Abstract
A shift toward a higher oxygen level in both ocean and atmosphere systems during the late Ediacaran to the early Cambrian has been suggested from multiple indirect proxies. However, the mechanism and magnitude of this oxidation remain unclear. To solve this issue, we measured carbon isotopes in both carbonate and organic matter as well as their trace element compositions for an Ediacaran-Cambrian sequence in the Lower Yangtze basin, South China. The δ13Corg and δ13Ccarb excursions of this sequence are coupled and can be compared with contemporaneous global carbon isotope curves. A 2‰ rise in Δ13Ccarb-org occurred from the late Ediacaran to the early Cambrian, suggesting a substantial increase in atmospheric oxygen level from 16% to 30% of the present atmospheric level (PAL). Furthermore, the distribution pattern of rare earth elements and the concentrations of water-insoluble elements in the carbonates indicate a sudden enhancement in chemical weathering of the continental crust during the early Cambrian, which may be a trigger for the rise of atmospheric O2 level. Both the supply of a large amount of nutrients due to the enhanced continental weathering and the contemporary increase of atmospheric oxygen concentrations may have promoted the appearance of large metazoans in the early Cambrian.
Similar content being viewed by others
Introduction
Fluctuations in atmospheric oxygen (O2) and seawater redox play a fundamental role in driving biological evolution throughout the geological time1,2. Atmospheric and seawater dissolved oxygen levels increase during the growth of organic carbon reservoirs because the CO2 flux into the atmosphere–ocean system is balanced by carbon burial with the release of oxygen3,4. The Cambrian bioradiation is thought to have occurred when the dissolved oxygen concentrations in the seawater exceed the relatively high oxygen requirements of animals5. Although multiple lines of evidence demonstrate a short-lived return to widespread seawater anoxia at the Ediacaran-Cambrian (E-C) boundary6,7,8,9,10,11,12, evidence from sedimentary Mo concentrations5,13,14 and isotopes15, U isotopes16,17 and rare earth element concentrations18 demonstrates that the global seawater became progressively oxygenated in the early Cambrian13,15,17,19,20,21. The mechanism and magnitude of raising atmospheric oxygen concentrations, however, remain to be fundamental yet unresolved questions.
A rise in atmospheric oxygen concentrations can lead to chemical changes in the Cambrian seawater and trigger the metazoan evolution15,22,23,24,25. However, the high level of atmospheric oxygen alone cannot explain the rapid rise in the biological complexity of Cambrian age26. The supply of essential nutrients, such as P, Ca, Fe, Na+, K+ and Mo, were also considered to have exerted a significant control on the evolution of myriad life forms21,27,28. It has been suggested that a suddenly enhanced weathering of the continental crust during the E-C transition acted as a trigger for the rise in atmospheric oxygen concentrations29,30,31. The enhanced weathering and the rise of O2 may pave the way for the development of large and complex Cambrian animals. Here, we present the geochemical proxies of carbonates, such as the carbon isotope compositions, rare earth elements (REE + Y) distribution patterns and water-insoluble element concentrations, to track the mechanism and magnitude of atmospheric oxygen fluctuation during the early Cambrian.
The target carbonates were sampled from an Ediacaran-Cambrian sequence at the Tangshan section in the Lower Yangtze basin, South China (Fig. 1). This region belongs to part of the Jiangnan orogen that is located between the Yangtze platform in the northwest and the Cathaysia foldbelt in the southeast32. During the E-C transition, the Yangtze platform was located at low to middle latitudes in the Northern Hemisphere (Fig. 1A)33. It was evolved from a passive continental margin basin in the early Neoproterozoic to a series of rift basins in the middle Neoproterozoic, with a continental shelf to a marine basin from the northwest to the southeast32. The carbonate platform was surrounded by narrow marginal transitional zones to the north (~800 km long) and to the southwest (>1600 km long)9,34,35, along which shallow-water carbonates sharply change into basinward black chert/shale successions35,36,37. These sedimentary sequences on the continental shelf are characterized by repeated transgression–regression events36. The post-Fortunian lower Cambrian on the Yangtze platform, from the shelf (i.e. Xiaotan, Shantan, Jinsha and Weng’an) to the slope (i.e. Songtao and Longbizui), started with black shale with similar lithologic markers or/and the first appearance of the trilobite T. niutitangensis9,38.
The Tangshan section is located in the eastern Anhui of South China (Fig. 1B). The strata there were deposited in a restricted shallow water39. From the bottom to the top (Fig. 2), the strata are composed of the Dengying Formation (DY) of the late Ediacaran, and the Lengquanwang Formation (LQW), the Bantang Formation (BT) and the Shanaoding Group (SAD) of the Cambrian40. Small shelly fossils, including Anabarites, have been found at the base of the LQW Formation, indicating that the E-C boundary is roughly between the DY and LQW formations.
The Lower Cambrian strata in South China are poor dated by isotope geochronology. The highly metalliferous black shales with high Ni-Mo concentrations from the Niutitang and Xiaoyanxi formations was once considered as close to the E-C boundary, but recent radiometric ages of 539.4 ± 2.9 Ma41, 532.3 ± 0.7 Ma42 and 536.3 ± 5.5 Ma34 from ash beds below the highly metalliferous black shales indicate that they are younger than the E-C boundary at 541.0 ± 1.0 Ma. The onset of the globally negative δ13C excursion at the lowermost Cambrian strata starts at the top of the Dengying Formation, referred to as the Basal Cambrian C-isotope Excursion (BACE)43,44,45,46, was constrained at 541.00 ± 0.13 Ma in Oman47,48.
Results
Petrographic observations indicate that most samples on the Tangshan section consist of microcrystalline dolomite with Mg/Ca ratios higher than 0.719. The dolostones show minor to dense packing equigranular and anhedral crystals under a microscope and in SEM images (Fig. 3a,b), indicating primary dolomite textures without significant alteration and metamorphism. Few samples with Mg/Ca ratios lower than 0.7 (Supplementary Table S1) are composed of both calcite and dolomite, as indicated by XRD measurement (Fig. 4). In addition, the SEM image of Sample 12CH128 with a Mg/Ca ratio of 0.38 shows a calcite assemblage among the dolomite grains (Fig. 3c,d), indicating dedolomitization during burial diagenesis.
Variations in δ13Ccarb values and REE + Y patterns for the Tangshan section have been discussed by Li et al.40 in detail. The δ13Corg values vary from −31.77 to −22.24‰ (n = 84), with four positive (DP1, LP1, LP2 and BP1) and four negative (ECN, LN1, LN2, and BN1) excursions on the profile (Fig. 2). The δ13Corg values for the DY Formation in the late Ediacaran vary from −28.15 to −22.24‰ (DP1), and they are positively correlated with the δ13Ccarb values. The first negative δ13Corg excursion (ECN) is coupled with BACE43,44,45,46 at the E-C transition (Fig. 3). An overlying positive δ13Corg excursion (LP1) from −28.17 to −26.29‰ occurs in the lower part of the LQW Formation in the early Cambrian, followed by a negative δ13Corg excursion (LN1) from −29.32 to −31.66‰. A second positive δ13Corg excursion (LP2) from −29.22 to −25.75‰ overlies LN1, followed by a third negative δ13Corg excursion (LN2) from −29.46 to −28.53‰ at the boundary between the LQW and BT formations. Upwards, the δ13Corg values vary with several δ13Ccarb fluctuations (Fig. 2), around −27.6 to −29.6‰ (including BP1 and BN1 for δ13Ccarb) from the upper BT formation to the SAD Formation. The difference between the carbonate and organic carbon isotope compositions (Δ13Ccarb-org = δ13Ccarb − δ13Corg) varies between 22.31‰ and 30.93‰, averaging 22.27‰ (Fig. 2). In particular, the Δ13Ccarb-org values in the early Cambrian LQW Formation (average 28.12‰) are abruptly higher than those in the late Ediacaran DY Formation (average 26.28‰) (Fig. 2).
Discussion
Preservation of primary geochemical features
Most samples from the Tangshan section are composed of dolomite with minor calcite. Dolomite grains show uniform equigranular and anhedral crystals with clear boundary (Fig. 3a,b, and Supplementary Table S1), implying insignificant alteration during post-depositional processes. Several samples were dedolomitized to calcite (Fig. 3c,d). In general, geochemical features such as carbon and oxygen isotope compositions and trace element concentrations are interpreted as primary geochemical signals related to the changes in seawater composition40. However, post-depositional processes such as microbial remineralization and diagenetic alteration have a potential to alter the primary δ13Corg values49,50. In particular, the shallow marine carbonates on the Tangshan section could be periodically exposed during sea-level oscillations40 and thus influenced by meteoric alteration50. Therefore, it is essential to assess the preservation of primary δ13Corg values. Although the preferential degradation of labile organic compounds can result in an elevation in δ13Corg values of the residual organic carbon, it is impossible to generate a 4 to 5‰ δ13Corg change in δ13Corg values of the residual organic carbon during degradation51. In this regard, the δ13Corg shift from −31.77 to −22.24‰ with a difference as large as 9‰ cannot be produced by the diagenetic alteration alone. In addition, the H/C ratios are higher than 0.2 (except several samples in the LQW and BT formations) and there is no correlations between H/C and δ13Corg (Fig. 5, and Supplementary Table S1), indicating insignificant effect by the diagenetic alteration.
Carbon isotope chemostratigraphy
The E-C boundary is defined by the first appearance of small shelly fossils, Treptichnus Pedum52,53,54. However, the relationship between the Ediacaran and Cambrian biotas is poorly known due to stratigraphic gaps, taphonomic bias, the restriction of metazoans to oxygenated habits, and the difficulty in integrating commonly disparate bio- and chemostratigraphic data5,31,46,55. The BACE, on the other hand, was followed by the rapid appearance and diversification of bilaterian animals in the early Cambrian according to a combined study of lithology, fossils and geochemical data5,47. Thus, it can be constrained and better correlated within the interval of characteristic Cambrian-type skeletal fossil distribution5,34,38,46,56,57.
The E-C successions on the Tangshan section can be divided into eight δ13Ccarb intervals, from bottom to top named DP1, ECN, LP1, LN1, LP1, LN2, BP1 and BN140. The δ13Corg values increase from −28.15 to −24.50‰ (DP1) in the Dengying Formation, which can be correlated in trend and magnitude with those of the Ganziping section in western Yunnan34, the Jiuqunao, Hezi’ao and Jijiapo sections in Hubei58. The DP1 are higher than the contemporaneous values for the Shantan section, the Songtao section49, the Longbuzi section59, the Yuantuwan-Lijiatuo section49 and the Ara Group60.
The δ13Corg excursions are coupled with the δ13Ccarb excursions in the late Ediacaran DY and the early Cambrian LQW formations (Fig. 5A). ECN, a negative δ13Corg excursion of −4.4‰ from −27.32 to −31.77‰ (Supplementary Table S1), can be identified at the E-C transition (Fig. 2), which can be correlated with the N1 at Xiaotan57 and Shatan49 of inner shelf, Anjiahe of outer shelf38 and Ganziping of shelf margin34. Corresponding shifts can also be found in the Jiuqunao, Hezi’ao and Jijiapo sections on the Yangtze platform56 (Fig. 6) and a shallow-water section in the western Anti-Atlas margin of Morocco46. The worldwide ECN or N1 negative δ13Ccarb excursion, named the BACE5,45, is consistent with the first appearance of small shelly fossils of Anabarites trisulcatus zone33, which was calibrated as the beginning of Cambrian61.
Above the ECN (or BACE), a first large positive excursion in both δ13Ccarb and δ13Corg (LP1) occurred (Fig. 5), followed by a second negative excursion in both δ13Ccarb and δ13Corg (LN1). The ECN and LN1 can be correlated with the shallow-water carbonates from the other sections on the Yangtze platform (Fig. 6), but they may be lost on some sections due to the unconformity and stratigraphic condensation in shallow-water settings37,62. The δ13Ccarb excursion in LN1 can be correlated with Cycle III on the Dovortsy section in Siberia63, N2 on the Jiuqunao section in Yichang58 and the Anjiahe section in Yangtze Gorge38, N2i on the Xiaotan section in Yunnan64 and N2 on the Jiuqunao, Hezi’ao and Jijiapo sections58. Upwards, the second positive δ13Ccarb and δ13Corg excursions (LP2) occurred and were overlain by the third negative δ13Ccarb and δ13Corg excursions (LN2), which can be correlated with N3 on the Anjiahe section37, C457 or N3 on the Xiaotan section37, N3 on the Shatan49, Daotuo37, Longbizui34 and Yuanwutan sections49. In the present study, the three negative δ13Ccarb and δ13Corg excursions (ECN, LN1 and LN2) are identified in the early Cambrian (Fig. 6). It should be noted that there is also a small negative δ13Ccarb excursion in the SAD Formation, but the δ13Corg values are relatively stable (Fig. 2).
In addition, the δ13Corg values largely vary between the different sections because Corg can be dominated by local influences associated with the mechanism and rate of carbon fixation, the degree of remineralization, terrestrial organic influx, and trophic structure38. Existing δ13Corg values from the earliest Cambrian in South China show a large variation from −31 to −37‰ (Fig. 6)34,38,49,57,58.
The Δ13Ccarb-org between carbonate and organic matter
The δ13Ccarb and δ13Corg values are coupled in the late Ediacaran DY Formation and the early Cambrian LQW Formation, but decoupled in the BT and SAD formations (Fig. 5A). In particular, the Δ13Ccarb-org values in the early Cambrian LQW Formation (average 28.12‰) are considerably higher than those in the late Ediacaran DY Formation (average 26.28‰) (Fig. 7), which also occurs on the other contemporary successions although the absolute Δ13Ccarb-org values vary with depositional settings (Fig. 7).
The evidence that the evolution of oxygenic photosynthesis had occurred by ~2.78 Ga is provide by the presence of 2-α methylhapanes from O2-producing cyanobacteria65 and sterols from O2-requiring eukaryotes66. Shortly before the Neoproterozoic glaciations, a significant diversification of eukaryotes occurred67. In the narrow time interval between the Sturtian and Marinoan glaciations, the steroid diversity and marine planktonic algae had a rapid rise68. The rise of algae would create nutrient, energy and photosynthesis, driving ecosystems towards larger and increasing complex organism68 and huge environmental changes15,17,20. Cultures show that the Δ13Ccarb-org values of marine algae are controlled by the environmental factors, such as temperature, growth rate, pH, CO2 and O2 concentration69. It was concluded that the Δ13Ccarb-org values between CO2 and cells are constant (~19‰) from 10 °C to 30 °C70. Seawater temperatures were proposed to be lower than 34 °C during the Cambrian71,72. In this regard, the temperature effect on Δ13Ccarb-org values between CO2 and cells can be negligible during the E-C transition. It is shown that a 2.7‰ change in carbon isotope fractionation can be associated with 0.1 shift in pH values69, which is dependent upon PCO2 and the influx of alkalinity. However, the [CO3]2− concentrations are relatively stable in the early Cambrian29. Therefore, the effect of pH variation on Δ13Ccarb-org values may not be important during the E-C transition.
Different photosynthesis species may be a factor related to 13C/12C fractionation, which may result from differences in growth rates and/or particulate organic carbon production3. Several experiments have indicated that the carbon isotope fractionation in algae or plants is primarily driven by the enzyme ribulose-1,5-bisphsphate Carboxylase/Oxygenase (RubisCO)73, which is responsible for the fixation of CO2 into organic compounds. The carbon fractionation between marine species and the partial pressure of atmospheric carbon dioxide (pCO2) is clearly CO2-dependent, with higher pCO2 leading to higher carbon isotope fractionation, which have been tested by dinoflagellate species3 and coccolithophorid with different growth rates4. Intrinsic carbon isotope fractionations associated with RubisCO have been estimated to be ~22‰ to 30‰74. It has been shown that the δ13C values of pCO2 are associated with the carbon isotope composition of carbonate75. The atmospheric concentration of CO2 gas is kept near the Henry’s Law equilibrium with surface seawater [HCO3]− through gas exchange during turbulent mixing75. Considering the limited variations in the pH, temperature and atmospheric pressure during the E-C transition29,71,72, the equilibrium constant of carbon isotope exchange between gaseous and dissolved CO2 varies in a small range76,77. Thus, the limited carbon isotope fractionation between the atmospheric CO2 and the marine dissolved inorganic carbon can be expected. The 13C-discrimination between carboxylating reactions (RubisCO) and atmospheric CO2 would likely be the reason for the abrupt changes in Δ13Ccarb-org during the E-C transition. However, the enhanced chemical weathering rate from the late Ediacaran DY Formation to the early Cambrian LQW Formation would cause a sudden CO2 drawdown and cooling, which is recorded by the CIA index30,37,78,79. It has been proven that the decrease of pCO2 would have induced a decrease in the Δ13Ccarb-org values3,4. In this way, the widespread and abrupt increase in Δ13Ccarb-org values during the E-C transition would not result from the variation in pCO2 (Fig. 7).
The lowered δ13Corg value can also be induced by the methanogenesis80 and the elevated Δ13Ccarb-org excursion could be caused by the methanotrophic biomass81. However, the anaerobic carbon monoxide dehydrogenases (CODH) occur under strongly anoxic conditions82. In addition, the least oxygenation of surface seawater5,13,14,15,16,17 and the model calculations83 indicate the sulfate concentrations during the early Cambrian were relatively high, which would reduce the role of methanogenesis. Thus, this possibility can safely be ruled out.
After the other environmental factors have been fully excluded, the best explanation for the abrupt increase in Δ13Ccarb-org values during the E-C transition is a rise in atmospheric O2. The dependence of biological carbon isotope fractionation on the partial pressure of atmospheric oxygen (pO2) has been demonstrated in laboratory experiments with vascular land plants, bryophytes, and marine phytoplankton84,85,86. Increased carbon isotopic discrimination for light isotopes (12C) is observed at increasing O2 levels because atmospheric CO2 levels rise within the plant of plankton cell as photorespiration begin to outpace photosynthesis, which in turn increases the carbon isotope fractionation during growth85. In this way, the pO2 increase can be expected by the increase of Δ13Ccarb-org values though the absolute values of pO2 are notoriously difficult to track87. Beerling et al.85 have modeled the atmospheric O2 dependence of the carbon isotope fractionation between fossil organic matter and CO2 (equation 4 in their paper), which was modified by Saltzman, et al.88 as:
where Δ13Cmeas is the measured value for the difference between δ13Ccarb and δ13Corg, and Δ13Cinitial is an estimate of the baseline value. In the present study, the average Δ13Ccarb-org value of 26.08‰ for the late Ediacaran DY formation is taken as the Δ13Cinitial value and the average Δ13Ccarb-org values of 28.07‰ for the early Cambrian LQW and BT formations are taken as the Δ13Cmeas value. A value of J = 5 is selected in the calculation based on the best fit to both the experimental data for isotopic discrimination in modern marine phytoplankton and the results of isotope mass balance model88. Then, the abrupt increase in the Δ13Ccarb-org values of approximately +2‰ (Fig. 7) indicates that the mass of oxygen in the atmosphere was doubled in the early Cambrian85,88. Although oxygen levels during the Late Neoproterozoic are difficult to constrain, it has been proposed that the oxygen levels are higher than 8% to 15% PAL near the end of the Neoproterozoic89. In this regard, the pO2 of early Cambrian would be at least 16 to 30% PAL, which is consistent with 10% to 40% PAL estimated for the Cambrian87. The decrease of Δ13Ccarb-org values in the SAD Formation, equivalent to the middle Tommotian (Fig. 6), could be generated by another shift in pO2 or pCO284,85,90 or the occurrence of new photosynthetic species91.
Continental trace element influx
During the E-C transition, a number of continental blocks were converged toward a large and complex continent, Gondwana27,92,93. In general, this supercontinent is considered as an assembly of four major continental blocks along three collisional orogens27. The convergent plate boundaries along the Gondwana-forming sutures could have served as effective sources for nutrient delivery94. An unconformity was formed globally62,95, followed by a worldwide transgression in the early Cambrian36, which would expand the area of shallow epicontinental seas, leading to an increased accommodation space and a sustained flux of continental weathered products to the seawater (Fig. 8). This unconformity is recorded by shallow-water marine carbonates directly overlying on the continental crystalline basement, or basin/slope organic-rich shales/chert unconformably overlying the Ediacaran carbonate shelf at the Shatan and Anjiahe sections, but it is not recorded by those in deep-water sequences on the Longbizui and Yanwutan sections34,37,49,58,96. The geochemical features, such as δ18Ocarb values, REE + Y patterns, water-insoluble element concentrations and Y/Ho ratios, suggest that the dolostones of the late Ediacaran DY Formation were deposited from normal seawater97,98, but those of the early Cambrian LQW Formation were influenced by a large amount of continental weathered materials, which is consistent with a rapid rise in Sr isotope ratios99,100,101,102. A number of elements, such as Be, Al, Sc, Co, Ga, Cs, REE, Hf and Th, are water-insoluble during chemical weathering given that they are concentrated in refractory minerals or strongly absorbed by clays103. After intensive weathering, these elements were enriched in weathered products, which would be concentrated in the impure freshwater, in river, stream and lake104, and ultimately into the seawater. In this way, the element concentrations and patterns of near-shore seawater would be altered by even a minor amount of terrigenous materials because the element compositions of terrigenous materials are quite different from that of seawater104,105. The weathered products, including a widespread nutrient supply, such as P, K, Fe, Ca, Mg and others106, could be transported to the shallow seawater via aeolian and fluvial systems across the global. Especially, the weathering flux of P into the seawater would increase, which could have promoted the activity of photosynthesis, leading a rise in atmospheric O228,107.
Implications for coincidence in continental weathering and oxygen rise
The early Cambrian witnessed the explosive radiation of animals5,22,56. Complex life requires two critical ingredients: nutrient and oxygen. Geological and geochemical studies indicate a constant increase of the oxygen level and a contraction of anoxic seawater during the E-C transition5,13,14,15,16,17,18. It is hypothesized that the oxygenation was triggered through increased weathering fluxes21,28, leading to an increase in oceanic primary production and organic carbon burial28,30,108. The species engaging in ecosystem engineering might also be responsible for the oxygenation109,110.
The emergence of huge landmass and the enhanced weathering of continental crust would transfer global scattering nutrients into the seawater through wind and river drainage in the early Cambrian (Fig. 8), which is indicated by the decrease of Y/Ho ratios and the increase of Th, V, Sc and REE concentrations (Fig. 2 and Supplementary Fig. S1). Although it is difficult to accurately constrain the proportions of fresh water in the surface seawater, the relative proportions of freshwater can be obtained on the basis of the δ18Ocarb values, the water-insoluble element concentrations, Y/Ho ratios (Supplementary Fig. S3). The mixing lines between δ18Ocarb values and Y/Ho ratios represent a result of mixing two end-members between seawater and freshwater, where the freshwater contains the terrigenous material from the chemical weathering into the seawater (Supplementary Fig. S3).
The chemical weathering not only serves as a negative feedback for CO2, but also provides the nutrients for feeding of the primary producers. A stable supply of essential nutrients would also exert a significant control on the evolution of myriad life forms27. Nutrients, such as P, Ca, Fe, Na+, K+ and Mo, are predominantly enriched in crustal rocks, such as granitoids and andesite27. These elements are critically important to build the ‘hard parts’ (bones, shells, teeth), whereas some of the other nutrients are essential for cell metabolism106. Enormous nutrient into the seawater would especially feed the planktons in the sea surface106, leading to an accelerated burst of photosynthesis67,111, which may be the Earth’s only major source of molecular oxygen, an oxygen source strong enough to sustain a major atmospheric oxygen increase112. The atmospheric O2 could can reach up to 16 to 30% PAL based on the abrupt increase of Δ13Ccarb-org values during the E-C transition (Fig. 7). Although anoxic seawaters may still have existed widely in the early Cambrian87,113, evidence from sedimentary Mo concentrations5,13,14 and isotopes15, U isotopes16,17 and its rare earth element concentrations18 indicate strong oxygenation of the deep seawater15, consistent with the rise of atmospheric oxygen concentrations113,114,115 and the appearance of filter-feeding sponges30,31,110.
Although the timing and magnitude of atmospheric O2 accumulation are difficult to be accurately constrained, it is supposed to be regulated by erosion and deposition related to tectonic processes30,116. An imbalance between the oxygen production and consumption would result in a gradual rise in the oxygen accumulation30,109,116. The high pO2 is necessary to keep the reaction at the seawater-atmosphere interface in order to produce collagen that paste plural cells together. However, even if the pO2 is high enough, metazoans cannot be synthesized if the nutrient elements that make up the metazoans were not supplied106. In this regard, the simultaneously enhanced weathering input would provide necessary nutrients for biological evolution. Therefore, the increasing amount of landmass, the enhancement of chemical weathering on the continental crust, the supply of a large amount of nutrients into the shelf and high oxygen level in the atmosphere all coincide in time for promoting the birth of large multi-cellular animals.
Conclusions
The Ediacaran-Cambrian (E-C) transition is a period with substantial biospheric, environmental, and tectonic changes. The δ13Corg profile on the Tangshan section in South China records three negative excursions during the early Cambrian, which can be compared with those of the other sections in the world. The abrupt increase in Δ13Ccarb-org values indicates a rise of atmospheric oxygen concentrations at the E-C boundary. Although there are several δ13Ccarb and δ13Corg excursions in the early Cambrian, the atmospheric O2 retained high during the deposition of the LQW and BT formations. Geochemical proxies, such as REE + Y patterns, water-insoluble element concentrations and Y/Ho ratios, indicate that an enhanced weathering of the continental crust occurred from the late Ediacaran DY Formation to the early Cambrian LQW Formation during the E-C transition. The enhancement of continental weathering delivers nutrients to the seawater for primary producers, which would accelerate photosynthesis and promote the rise of atmospheric pO2. The enhanced weathering along with the rise of atmospheric oxygen, may have boosted the emergence of complex life in the early Cambrian.
Methods
Some samples and the method of trace element concentrations have been described in Li et al.40. The methods of this study are as following.
XRD
The Mg/Ca ratios of all samples were analyzed by Li et al.40. Only several sample powders with special Mg/Ca ratios were measured in the present study by X-ray diffraction (XRD). The XRD analysis was performed at University of Science and Technology of China (USTC) in Hefei using a TTR-III diffractometer (operating at 40 kV, 200 mA) equipped with a fixed graphite monochromator and a Cu target tube in the 2θ with a steep size of 0.02 degree. Meanwhile, the microstructure of the carbonate was observed by the scanning electron microscope (SEM) at USTC.
Oxygen and carbon isotopes
The methods for the analyses of carbonate C and O isotopes (δ13Ccarb and δ18Ccarb) were presented in Li et al.40. The organic carbon isotope compositions (δ13Corg) were measured on 84 samples. The 50% (v/v) guarantee reagent hydrochloric acid had been used at 60 °C for 24 h to remove the all carbonate components in whole-rock powders, and the undissolved part was washed by ultra-pure water and extracted for three times to achieve neutral solution conditions19. The undissolved part of the carbonate were dried at 70 °C and weighed, and then they were analyzed to get the TC, TH and TN of the residue by the Elementar vario EL cube at USTC. The analytical precisions of TC, TH and TN are better than ±0.3%. The organic carbon isotopes of the residues were analyzed via online EA-1112 link to Finigan Delta-Plus XL at State Key Laboratory of Organic Geochemistry in Guangzhou Institute of Geochemistry, Chinese Academy of Sciences (CAS), Guangzhou. The results are calibrated from the standard carbon black δ13CPDB = −36.91‰ and given in the standard delta notation as per mil difference to the Vienna PDB standard (VPDB). Reproducibility was generally better than ±0.5‰.
References
Anbar, A. D. & Knoll, A. H. Proterozoic ocean chemistry and evolution: a bioionorganic bridge? Science 297, 1137–1142 (2002).
Berner, R. A., VandenBrooks, J. M. & Ward, P. D. Oxygen and Evolution. Science 316, 557–558 (2007).
Hoins, M. et al. Stable carbon isotope fractionation of organic cyst-forming dinoflagellates: Evaluating the potential for a CO2 proxy. Geochim. Cosmochim. Acta 160, 267–276 (2015).
Rost, B., Zondervan, I. & Riebesell, U. Light-dependent carbon isotope fractionation in the coccolithophorid Emiliania huxleyi. Limnol. Oceanogr. 47, 120–128 (2002).
Zhu, M., Zhuravlev, A. Y., Wood, R. A., Zhao, F. & Sukhov, S. S. A deep root for the Cambrian explosion: Implications of new bio- and chemostratigraphy from the Siberian Platform. Geology 45, 459–462 (2017).
Canfield, D. E. et al. Ferruginous Conditions Dominated Later Neoproterozoic Deep-Water Chemistry. Science 321, 949–952 (2008).
Kimura, H. & Watanabe, Y. Oceanic anoxia at the Precambrian-Cambrian boundary. Geology 29, 995–998 (2001).
Liu, K., Feng, Q., Shen, J., Khan, M. & Planavsky, N. J. Increased productivity as a primary driver of marine anoxia in the Lower Cambrian. Palaeogeogr. Palaeoclimatol. Palaeoecol. 491, 1–9 (2018).
Jin, C. et al. A highly redox-heterogeneous ocean in South China during the early Cambrian (∼529–514 Ma): Implications for biota-environment co-evolution. Earth Planet. Sci. Lett. 441, 38–51 (2016).
Guo, Q. et al. Trace element chemostratigraphy of two Ediacaran-Cambrian successions in South China: Implications for organosedimentary metal enrichment and silicification in the Early Cambrian. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 194–216 (2007).
Wei, G.-Y. et al. Marine redox fluctuation as a potential trigger for the Cambrian explosion. Geology 46, 587–590 (2018).
Wen, H., Fan, H., Zhang, Y., Cloquet, C. & Carignan, J. Reconstruction of early Cambrian ocean chemistry from Mo isotopes. Geochim. Cosmochim. Acta 164, 1–16 (2015).
Scott, C. et al. Tracing the stepwise oxygenation of the Proterozoic ocean. Nature 452, 456–459 (2008).
Chang, C. et al. Characterization of trace elements and carbon isotopes across the Ediacaran-Cambrian boundary in Anhui Province, South China: Implications for stratigraphy and paleoenvironment reconstruction. J. Asian Earth Sci (2016).
Chen, X. et al. Rise to modern levels of ocean oxygenation coincided with the Cambrian radiation of animals. Nat Commun 6 (2015).
Tostevin, R. et al. Uranium isotope evidence for an expansion of anoxia in terminal Ediacaran oceans. Earth Planet. Sci. Lett. 506, 104–112 (2019).
Kendall, B. et al. Uranium and molybdenum isotope evidence for an episode of widespread ocean oxygenation during the late Ediacaran Period. Geochim. Cosmochim. Acta 156, 173–193 (2015).
Ling, H.-F. et al. Cerium anomaly variations in Ediacaran–earliest Cambrian carbonates from the Yangtze Gorges area, South China: Implications for oxygenation of coeval shallow seawater. Precambr. Res. 225, 110–127 (2013).
Krause, A. J. et al. Stepwise oxygenation of the Paleozoic atmosphere. Nature Communications 9, 4081 (2018).
Wallace, M. W. et al. Oxygenation history of the Neoproterozoic to early Phanerozoic and the rise of land plants. Earth Planet. Sci. Lett. 466, 12–19 (2017).
Mills, B., Lenton, T. M. & Watson, A. J. Proterozoic oxygen rise linked to shifting balance between seafloor and terrestrial weathering. Proceedings of the National Academy of Sciences of the United States of America 111, 9073–9078 (2014).
Erwin, D. H. et al. The Cambrian Conundrum: Early Divergence and Later Ecological Success in the Early History of Animals. Science 334, 1091–1097 (2011).
Sperling, E. A. et al. Oxygen, ecology, and the Cambrian radiation of animals. Proceedings of the National Academy of Sciences 110, 13446–13451 (2013).
Smith, M. P. & Harper, D. A. T. Causes of the Cambrian Explosion. Science 341, 1355–1356 (2013).
Mills, D. B. et al. Oxygen requirements of the earliest animals. Proceedings of the National Academy of Sciences 111, 4168–4172 (2014).
Mills, D. B. & Canfield, D. E. Oxygen and animal evolution: Did a rise of atmospheric oxygen “trigger” the origin of animals? BioEssays 36, 1145–1155 (2014).
Santosh, M., Maruyama, S., Sawaki, Y. & Meert, J. G. The Cambrian Explosion: Plume-driven birth of the second ecosystem on Earth. Gondwana Res. 25, 945–965 (2014).
Planavsky, N. J. et al. The evolution of the marine phosphate reservoir. Nature 467, 1088 (2010).
Riding, R. & Liang, L. Seawater chemistry control of marine limestone accumulation over the past 550 million years. Vol. 37 (2005).
Lenton, T. M. & Watson, A. J. Biotic enhancement of weathering, atmospheric oxygen and carbon dioxide in the Neoproterozoic. Geophys. Res. Lett. 31 (2004).
Erwin, D. H. & Tweedt, S. Ecological drivers of the Ediacaran-Cambrian diversification of Metazoa. Evol. Ecol. 26, 417–433 (2012).
Wang, J. & Li, Z. X. History of Neoproterozoic rift basins in South China: implications for Rodinia break-up. Precambr. Res. 122, 141–158 (2003).
Maloof, A. et al. The earliest Cambrian record of animals and ocean geochemical change. Bulletin of the Geological Society of America 122, 1731–1774 (2010).
Chen, D., Wang, J., Qing, H., Yan, D. & Li, R. Hydrothermal venting activities in the Early Cambrian, South China: Petrological, geochronological and stable isotopic constraints. Chem. Geol. 258, 168–181 (2009).
Guo, Q., Deng, Y., Hippler, D., Franz, G. & Zhang, J. REE and trace element patterns from organic-rich rocks of the Ediacaran–Cambrian transitional interval. Gondwana Res. 36, 94–106 (2016).
Steiner, M., Wallis, E., Erdtmann, B.-D., Zhao, Y. & Yang, R. Submarine-hydrothermal exhalative ore layers in black shales from South China and associated fossils — insights into a Lower Cambrian facies and bio-evolution. Palaeogeogr. Palaeoclimatol. Palaeoecol. 169, 165–191 (2001).
Zhai, L., Wu, C., Ye, Y., Zhang, S. & Wang, Y. Fluctuations in chemical weathering on the Yangtze Block during the Ediacaran–Cambrian transition: Implications for paleoclimatic conditions and the marine carbon cycle. Palaeogeogr. Palaeoclimatol. Palaeoecol. 490, 280–292 (2018).
Ishikawa, T. et al. Irreversible change of the oceanic carbon cycle in the earliest Cambrian: High-resolution organic and inorganic carbon chemostratigraphy in the Three Gorges area, South China. Precambr. Res. 225, 190–208 (2013).
Sun, C. Diachronism of the bantang delitic dolomite in the chaohu lake area Anhui. Regional Geology of China 4, 380–383 (Chinese with English Abstract) (1988).
Li, W.-P., Zheng, Y.-F. & Zhao, Y.-Y. Geochemical evidence from marine carbonate for enhanced terrigenous input into seawater during the Ediacaran-Cambrian transition in South China. Precambr. Res. 291, 83–97 (2017).
Compston, W., Zhang, Z., Cooper, J. A., Ma, G. & Jenkins, R. J. F. Further SHRIMP geochronology on the early Cambrian of South China. Am. J. Sci. 308, 399–420 (2008).
Jiang, S.-Y. et al. Early Cambrian ocean anoxia in South China. Nature 459, E5 (2009).
Zhu, M.-Y., Babcock, L. E. & Peng, S.-C. Advances in Cambrian stratigraphy and paleontology: Integrating correlation techniques. paleobiology, taphonomy and paleoenvironmental reconstruction. Palaeoworld 15, 217–222 (2006).
Landing, E., Peng, S., Babcock, L., Geyer, G. & Moczydlowska-Vidal, M. Global standard names for the lowermost Cambrian series and stage. Episodes 30, 287–289 (2007).
Zhu, M. Y., Strauss, H. & Shields, G. A. From snowball earth to the Cambrian bioradiation: Calibration of Ediacaran-Cambrian earth history in South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 1–6 (2007).
Maloof, A. C. et al. Constraints on early Cambrian carbon cycling from the duration of the Nemakit-Daldynian-Tommotian boundary 13C shift, Morocco. Geology 38, 623–626 (2010).
Amthor, J. E. et al. Extinction of Cloudina and Namacalathus at the Precambrian-Cambrian boundary in Oman. Geology 31, 431–434 (2003).
Bowring, S. et al. Geochronologic constraints on the chronostratigraphic framework of the Neoproterozoic Huqf Supergroup of Oman. Am. J. Sci. 307, 1097–1145 (2007).
Guo, Q. et al. Carbon isotopic evolution of the terminal Neoproterozoic and early Cambrian: Evidence from the Yangtze Platform, South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 140–157 (2007).
Oehlert, A. M. & Swart, P. K. Interpreting carbonate and organic carbon isotope covariance in the sedimentary record. Nature Communications 5, 4672 (2014).
Lehmann, M. F., Bernasconi, S. M., Barbieri, A. & McKenzie, J. A. Preservation of organic matter and alteration of its carbon and nitrogen isotope composition during simulated and in situ early sedimentary diagenesis. Geochim. Cosmochim. Acta 66, 3573–3584 (2002).
Landing, E. Precambrian-Cambrian boundary global stratotype ratified and a new perspective of Cambrian time. Geology 22, 179–182 (1994).
Brasier, M., Cowie, J. & Taylor, M. Decision on the Precambrian-Cambrian boundary stratotype. Vol. 17 (1994).
Buatois, L. A. Treptichnus pedum and the Ediacaran–Cambrian boundary: significance and caveats. Geol. Mag. 155, 174–180 (2018).
Schiffbauer, J. D. Research focus: The age of tubes: A window into biological transition at the Precambrian-Cambrian boundary. Geology 44, 975–976 (2016).
Smith, E. F. et al. The end of the Ediacaran: Two new exceptionally preserved body fossil assemblages from Mount Dunfee, Nevada, USA. Geology 44, 911–914 (2016).
Cremonese, L. et al. Marine biogeochemical cycling during the early Cambrian constrained by a nitrogen and organic carbon isotope study of the Xiaotan section. South China. Precambr. Res. 225, 148–165 (2013).
Jiang, G. et al. The origin of decoupled carbonate and organic carbon isotope signatures in the early Cambrian (ca. 542–520Ma) Yangtze platform. Earth Planet. Sci. Lett. 317–318, 96–110 (2012).
Wang, J., Chen, D., Yan, D., Wei, H. & Xiang, L. Evolution from an anoxic to oxic deep ocean during the Ediacaran-Cambrian transition and implications for bioradiation. Chem. Geol. 306-307, 129–138 (2012).
Fike, D. A. Carbon and Sulfur Isotopic Constraints on Ediacaran Biogeochemical Processes, Huqf Supergroup, Sultanate of Oman ph.D. thesis, Cambridge, Massachusetts, (2007).
Kaufman, A. et al. Integrated chronostratigraphy of Proterozoic-Cambrian boundary beds in the western Anabar region, Northern Siberia. Geol. Mag. 133, 509–533 (1996).
Peters, S. E. & Gaines, R. R. Formation of the ‘Great Unconformity’ as a trigger for the Cambrian explosion. Nature 484, 363–366 (2012).
Magaritz, M., Kirschvink, J. L., Latham, A. J., Zhuravlev, A. Y. & Rozanov, A. Y. Precambrian/Cambrian boundary problem: Carbon isotope correlations for Vendian and Tommotian time between Siberia and Morocco. Geology 19, 847–850 (1991).
Li, D. et al. Carbon and strontium isotope evolution of seawater across the Ediacaran–Cambrian transition: Evidence from the Xiaotan section, NE Yunnan, South China. Precambrian Research 225, 128–147 (2013).
Brocks, J. J., Logan, G. A., Buick, R. & Summons, R. E. Archean molcular fossils and the early rise of eukaryotes. Science 285, 1033–1036 (1999).
Summons, R. E., Jahnke, L. L., Hope, J. M. & Logan, G. A. 2-Methylhopanoids as biomarkers for cyanobacterial oxygenic photosynthesis. Nature 400, 554–557 (1999).
Lenton, T. M., Boyle, R. A., Poulton, S. W., Shields-Zhou, G. A. & Butterfield, N. J. Co-evolution of eukaryotes and ocean oxygenation in the Neoproterozoic era. Nature Geoscience 7, 257 (2014).
Brocks, J. J. et al. The rise of algae in Cryogenian oceans and the emergence of animals. Nature 548, 578 (2017).
Hinga, K. R., Arthur, M. A., Pilson, M. E. Q. & Whitaker, D. Carbon isotope fractionation by marine phytoplankton in culture: The effects of CO2 concentration, pH, temperature, and species. Global Biogeochem. Cycles 8, 91–102 (1994).
Degens, E. T., Guillard, R. R. L., Sackett, W. M. & Hellebust, J. A. Metabolic fractionation of carbon isotopes in marine plankton; 1, Temperature and respiration experiments. Deep-Sea Research and Oceanographic Abstracts 15, 1–9 (1968).
Jaffrés, J. B. D., Shields, G. A. & Wallmann, K. The oxygen isotope evolution of seawater: a critical review of a long-standing controversy and an improved geological water cycle model for the past 3.4 billion years. Earth Sci. Rev. 83, 83–122 (2007).
Robert, F. & Chaussidon, M. A palaeotemperature curve for the Precambrian oceans based on silicon isotopes in cherts. Nature 443, 920–921 (2006).
Roeske, C. A. & O’Leary, M. H. Carbon isotope effects on enzyme-catalyzed carboxylation of ribulose bisphosphate. Biochemistry 23, 6275–6284 (1984).
Scott, K. M. et al. Kinetic isotope effect and biochemical characterization of form IA RubisCO from the marine cyanobacterium Prochlorococcus marinus MIT9313. Limnol. Oceanogr. 52, 2199–2204 (2007).
Ripperdan, E. L. In Stable Isotope Geochemistry Vol. 43 (eds Valley, J. W. & Cole, D. R.) 637–662 (Blackburg, Virginia, 2001).
Myrttinen, A., Becker, V. & Barth, J. A. C. A review of methods used for equilibrium isotope fractionation investigations between dissolved inorganic carbon and CO2. Earth-Sci. Rev. 115, 192–199 (2012).
Szaran, J. Carbon isotope fractionation between dissolved and gaseous carbon dioxide. Chem. Geol. 150, 331–337 (1998).
Donnadieu, Y., Goddéris, Y., Ramstein, G., Nédélec, A. & Meert, J. A. ‘snowball Earth’ climate triggered by continental break-up through changes in runoff. Nature 428, 303 (2004).
Dessert, C., Dupré, B., Gaillardet, J., François, L. M. & Allègre, C. J. Basalt weathering laws and the impact of basalt weathering on the global carbon cycle. Chem. Geol. 202, 257–273 (2003).
Qu, Y., Crne, A. E., Lepland, A. & van Zuilen, M. A. Methanotrophy in a Paleoproterozoic oil field ecosystem, Zaonega Formation, Karelia, Russia. Geobiology 10, 467–478 (2012).
Zerkle, A. L., House, C. H. & Brantley, S. L. Biogeochemical signatures through time as inferred from whole microbial genomes. Am. J. Sci. 305, 467–502 (2005).
Techtmann, S., Colman, A., Lebedinsky, A., Sokolova, T. & Robb, F. Evidence for Horizontal Gene Transfer of Anaerobic Carbon Monoxide Dehydrogenases. Frontiers in Microbiology 3 (2012).
Schrag, D. P., Higgins, J. A., Macdonald, F. A. & Johnston, D. T. Authigenic Carbonate and the History of the Global Carbon Cycle. Science 339, 540–543 (2013).
Berner, R. A. et al. Isotope Fractionation and Atmospheric Oxygen: Implications for Phanerozoic O2 Evolution. Science 287, 1630–1633 (2000).
Beerling, D. J. et al. Carbon isotope evidence implying high O2/CO2 ratios in the Permo-Carboniferous atmosphere. Geochim. Cosmochim. Acta 66, 3757–3767 (2002).
Fletcher, B. J., Brentnall, S. J., Quick, W. P. & Beerling, D. J. BRYOCARB: A process-based model of thallose liverwort carbon isotope fractionation in response to CO2, O2, light and temperature. Geochim. Cosmochim. Acta 70, 5676–5691 (2006).
Sperling, E. A. et al. Statistical analysis of iron geochemical data suggests limited late Proterozoic oxygenation. Nature 523, 451 (2015).
Saltzman, M. R. et al. Pulse of atmospheric oxygen during the late Cambrian. Proceedings of the National Academy of Sciences (2011).
Canfield, D. E. The early hisitory of atmospheric oxygen. Annual Review of Earth and Planetary Sciences: homage to Robert M. Garrels 33, 1–36 (2005).
Kump, L. R. & Arthur, M. A. Interpreting carbon-isotope excursions: carbonates and organic matter. Chem. Geol. 161, 181–198 (1999).
Zhang, X. et al. Triggers for the Cambrian explosion: Hypotheses and problems. Gondwana Res. 25, 896–909 (2014).
Meert, J. G. & Lieberman, B. S. The Neoproterozoic assembly of Gondwana and its relationship to the Ediacaran-Cambrian radiation. Gondwana Res. 14, 5–21 (2008).
Collins, A. S., Clark, C. & Plavsa, D. Peninsular India in Gondwana: The tectonothermal evolution of the Southern Granulite Terrain and its Gondwanan counterparts. Gondwana Res. 25, 190–203 (2014).
Squire, R. J., Campbell, I. H., Allen, C. M. & Wilson, C. J. L. Did the Transgondwanan Supermountain trigger the explosive radiation of animals on Earth? Earth Planet. Sci. Lett. 250, 116–133 (2006).
Keller, C. B. et al. Neoproterozoic glacial origin of the Great Unconformity. Proceedings of the National Academy of Sciences 116, 1136–1145 (2019).
Ishikawa, T. et al. Carbon isotope chemostratigraphy of a Precambrian/Cambrian boundary section in the Three Gorge area, South China: Prominent global-scale isotope excursions just before the Cambrian Explosion. Gondwana Research 14, 193–208 (2008).
Zhang, J. & Nozaki, Y. Rare earth elements and yttrium in seawater: ICP-MS determinations in the East Caroline, Coral Sea, and South Fiji basins of the western South Pacific Ocean. Geochim. Cosmochim. Acta 60, 4631–4644 (1996).
Kamber, B. S. & Webb, G. E. The geochemistry of late Archaean microbial carbonate: implications for ocean chemistry and continental erosion history. Geochim. Cosmochim. Acta 65, 2509–2525 (2001).
Cox, G. M. et al. Continental flood basalt weathering as a trigger for Neoproterozoic Snowball Earth. Earth Planet. Sci. Lett. 446, 89–99 (2016).
van der Meer, D. G. et al. Reconstructing first-order changes in sea level during the Phanerozoic and Neoproterozoic using strontium isotopes. Gondwana Res. 44, 22–34 (2017).
Shields, G. & Stille, P. Diagenetic constraints on the use of cerium anomalies as palaeoseawater redox proxies: An isotopic and REE study of Cambrian phosphrorites. Chem. Geol. 175, 29–48 (2001).
Melezhik, V. A., Gorokhov, I. M., Kuznetsov, A. B. & Fallick, A. E. Chemostratigraphy of Neoproterozoic carbonates: implications for ‘bind dating’. Terra Nova 13, 1–11 (2001).
Zhao, M.-Y. & Zheng, Y. F. Marine carbonate records of terrigenous input into Paleotethyan seawater: Geochemical constraints from Carboniferous limestones. Geochim. Cosmochim. Acta 141, 508–531 (2014).
Lawrence, M. G., Greig, A., Collerson, K. D. & Kamber, B. S. Rare Earth Element and Yttrium Variability in South East Queensland Waterways. Aquatic Geochemistry 12, 39–72 (2006).
Nothdurft, L. D., Webb, G. E. & Kamber, B. S. Rare earth element geochemistry of Late Devonian reefal carbonates, Canning Basin, Western Australia: confirmation of a seawater REE proxy in ancient limestones. Geochim. Cosmochim. Acta 68, 263–283 (2004).
Maruyama, S. et al. The naked planet Earth: Most essential pre-requisite for the origin and evolution of life. Geoscience Frontiers 4, 141–165 (2013).
Reinhard, C. T. et al. Evolution of the global phosphorus cycle. Nature 541, 386 (2016).
Canfield, D. E., Poulton, S. W. & Marbonne, G. M. Late-Neoproterozoic deep-ocean oxygenation and the rise of animal life. Science 315, 92–95 (2007).
Tatzel, M., von Blanckenburg, F., Oelze, M., Bouchez, J. & Hippler, D. Late Neoproterozoic seawater oxygenation by siliceous sponges. Nature. Communications 8, 621 (2017).
Mángano, M. G. & Buatois, L. A. Decoupling of body-plan diversification and ecological structuring during the Ediacaran–Cambrian transition: evolutionary and geobiological feedbacks. Proceedings: Biological Sciences 281, 1–9 (2014).
Han, T. M. & Runnegar, B. Megascopic eukaryotic algae from the 2.1-billion-year-old negaunee iron-formation, Michigan. Science 257, 232–235 (1992).
Johnston, D. T., Wolfe-Simon, F., Pearson, A. & Knoll, A. H. Anoxygenic photosynthesis modulated Proterozoic oxygen and sustained Earth’s middle age. Proceedings of the National Academy of Sciences 106, 16925–16929 (2009).
Goldberg, T., Strauss, H., Guo, Q. & Liu, C. Reconstructing marine redox conditions for the Early Cambrian Yangtze Platform: Evidence from biogenic sulphur and organic carbon isotopes. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 175–193 (2007).
Jin, C. et al. Spatiotemporal variability of ocean chemistry in the early Cambrian, South China. Science China Earth Sciences 57, 579–591 (2014).
Zhang, J., Fan, T., Algeo, T. J., Li, Y. & Zhang, J. Paleo-marine environments of the Early Cambrian Yangtze Platform. Palaeogeogr. Palaeoclimatol. Palaeoecol. 443, 66–79 (2016).
Marais, D. J. D., Strauss, H., Summons, R. E. & Hayes, J. M. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature 359, 605 (1992).
Acknowledgements
This study was supported by funds from the National Natural Science Foundation of China (41873006) and the Chinese Academy of Sciences (XDB18000000). Thanks are due to Yan’an Shen, Jun Yan and Jing Huang for their assistance with field sampling and discussion, and to Ping’an Peng for his assistance with the organic carbon isotope analysis in Guangzhou.
Author information
Authors and Affiliations
Contributions
Weiping Li has undertaken this work and completed the analysis of all data. Yanyan Zhao has drafted and finished the manuscript. Mingyu Zhao has made contributions to the interpretation of data and revised it critically for important content. Xiang-Ping Zha has assisted the carbon and oxygen isotope analyses of carbonates. Yong-Fei Zheng has designed this research made substantial contributions to the conception. All authors approve this submission.
Corresponding authors
Ethics declarations
Competing Interests
The authors declare no competing interests.
Additional information
Publisher’s note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Supplementary information
Rights and permissions
Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to the original author(s) and the source, provide a link to the Creative Commons license, and indicate if changes were made. The images or other third party material in this article are included in the article’s Creative Commons license, unless indicated otherwise in a credit line to the material. If material is not included in the article’s Creative Commons license and your intended use is not permitted by statutory regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of this license, visit http://creativecommons.org/licenses/by/4.0/.
About this article
Cite this article
Li, WP., Zhao, YY., Zhao, MY. et al. Enhanced weathering as a trigger for the rise of atmospheric O2 level from the late Ediacaran to the early Cambrian. Sci Rep 9, 10630 (2019). https://doi.org/10.1038/s41598-019-47142-3
Received:
Accepted:
Published:
DOI: https://doi.org/10.1038/s41598-019-47142-3
- Springer Nature Limited