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.

Figure 1
figure 1

(A) Tentative paleogeographic reconstruction for ca. 542 Ma (modified after Maloof et al.33). Cratons are labeled: LAUR—Laurentia, AM—Amazonia, WA—West Africa, BALT—Baltica, SIB—Siberia, MONG—Mongolia, and SC—South China. (B) Depositional environments during the Ediacaran-Cambrian transition on the Yangtze platform and geological settings of the Tangshan section at Chaohu in eastern Anhui, South China (modified after Li et al.40). The insert map is the location of Yangtze platform and Cathaysia foldbelt. Red box and triangle denote the studied area and sample section, respectively.

Full size image

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.

Figure 2
figure 2

Lithological and geochemical profiles for dolostone from the Ediacaran to early Lower Cambrian strata in the Tangshan section.

Full size image

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.

Figure 3
figure 3

The SEM photograph of 12CH128 in the early Cambrian, (a) and (b) the compacted micrystalline dolomite. (b) and (c) the different forms of autogenetic calcite filled in the dolostone.

Full size image
Figure 4
figure 4

XRD results of the selected dolostone samples on the Tangshan section. The dolomite and calcite have characteristic peak below 30 and above 30, respectively. Dolostones 12CH122, 12CH123 and 12CH128 comprise much calcite crystal in the whole power.

Full size image

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.

Figure 5
figure 5

Correlations between δ13Corg values and (A) δ13Ccarb values, (B) H/C ratios, (C) δ18Ocarb values for dolostone on the Tangshan section. The arrows in (B) represent the high H/C ratios for dolostone in the early Cambrian LQW formation.

Full size image

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.

Figure 6
figure 6

Variations in organic carbon isotope compositions across the Ediacaran-Cambrian transition on the Tangshan section and their stratigraphic correlations with those at Xiaotan57, Shatan49, Anjiahe38, Ganziping34, Yangtze Gorge58 and the western Anti-Atlas margin of Morocco46.

Full size image

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).

Figure 7
figure 7

Plot of abrupt change in Δ13Ccarb-org values for the Tangshan section (this study), the Xiaotan section57,64, the Anjiahe section38 and the compiled data for Jiuqunao, Hezi’ao and Jijiapo sections at the Ediacaran-Cambrian boundary (the solid red line). The yellow areas indicate the variation of Δ13Ccarb-org values, and the blue areas are the correlation between δ13Corg values in Fig. 6.

Full size image

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:

$${{\rm{\Delta }}}^{{\rm{13}}}{{\rm{C}}}_{{\rm{meas}}}={{\rm{\Delta }}}^{{\rm{13}}}{{\rm{C}}}_{{\rm{initial}}}+{\rm{J}}\ast [\frac{{{\rm{M}}}_{{\rm{o2}}}}{38}]-{\rm{1}}$$
(1)

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.

Figure 8
figure 8

A conceptual model for the emergence of huge landmass and enhanced weathering, which results in the global scattering of nutrients into the seawater through river drainage and wind. This model coincide with the rise of atmospheric pO2 due to the photosynthesis in the shallow seawater during Ediacaran-Cambrian transition.

Full size image

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‰.