019). This weathering crust was formed during a critical time period in Earth history at the Archean-to-Proterozoic transition when sequential global events paved the way for the establishment of the aerobic atmosphere of the modern-type Earth (Bekker et al., 2004, Melezhik et al., 2005; Holland, 2006). The key event was the shift from anoxic to oxic surface environmental conditions, termed the “Great Oxidation Event” (GOE; Holland, 2006; Lyons et al., 2014). The age of the GOE is bracketed between ca. 2.4 and 2.3 Ga (e.g., Bekker et al., 2004) but Gumsley et al. (2017) have recently suggested onset of the GOE between ca. 2.46 and 2.426 Ga. The GOE was also associated with the initiation of the first low-latitude Huronian glaciations between 2.45 and 2.22 Ga (Young, 2019; Evans et al., 1997; Bekker et al., 2004). The atmospheric composition and environmental conditions before and after the GOE are not well-constrained and are thus still highly debated (Bekker et al., 2004; Canfield, 2005; Kaufman et al., 2007; Lyons et al., 2014; Planavsky et al., 2014; Zahnle and Catling, 2014). The Kuksha paleoweathering crust, immediately predating and/or overlapping in age with the GOE thus, could provide important insights into the atmospheric-climatic conditions of this key time interval. The aim of this paper is first, to estimate the atmospheric pCO2 concentrations at the time of the weathering and, secondly, to determine environmental conditions (mean annual precipitation – MAP; mean annual temperature – MAT) during the formation of the Kuksha weathering crust at the beginning of the GOE using geochemical climofunctions.
Section snippets
Geological setting
The Kuksha paleoweathering crust developed on the basalt of the Kuksha Volcanic Formation at the Imandra-Varzuga Greenstone Belt in the northeastern part of the Fennoscandian Shield (Melezhik, 2013). The Imandra-Varzuga Greenstone Belt is a 350 km-long and up to 10–50 km-wide zone in the southeastern part of the ca. 800 km-long system of supracrustal belts extending from northern Norway through northern Finland to northwestern Russia (Fig. 1) (Chashchin et al., 2008). The Kuksha succession
Materials and methods
Altogether 32 samples for petrological, whole rock mineralogical and geochemical analysis were collected from drillcore 1A drilled during the International Fennoscandian Arctic Russia - Drilling Early Earth Project (FAR-DEEP) (Melezhik, 2013). The whole-rock petrographic, mineralogical and geochemical data provided by Soomer et al. (2019) described a ca. 14 m-thick profile of the 1A core from unweathered parent rock (194–200 m depth) to the paleoweathering surface (at ~185 m depth).
The details
Molecular indices of paleoclimate conditions
When one applies MAT and MAP climofunctions to ancient paleosols it is important to remember that these should be used with great caution. Firstly, most of the climofunctions used in paleosol studies were developed for application to specific soil types and using specific horizons (e.g., Bw or Bt horizons). Thus, these proxies may not be directly applicable to other, unspecified soil types and horizons. Secondly, classification and characterization of deep-time paleosols is complicated because
Conclusions
The c. ~2.45 Ga old Kuksha paleoweathering crust developed at the time of the breakup of the Kenorland (Superia) supercontinent and followed by series of large-scale Huronian glaciation(−s) largely coeval with the GOE. The climofunctions, although not generally directly applied to this age paleosols, suggest that Kuksha paleoweathering crust formed under cool temperate climate with mean annual temperatures ~12–13 °C and mean annual precipitation between 700 and 1100 mm yr−1, roughly
Declaration of competing interest
The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.
Acknowledgements
This study was supported by Estonian Research Council grants PUT1511, PRG447 and the Estonian Centre of Analytical Chemistry.
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2022, Precambrian ResearchCitation Excerpt :Kanzaki and Murakami (2015) estimated an even higher paleoatmospheric pCO2 level in the range of 160 – 490 PIAL, at ca. 2460 Ma. Somelar et al. (2020), however, estimated a much lower level of pCO2 at 1–10 PIAL in the Kuksha paleosol, which was interpreted to record CO2 drawdown prior to the onset of the Huronian glaciation (Young, 2019). Further, this low pCO2 estimate agrees with the mass-balance modeling studies of Sheldon (2006), Mitchell and Sheldon (2010), and Medaris et al. (2017), suggesting a drop in CO2 levels below 20 PIAL between 2500 and 1800 Ma and further down to only 4–6 PIAL between 1800 and 1000 Ma.