top

Abstract

Figures and Tables

Supplementary Figures and Tables

top

Letter

In contrast to many geochemical proxies that evaluate local ancient marine redox including iron speciation and trace metal (Mo, U, V) enrichments, the uranium isotope composition (δ²³⁸U, the per mille deviation of the ²³⁸U/²³⁵U ratio relative to CRM 145 standard) of seawater can be used to evaluate ocean oxygenation at a globally integrated scale. This is possible due to the long residence time and uniform δ²³⁸U of uranium in the modern ocean and predicted for the Cambrian ocean (Weyer et al., 2008

Weyer, S., Anbar, A.D., Gerdes, A., Gordon, G.W., Algeo, T.J., Boyle, E.A. (2008) Natural fractionation of 238U/235U. Geochimica et Cosmochimica Acta 72, 345–359.

; Dahl et al., 2014

Dahl, T.W., Boyle, R.A., Canfield, D.E., Connelly, J.N., Gill, B.C., Lenton, T.M., Bizzarro, M. (2014) Uranium isotopes distinguish two geochemically distinct stages during the later Cambrian SPICE event. Earth and Planetary Science Letters 401, 313–326.

; Tissot and Dauphas, 2015

Tissot, F.L.H., Dauphas, N. (2015) Uranium isotopic compositions of the crust and ocean: Age corrections, U budget and global extent of modern anoxia. Geochimica et Cosmochimica Acta 167, 113–143.

). The δ²³⁸U proxy has been utilised to track past global ocean redox during three known oceanic anoxic events (Montoya-Pino et al., 2010

Montoya-Pino, C., Weyer, S., Anbar, A.D., Pross, J., Oschmann, W., van de Schootbrugge, B., Arz, H.W. (2010) Global enhancement of ocean anoxia during Oceanic Anoxic Event 2: A quantitative approach using U isotopes. Geology 38, 315–318.

; Brennecka et al., 2011

Brennecka, G.A., Herrmann, A.D., Algeo, T.J., Anbar, A.D. (2011) Rapid expansion of oceanic anoxia immediately before the end-Permian mass extinction. Proceedings of the National Academy of Sciences 108, 17631–17634.

; Dahl et al., 2014

Dahl, T.W., Boyle, R.A., Canfield, D.E., Connelly, J.N., Gill, B.C., Lenton, T.M., Bizzarro, M. (2014) Uranium isotopes distinguish two geochemically distinct stages during the later Cambrian SPICE event. Earth and Planetary Science Letters 401, 313–326.

; Elrick et al., 2016

Elrick, M., Polyak, V., Algeo, T.J., Romaniello, S., Asmerom, Y., Herrmann, A.D., Anbar, A.D., Zhao, L., Chen, Z.-Q. (2016) Global-ocean redox variation during the middle-late Permian through Early Triassic based on uranium isotope and Th/U trends of marine carbonates. Geology, G38585.1.

; Lau et al., 2016

Lau, K.V., Maher, K., Altiner, D., Kelley, B.M., Kump, L.R., Lehrmann, D.J., Silva-Tamayo, J.C., Weaver, K.L., Yu, M., Payne, J.L. (2016) Marine anoxia and delayed Earth system recovery after the end-Permian extinction. Proceedings of the National Academy of Science USA 113, 2360–2365.

), where anoxic water masses expanded over larger areas of the seafloor and caused negative δ²³⁸U excursions. Here, we use uranium isotopes to identify a transient global oxygenation episode during the radiation of animals in the Cambrian.

Our new δ²³⁸U data of carbonate-associated uranium comes from limestones collected from the Siberian Platform across the provisional Cambrian Stage 2–3 boundary (~521 to 520 million years ago) (Fig. 1), when animals that shed their exoskeleton (ecdysozoa) began to diversify (Maloof et al., 2010

Maloof, A.C., Porter, S.M., Moore, J.L., Dudas, F.O., Bowring, S.A., Higgins, J.A., Fike, D.A., Eddy, M.P. (2010) The earliest Cambrian record of animals and ocean geochemical change. Geological Society of America Bulletin 122, 1731–1774.

; Kouchinsky et al., 2012

Kouchinsky, A., Bengtson, S., Runnegar, B., Skovsted, C., Steiner, M., Vendrasco, M. (2012) Chronology of early Cambrian biomineralization. Geological Magazine 149, 221–251.

). A perturbation in the marine carbon cycle is expressed at this time as a large positive carbon isotope excursion recognised globally and in all studied sections; geological maps and stratigraphic sections are shown in Figures S-1 and S-2 (Maloof et al., 2010

Maloof, A.C., Porter, S.M., Moore, J.L., Dudas, F.O., Bowring, S.A., Higgins, J.A., Fike, D.A., Eddy, M.P. (2010) The earliest Cambrian record of animals and ocean geochemical change. Geological Society of America Bulletin 122, 1731–1774.

). This excursion serves as an important stratigraphic marker, although little is known about the biogeochemical significance of the event. The end-Stage 2 samples carry low δ²³⁸U values of –0.65 ‰, increasing stratigraphically to a value of –0.45 ‰ that approaches the modern oxygenated oceans, –0.39 ± 0.01 ‰ (Tissot and Dauphas, 2015

Tissot, F.L.H., Dauphas, N. (2015) Uranium isotopic compositions of the crust and ocean: Age corrections, U budget and global extent of modern anoxia. Geochimica et Cosmochimica Acta 167, 113–143.

), before again returning in two steps to –0.7 ‰. This positive δ²³⁸U excursion of +0.25 ‰ coincides with the positive carbon isotope excursion, suggesting that they are both linked to the changes in global seawater chemistry. The samples display no systematic correlation between the δ²³⁸U excursion and indicators of dolomitisation (Mg/Ca, dolomite), pore water redox conditions (total organic carbon content), diagenetic alteration (Mn/Sr, δ¹⁸O) detrital input (Al/Ca, clay content) and primary carbonate mineralogy (Sr/Ca; see Supplementary Information S2) that might produce such a positive δ²³⁸U excursion, offset from contemporaneous seawater. That said, our samples consist of (abiotic) micrite with (biotic) shells made of secondary calcite (Fig. S-3). The difference between abiotic and biotic precipitation of calcite δ²³⁸U is predicted to induce a ~0.1 ‰ offset from seawater (Chen et al., 2016

Chen, X., Romaniello, S.J., Herrmann, A.D., Wasylenki, L.E., Anbar, A.D. (2016) Uranium isotope fractionation during coprecipitation with aragonite and calcite. Geochimica et Cosmochimica Acta 188, 189–207.

). As we cannot determine the ratio of U derived from abiotic to biotic sources, we conclude that at this level of confidence, an overall positive δ²³⁸U trend is observed in the stratigraphy that we ascribe to secular changes of open marine δ²³⁸U in early Cambrian seawater. We note that our data set is limited to only one section and predict the same trend can be observed in other marine deposits with authigenic U enrichments.

Full size image | Download in Powerpoint

The positive excursion of seawater δ²³⁸U represents a global oxygenation period in the oceans that dramatically decreased the fraction (f_U) of total U burial occurring in anoxic marine settings. A simple isotope mass balance calculation constrains f_U from the δ²³⁸U of seawater (δ_SW) (see derivation in Supplementary Information S4). This calculation assumes an isotopically constant U input from rivers and a constant isotope fractionation between seawater and the U sinks. The modern ocean is at steady state with δ²³⁸U at –0.39 ± 0.01 ‰ and estimates for the average oceanic input range from –0.30 ‰ to –0.27 ‰ (Tissot and Dauphas, 2015

Tissot, F.L.H., Dauphas, N. (2015) Uranium isotopic compositions of the crust and ocean: Age corrections, U budget and global extent of modern anoxia. Geochimica et Cosmochimica Acta 167, 113–143.

; Noordmann et al., 2016

Noordmann, J., Weyer, S., Georg, R.B., Jons, S., Sharma, M. (2016) 238U/235U isotope ratios of crustal material, rivers and products of hydrothermal alteration: new insights on the oceanic U isotope mass balance. Isotopes in Environmental Health Studies 52, 141–163.

). If the anoxic proportion of global U burial (f_U) today is 12–25 % (Morford and Emerson, 1999

Morford, J.L., Emerson, S. (1999) The geochemistry of redox sensitive trace metals in sediments. Geochimica Et Cosmochimica Acta 63, 1735–1750.

; Dunk et al., 2002

Dunk, R.M., Mills, R.A., Jenkins, W.J. (2002) A reevaluation of the oceanic uranium budget for the Holocene. Chemical Geology 190, 45–67.

; Noordmann et al., 2016

Noordmann, J., Weyer, S., Georg, R.B., Jons, S., Sharma, M. (2016) 238U/235U isotope ratios of crustal material, rivers and products of hydrothermal alteration: new insights on the oceanic U isotope mass balance. Isotopes in Environmental Health Studies 52, 141–163.

), and anoxic settings impart a +0.5 ± 0.1 ‰ net isotope offset from overlying seawater, this implies that the average isotope fractionation between seawater and all other oxic U sinks (Δ_OTHER) is –0.02 ± 0.05 ‰. Given this formulation, we calculate that 68 ± 18 % of marine U burial before the Stage 2–3 event occurred in anoxic parts of the oceans when δ²³⁸U of seawater was –0.65 ‰. This scenario compares to the ocean state during the anoxic expansion associated with the end-Permian extinction (Brennecka et al., 2011

Brennecka, G.A., Herrmann, A.D., Algeo, T.J., Anbar, A.D. (2011) Rapid expansion of oceanic anoxia immediately before the end-Permian mass extinction. Proceedings of the National Academy of Sciences 108, 17631–17634.

). Further, the anoxic U burial fraction declined to 29 ± 12 % at the peak of the event and then afterward returned to a more reducing ocean state with 78 ± 20 % anoxic U burial (Fig. S-2). The stated uncertainties for f_U in the Cambrian are propagated errors from Δ_OTHER, Δ_ANOX, and δ_IN reflecting a range of plausible parameterisations of the marine U cycle (see sensitivity analysis in Supplementary Information S4 for details). These data consistently require the Cambrian Stage 2–3 event to have occurred when the oceans were far more reducing than today, culminating at an oxygenation state similar to the modern ocean.

Existing palaeoredox studies also point to ocean oxygenation during the Early Cambrian (Fig. 2). Specifically, the molybdenum isotope (δ⁹⁸Mo) record from shales and phosphorites show several fluctuations recorded in ~560 to 520 Myr old stratigraphic sections in China (Wille et al., 2008

Wille, M., Nagler, T.F., Lehmann, B., Schroder, S., Kramers, J.D. (2008) Hydrogen sulphide release to surface waters at the Precambrian/Cambrian boundary. Nature 453, 767–769.

; Wen et al., 2011

Wen, H., Carignan, J., Zhang, Y., Fan, H., Cloquet, C., Liu, S. (2011) Molybdenum isotopic records across the Precambrian-Cambrian boundary. Geology 39, 775–778.

; Xu et al., 2012

Xu, L., Lehmann, B., Mao, J., Nägler, T.F., Neubert, N., Böttcher, M.E., Escher, P. (2012) Mo isotope and trace element patterns of Lower Cambrian black shales in South China: Multi-proxy constraints on the paleoenvironment. Chemical Geology 318-319, 45–59.

; Chen et al., 2015

Chen, X., Ling, H.F., Vance, D., Shields-Zhou, G.A., Zhu, M., Poulton, S.W., Och, L.M., Jiang, S.Y., Li, D., Cremonese, L., Archer, C. (2015) Rise to modern levels of ocean oxygenation coincided with the Cambrian radiation of animals. Nature Communications 6, 7142.

; Kendall et al., 2015

Kendall, B., Komiya, T., Lyons, T.W., Bates, S.M., Gordon, G.W., Romaniello, S.J., Jiang, G., Creaser, R.A., Xiao, S., McFadden, K., Sawaki, Y., Tahata, M., Shu, D., Han, J., Li, Y., Chu, X., Anbar, A.D. (2015) Uranium and molybdenum isotope evidence for an episode of widespread ocean oxygenation during the late Ediacaran Period. Geochimica et Cosmochimica Acta 156, 173–193.

; Wen et al., 2015

Wen, H., Fan, H., Zhang, Y., Cloquet, C., Carignan, J. (2015) Reconstruction of early Cambrian ocean chemistry from Mo isotopes. Geochimica et Cosmochimica Acta 164, 1–16.

). The last positive δ⁹⁸Mo excursion is broadly correlated to the first appearance of trilobites in China, strong Mo enrichments and with the oxygenation event reported here (Dahl et al., 2010

Dahl, T.W., Hammarlund, E.U., Anbar, A.D., Bond, D.P.G., Gill, B.C., Gordon, G.W., Knoll, A.H., Nielsen, A.T., Schovsbo, N.H., Canfield, D.E. (2010) Devonian rise in atmospheric oxygen correlated to the radiations of terrestrial plants and large predatory fish. Proceedings of the National Academy of Sciences 107, 17911–17915.

; Xu et al., 2012

Xu, L., Lehmann, B., Mao, J., Nägler, T.F., Neubert, N., Böttcher, M.E., Escher, P. (2012) Mo isotope and trace element patterns of Lower Cambrian black shales in South China: Multi-proxy constraints on the paleoenvironment. Chemical Geology 318-319, 45–59.

; Chen et al., 2015

Chen, X., Ling, H.F., Vance, D., Shields-Zhou, G.A., Zhu, M., Poulton, S.W., Och, L.M., Jiang, S.Y., Li, D., Cremonese, L., Archer, C. (2015) Rise to modern levels of ocean oxygenation coincided with the Cambrian radiation of animals. Nature Communications 6, 7142.

; Wen et al., 2015

Wen, H., Fan, H., Zhang, Y., Cloquet, C., Carignan, J. (2015) Reconstruction of early Cambrian ocean chemistry from Mo isotopes. Geochimica et Cosmochimica Acta 164, 1–16.

; Jin et al., 2016

Jin, C., Li, C., Algeo, T.J., Planavsky, N.J., Cui, H., Yang, X., Zhao, Y., Zhang, X., Xie, S. (2016) A highly redox-heterogeneous ocean in South China during the early Cambrian (∼529–514 Ma): Implications for biota-environment co-evolution. Earth and Planetary Science Letters 441, 38–51.

). The coincident positive δ⁹⁸Mo and δ²³⁸U excursions point to a widespread oxygenation episode in the earliest Cambrian Stage 3 oceans where O₂-rich waters expanded and affected the Mo and U isotope composition of seawater as the overall burial fluxes of Mo and U into anoxic and euxinic settings decreased. We can predict the δ⁹⁸Mo trajectory of seawater from the oceanic δ²³⁸U trajectory, if we assume the anoxic burial proportions of total Mo and U burial (f_Mo, f_U) are correlated. We adopt a power law relationship f_Mo = f_U^α that satisfies f_Mo = f_U = 1 during extreme anoxia and f_Mo = f_U = 0 for extreme oxia. Using the modern ocean state as a calibration point, we find that α = 1.34 ± 0.38 (see Supplementary Information for details). Based on this relationship, average seawater δ⁹⁸Mo should have increased from 1.40 ‰ in the Cambrian Stage 2 to a peak at 2.0 ‰ before returning to 1.10 ‰ during the δ²³⁸U excursion. This prediction is in good agreement with the maximum values observed in the δ⁹⁸Mo record during this time interval (Lehmann et al., 2007

Lehmann, B., Nägler, T.F., Holland, H.D., Wille, M., Mao, J., Pan, J., Ma, D., Dulski, P. (2007) Highly metalliferous carbonaceous shale and Early Cambrian seawater. Geology 35, 403.

; Chen et al., 2015

Chen, X., Ling, H.F., Vance, D., Shields-Zhou, G.A., Zhu, M., Poulton, S.W., Och, L.M., Jiang, S.Y., Li, D., Cremonese, L., Archer, C. (2015) Rise to modern levels of ocean oxygenation coincided with the Cambrian radiation of animals. Nature Communications 6, 7142.

; Wen et al., 2015

Wen, H., Fan, H., Zhang, Y., Cloquet, C., Carignan, J. (2015) Reconstruction of early Cambrian ocean chemistry from Mo isotopes. Geochimica et Cosmochimica Acta 164, 1–16.

) (Fig. 2). Collectively, the Mo and U isotopes indicate a transient, rather than a persistent change in ocean oxygenation at the beginning of Stage 3. This implies that earlier positive δ⁹⁸Mo excursions (~2 ‰) in Terreneuvian phosphorite deposits (Wen et al., 2011

Wen, H., Carignan, J., Zhang, Y., Fan, H., Cloquet, C., Liu, S. (2011) Molybdenum isotopic records across the Precambrian-Cambrian boundary. Geology 39, 775–778.

), and perhaps in the latest Ediacaran (Kendall et al., 2015

Kendall, B., Komiya, T., Lyons, T.W., Bates, S.M., Gordon, G.W., Romaniello, S.J., Jiang, G., Creaser, R.A., Xiao, S., McFadden, K., Sawaki, Y., Tahata, M., Shu, D., Han, J., Li, Y., Chu, X., Anbar, A.D. (2015) Uranium and molybdenum isotope evidence for an episode of widespread ocean oxygenation during the late Ediacaran Period. Geochimica et Cosmochimica Acta 156, 173–193.

), represent episodic events rather than persistent ocean oxygenation. Similarly, detailed studies of the bottom water redox conditions in the Nanhua Basin, South China, suggest that oxygenated waters also invaded shallower part of the basin later in the Stage 3 (Jin et al., 2016

Jin, C., Li, C., Algeo, T.J., Planavsky, N.J., Cui, H., Yang, X., Zhao, Y., Zhang, X., Xie, S. (2016) A highly redox-heterogeneous ocean in South China during the early Cambrian (∼529–514 Ma): Implications for biota-environment co-evolution. Earth and Planetary Science Letters 441, 38–51.

). As such, the oxygenation history of the early Cambrian ocean appears more dynamic than previously thought (e.g., Dahl et al., 2010

Dahl, T.W., Hammarlund, E.U., Anbar, A.D., Bond, D.P.G., Gill, B.C., Gordon, G.W., Knoll, A.H., Nielsen, A.T., Schovsbo, N.H., Canfield, D.E. (2010) Devonian rise in atmospheric oxygen correlated to the radiations of terrestrial plants and large predatory fish. Proceedings of the National Academy of Sciences 107, 17911–17915.

; Sperling et al., 2013

Sperling, E.A., Frieder, C.A., Raman, A.V., Girguis, P.R., Levin, L.A., Knoll, A.H. (2013) Oxygen, ecology, and the Cambrian radiation of animals. Proceedings of the National Academy of Sciences 110, 13446–13451.

; Chen et al., 2015

Chen, X., Ling, H.F., Vance, D., Shields-Zhou, G.A., Zhu, M., Poulton, S.W., Och, L.M., Jiang, S.Y., Li, D., Cremonese, L., Archer, C. (2015) Rise to modern levels of ocean oxygenation coincided with the Cambrian radiation of animals. Nature Communications 6, 7142.

).

Full size image | Download in Powerpoint

The oxygenation event coincides with global changes in the marine carbon and sulphur cycles. We also report sulphur isotope data from carbonate-associated sulphate (δ³⁴S_CAS) from three distinct stratigraphic sections in Siberia that show a similar systematic positive δ³⁴S_CAS isotope excursion coinciding with the δ²³⁸U and δ¹³C excursions (Fig. 1). The simultaneous excursions across the Cambrian Stage 2-3 boundary suggest the C, S and U cycles responded to the same global biogeochemical event. Similar parallel positive carbon and sulphur isotope excursions during the late Cambrian SPICE event (Gill et al., 2011

Gill, B.C., Lyons, T.W., Young, S.A., Kump, L.R., Knoll, A.H., Saltzman, M.R. (2011) Geochemical evidence for widespread euxinia in the later Cambrian ocean. Nature 469, 80–83.

; Dahl et al., 2014

Dahl, T.W., Boyle, R.A., Canfield, D.E., Connelly, J.N., Gill, B.C., Lenton, T.M., Bizzarro, M. (2014) Uranium isotopes distinguish two geochemically distinct stages during the later Cambrian SPICE event. Earth and Planetary Science Letters 401, 313–326.

) and the Botoman Sinsk event were interpreted to represent a period of enhanced organic carbon and pyrite burial (Zhuravlev and Wood, 1996

Zhuravlev, A., Wood, R. (1996) Anoxia as the cause of the mid-Early Cambrian (Botomian) extinction event. Geology 24, 311–314.

). While these events are linked to expanding ocean anoxia and animal extinctions, the positive δ²³⁸U excursion reported here reveals a distinct driver for the environmental change.

The Cambrian Stage 2–3 event coincides with the onset of the first major diversification of arthropods, which predates by a few million years the first appearance of macrozooplankton and suspension-feeding anomalocarids (Hou, 2004

Hou, X.G., Aldridge, R.J., Bergström, J., Siveter, D.J., Siveter, D.J., Feng, X.H. (2004) The Cambrian fossils of Chengjiang, China - the flowering of early animal life. Blackwell Publishing, Malden, Oxford.

; Stein et al., 2009

Stein, M., Peel, J., Siveter, D., Williams, M. (2009) Isoxys (Arthropoda) with preserved soft anatomy from the Sirius Passet Lagerstätte, lower Cambrian of North Greenland. Lethaia 43, 258–265.

; Vinther et al., 2014

Vinther, J., Stein, M., Longrich, N.R., Harper, D.A. (2014) A suspension-feeding anomalocarid from the Early Cambrian. Nature 507, 496–499.

). Therefore, we consider that this oxygenation episode reflects the first invasion of larger zooplankton in the pelagic zone that triggered an increase in the sinking rate and compaction of organic matter. Today, the export of organic matter from the photic zone (<200 m depth) — a process referred to as the biological pump — occurs as sinking molts, faecal matter, carcasses, and skeletons (Alldredge et al., 1993

Alldredge, A.L., Passow, U., Logan, B.E. (1993) The abundance and significance of a class of large, transparent organic particles in the ocean. Deep Sea Research I 40, 1131–1140.

; Hedges and Keil, 1995

Hedges, J.L., Keil, R.G. (1995) Sedimentary organic matter preservation: an assessment and speculative synthesis. Marine Chemistry 49, 81–115.

). The sinking velocity of a particle in the ocean is a quadratic function of its size. Therefore, a small increase in the mean size of particulate organic matter would have caused the rates of water column remineralisation to decrease, so that less O₂ was consumed in the water column and a more gentle O₂ gradient was established below the photic zone (Fig. 3, Meyer et al., 2016

Meyer, K.M., Ridgwell, A., Payne, J.L. (2016) The influence of the biological pump on ocean chemistry: implications for long-term trends in marine redox chemistry, the global carbon cycle, and marine animal ecosystems. Geobiology 14, 207–219.

). Consequently, more organic matter would have been exported to the seafloor resulting in enhanced rates of organic carbon and pyrite burial.

Full size image | Download in Powerpoint

To quantify the potential consequences of faster sinking rates on organic carbon export (J_RAIN) onto the seafloor, we utilise the δ¹³C and δ³⁴S data and a simple biogeochemical model for the coupled marine C and S cycles. Organic carbon export fuels both organic carbon burial (J_ORG = J_RAIN - J_REMIN) and remineralisation in sediments (J_REMIN). In this model, pyrite burial is proportional to the remineralisation flux (J_PY = α · J_REMIN) because microbial sulphate reduction accounts for the major part of organic remineralisation in sediments and a fraction of its byproduct, sulphide, reacts with available Fe compounds to form pyrite (see Supplementary Information for model details). With this formulation (α constant), we find that a 3–6 fold increase in the organic carbon rain rate is sufficient to increase δ¹³C (0 to 3 ‰) and δ³⁴S (28 to 40 ‰) simultaneously. This can be achieved with only a 1.7–2.5 fold increase in the mean of the size distribution of sinking organic matter particulates, assuming similar density and shape for the particulates before and after the invasion of larger zooplankton. We consider this as a minimum estimate, if a greater portion of larger faecal pellets do not aggregate (Butterfield, 1997

Butterfield, N.J. (1997) Ecology and the Proterozoic-Phanerozoic Transition. Paleobiology 23, 247–262.

). However, it shows that even a modest size increase potentially influences the global biogeochemical cycles.

The evolutionary history of animals over the Neoproterozoic-Cambrian transition suggests a stepwise increase in their overall size and their digestive tracts, which would have enhanced the biological pump (Logan et al., 1995

Logan, G.A., Hayes, J.M., Hieshima, G.B., Summons, R.E. (1995) Terminal Proterozoic reorganization of biogeochemical cycles. Nature 376, 53–56.

; Butterfield, 2009

Butterfield, N.J. (2009) Oxygen, animals and oceanic ventilation: an alternative view. Geobiology 7, 1–7.

; Lenton et al., 2014

Lenton, T.M., Boyle, R.A., Poulton, S.W., Shields-Zhou, G.A., Butterfield, N.J. (2014) Co-evolution of eukaryotes and ocean oxygenation in the Neoproterozoic era. Nature Geoscience 7, 257–265.

). Biomarker evidence suggests a fundamental shift in the preservation state of marine organic matter with abundant faecal matter in sedimentary rocks younger than ~517 Ma compared to similar rocks older than ~565 Ma (Logan et al., 1995

Logan, G.A., Hayes, J.M., Hieshima, G.B., Summons, R.E. (1995) Terminal Proterozoic reorganization of biogeochemical cycles. Nature 376, 53–56.

). Although the abundance of pelagic animal fauna through the Cambrian Stage 2–3 interval is not known, it is apparent that the maximum size of pelagic heterotrophic organisms and their digestive tracts increased by three orders of magnitude from the Ediacaran into the Cambrian. Heterotrophic consumers such as micro- and mesozooplankton (20–200 µm) evolved in the latest Ediacaran (635–542 Ma) (Perrier et al., 2015

Perrier, V., Williams, M., Siveter, D.J. (2015) The fossil record and palaeoenvironmental significance of marine arthropod zooplankton. Earth-Science Reviews 146, 146–162.

). The fossil record indicates that animals of probable chaetognath affinity ('protoconodonts') a few millimetres in size or larger swam in the early Fortunian ocean (>535 Ma), and that large macrozooplankton (i.e. bivalved arthropod Isoxys) up to 45 mm had appeared in the earliest Stage 3 (~520 Ma) (Ivantsov, 1990

Ivantsov, A.Y. (1990) The first findings of phyllocarida in the lower Cambrian deposits of Yakutia Russian SFSR USSR. Paleontologicheskii Zhurnal, 130–132.

; Hou, 2004

Hou, X.G., Aldridge, R.J., Bergström, J., Siveter, D.J., Siveter, D.J., Feng, X.H. (2004) The Cambrian fossils of Chengjiang, China - the flowering of early animal life. Blackwell Publishing, Malden, Oxford.

; Stein et al., 2009

Stein, M., Peel, J., Siveter, D., Williams, M. (2009) Isoxys (Arthropoda) with preserved soft anatomy from the Sirius Passet Lagerstätte, lower Cambrian of North Greenland. Lethaia 43, 258–265.

). However, it says nothing about when various types of zooplankton became ecologically important. Rather, we suggest that a stepwise increase in the size of zooplankton is a plausible trigger for the reorganisation of Earth's marine biogeochemical cycles and, consequently, the earliest Cambrian Stage 3 oxygenation episode. By implication, earlier steps in animal evolution may have also led to oxygenation events preserved in the geochemical record during the Ediacaran–Cambrian transition (see below).

Given that the geochemical evidence suggests that the Cambrian Stage 2–3 oxygenation episode was brief, a stabilising feedback must have acted to counterbalance marine oxygenation. Current models for the evolving biogeochemical cycles in the Cambrian have not included such a rapid feedback mechanism (Bergman et al., 2004

Bergman, N.M., Lenton, T.M., Watson, A.J. (2004) COPSE: A new model of biogeochemical cycling over Phanerozoic time. American Journal of Science 304, 397–437.

; Berner, 2006

Berner, R.A. (2006) GEOCARBSULF: A combined model for Phanerozoic atmospheric O2 and CO2. Geochimica et Cosmochimica Acta 70, 5653–5664.

). Evidence from bioturbation indices (Mangano and Buatois, 2014

Mangano, M.G., Buatois, L.A. (2014) Decoupling of body-plan diversification and ecological structuring during the Ediacaran-Cambrian transition: evolutionary and geobiological feedbacks. Proceedings of the Royal Society B. Biological Science 281, 20140038.

) and Earth system modelling (Canfield and Farquhar, 2009

Canfield, D.E., Farquhar, J. (2009) Animal evolution, bioturbation, and the sulfate concentration of the oceans. Proceedings of the National Academy of Sciences USA 106, 8123–8127.

; Boyle et al., 2014

Boyle, R.A., Dahl, T.W., Dale, A.W., Shields-Zhou, G.A., Zhu, M., Brasier, M.D., Canfield, D.E., Lenton, T.M. (2014) Stabilization of the coupled oxygen and phosphorus cycles by the evolution of bioturbation. Nature Geoscience 7, 671–676.

) suggest that the Cambrian Stage 3 oxygenation episode occurred during an apparent atmospheric pO₂ decline resulting from increased mixing of marine sediments by animals. Since the early Cambrian Stage 2 (~530 Ma), animals had evolved the ability to burrow deeper into sediments acting to lower marine P availability, organic productivity, organic carbon burial, and hence the main source of atmospheric pO₂ (Boyle et al., 2014

Boyle, R.A., Dahl, T.W., Dale, A.W., Shields-Zhou, G.A., Zhu, M., Brasier, M.D., Canfield, D.E., Lenton, T.M. (2014) Stabilization of the coupled oxygen and phosphorus cycles by the evolution of bioturbation. Nature Geoscience 7, 671–676.

). The emerging larger pelagic fauna accelerated this feedback mechanism, since the enhanced biological pump would have both increased the food supply and food quality for benthic heterotrophic organisms and led to more fully oxygenated conditions in the water column below the photic zone, thereby opening new ecospace for sediment-mixing animals over wider areas of the continental shelves. Subsequently, this focusing of organic matter at the seafloor increased overall rate of bioturbation, organic carbon remineralisation and oxygen consumption over larger areas of the seafloor and, ultimately, a decline in atmospheric pO₂ that again limited the size of the oxygenated zones in the oceans (Fig. 3). It is notable that for this sequence of feedbacks to respond over the time scale of the Cambrian Stage 2–3 oxygenation episode (1.3 ± 0.8 Myr (Maloof et al., 2010

Maloof, A.C., Porter, S.M., Moore, J.L., Dudas, F.O., Bowring, S.A., Higgins, J.A., Fike, D.A., Eddy, M.P. (2010) The earliest Cambrian record of animals and ocean geochemical change. Geological Society of America Bulletin 122, 1731–1774.

), see calculation in the Supplementary Information), atmospheric O₂ inventory must have been significantly smaller than today in order to produce an excursion of the right duration. We derive an order of magnitude estimate for the atmospheric pO₂ level at the Cambrian Stage 2-3 boundary of between 4 ± 2 and 7 ± 4 atm % from the duration of the falling limb isotope excursions, assuming the global burial rate of marine organic carbon was the same as today. We also assume that anoxia returns as atmospheric pO₂ levels decline due to less organic carbon burial over the course of ~1/4 to ~1/2 the duration of the full δ¹³C excursion (~325 ± 200 and ~650 ± 400 kyr). This atmospheric pO₂ estimate scales linearly with global organic carbon burial flux, and requires that the acceleration of the biological pump (by faecal pellets) and the subsequent migration of the sediment-dwelling taxa are essentially instantaneous (<<100 kyr). This atmospheric pO₂ level is well above the metabolic need for some animals (Pasteur limit ~0.2 atm %) (Mills and Canfield, 2014

Mills, D.B., Canfield, D.E. (2014) Oxygen and animal evolution: Did a rise of atmospheric oxygen "trigger" the origin of animals? Bioessays 36, 1145–1155.

), but not sufficiently high to oxygenate the deep oceans permanently (Lyons et al., 2014

Lyons, T.W., Reinhard, C.T., Planavsky, N.J. (2014) The rise of oxygen in Earth's early ocean and atmosphere. Nature 506, 307–315.

) and it conforms with the idea that animal ecosystems could have become self-limiting in terms of determining the size of the habitable ecospace in the oceans (Sperling et al., 2013

Sperling, E.A., Frieder, C.A., Raman, A.V., Girguis, P.R., Levin, L.A., Knoll, A.H. (2013) Oxygen, ecology, and the Cambrian radiation of animals. Proceedings of the National Academy of Sciences 110, 13446–13451.

; Boyle et al., 2014

Boyle, R.A., Dahl, T.W., Dale, A.W., Shields-Zhou, G.A., Zhu, M., Brasier, M.D., Canfield, D.E., Lenton, T.M. (2014) Stabilization of the coupled oxygen and phosphorus cycles by the evolution of bioturbation. Nature Geoscience 7, 671–676.

).

top

Acknowledgements

We thank N. Jensen, T. Balic-Zunic, H. Almind, N.J.B. Kristensen (UCPH) for laboratory assistance, S. Bengtson (SMNH) for field assistance in 1996 with AK, and T.J. Algeo, C. Li and one anonymous reviewer for insightful comments. Funding for this project was provided by a grant from the VILLUM Foundation (VKR023127) to TWD, and from the Danish Agency for Science, Technology and Innovation (grant number 12-125692) to JNC as well as grants from the Danish National Research Foundation (grant number DNRF97) and from the European Research Council (ERC Consolidator grant agreement 616027-STARDUST2ASTEROIDS) to MB. Fieldwork was financially supported by grants from the Royal Swedish Academy of Sciences (KVA), the Swedish Natural Science Research Council (NFR), and the Swedish Research Council (VR).

Editor: Liane G. Benning

top

Author Contributions

AK collected the samples. TWD conceived the idea, designed the study, interpreted the data and created the models. TWD, JNC, BCG, SFM and MB analysed the samples and wrote the manuscript with significant contributions from all co-authors.

top

References

top

Supplementary Information

S1. Methods, Samples and Geological Settings

Samples and geological settings
The material was obtained by AK during a field expedition in 1996 along the 1) Bolshaya-Kuonamka River (Anabar Uplift, Siberian Platform) (Kouchinsky et al., 2001) and 2) Malaya-Kuonamka River (Anabar Uplift) (Kouchinsky et al., 2001), and by Vladimir Pavlov and Vladimir Vodovozov in 2004 (Institute of Physics of the Earth, Moscow) from 3) Sukharikha River (Igarka region, western margin of the Siberian Platform) (Kouchinsky et al., 2007; see Fig. S-1). The stratigraphic sections at both localities consist of open-shelf marine facies (Rowland et al., 1998). The Tommotian-Atdabanian (Cambrian Stage 2–3 in Siberia) interval is found within in the Krasnoporog Formation (Rowland et al., 1998) and the Emyaksin (Kouchinsky et al., 2001) Formations at the Sukharikha River and the sections from the Anabar Uplift, respectively. The Krasnoporog Formation represents shallow shelf environments consisting primarily of burrow-mottled thick-bedded pink and grey lime mudstones with two archaeocyath-rich intervals at the base of the Atdabanian portion of the section (Rowland et al., 1998). The Emyaksin Formation dominantly consists of argillaceous lime mudstones, although the clay mineral content is low (<10 wt. %, Extended Data 1 and 2). The Emyaksin Formation belongs to the Yudoma-Olenek facies region (or belt) of the lower Cambrian deposits of the Siberian Platform (e.g., Varlamov et al., 2008 and references therein). A detailed interpretation of the depositional environments is provided for the analogous Erkeket and Medvezhya Formations of the northern part of the Yudoma-Olenek facies region (Kaufman and Knoll, 1995; Kaufman et al., 1996).

The sections are correlated to one another and to other sections globally via chemo- and biostratigraphy (Fig. S-2; Kouchinsky et al., 2012, 2015). Briefly, the Tommotian-Atdabanian is biostratigraphically defined by archaeocyaths located at 12–19 m above the base of the Krasnoporog Formation (Kouchinsky et al., 2012). This broadly coincides with where δ¹³C crosses 0 ‰ on the rising limb of the positive δ¹³C excursion (peak 'IV'). This δ¹³C excursion, which reaches +3 to +4 ‰ at its peak, is a globally recognised chemostratigraphic marker and as bracketed by ²⁰⁶Pb–²³⁸U zircon dates from two ash beds in Morocco (520.93 ± 0.14 Ma and 517 ± 1.5 Ma) (Maloof et al., 2010b). No erosional surface was revealed in the Krasnoporog Formation during field studies (Kouchinsky et al., 2007). At a constant sedimentation rate, we calculate the carbon isotope excursion (as defined by the two points where δ¹³C crosses 0 ‰) spans 1.3 ± 0.8 Myr. The δ³⁴S and δ²³⁸U isotope excursions presented here both cover a similar time span. Profallotaspis jakutensis, one of earliest trilobites, occurs ~10 metres above the point where the δ¹³C crosses -0.5 ‰ on the rising limb of the carbon isotope excursion at Zhurinskii Mys, Lena River (Rozanov et al., 2008).

Full size image | Download in Powerpoint

Full size image | Download in Powerpoint

Full size image | Download in Powerpoint

Methods
δ²³⁸U_CAU. We calibrated (see Supplementary Information) and utilised a soft leaching procedure to extract mainly carbonate-associated U (CAU) from our samples and to avoid U associated with authigenic carbonate fluorapatite and ferric oxyhydroxide-associated P (Jeppsson et al., 1999), or any detrital or organic-bound U. Samples were cut in ~10–20 g pieces and secondary veins, if present, were removed using a water-cooled saw. The rock surface was first leached in 10 % acetic acid to remove surface contaminants, dried and then crushed in a shatter box to a fine powder. For analysis of carbonate-associated uranium, approximately 1–3 g of the fine powder were weighed out in 50 mL centrifuge tubes and leached using a mild leaching procedure: 40 mL 10 % acetic acid were added 10 mL at a time to dissolve the carbonate preferentially. All labware in contact with dissolved sample material were acid-cleaned in the same acid prior to use. The mixture was shaken several times and left overnight at 50 °C in a warm bath, then ultrasonicated and centrifuged for 5 min at 2500 rpm before the supernatant was taken up by syringe and filtered through 0.22 µm filter to separate the liquid from the solids. Following the removal of the supernatant, the residue was washed with 5 mL 10 % acetic acid, centrifuged before the supernatant was again transferred to the leachate to ensure maximum transfer of the carbonate-associated uranium. Still, calcite is not quantitatively extracted from our sample powders. Extraction yields were 57 ± 19 % as calculated from the bulk calcite contents (83 ± 9 wt. % CaCO₃; Extended Data 2) and the Ca concentration in the leachates (19 ± 4 wt. % Ca).

U and Th concentrations were measured in a small fraction (~1 %) of the leachate on a Thermo X-Series 2, quadrupole inductively coupled mass spectrometer (q-ICPMS). U isotope analyses were performed on the remaining fraction (~99 %) or a smaller aliquot with 60–600 ng U. The leachate was transferred to 60 mL Teflon beakers dried down on a hot plate and redissolved in 7 N nitric acid in a 16 mL Teflon beaker. Samples were spiked prior to chemical purification using the IRMM-3636 ²³³U–²³⁶U double spike (Condon et al., 2010) and left for >48 hours on a hotplate first in hydrochloric acid and later in nitric acid. This allowed the U from the spike to equilibrate isotopically with U in the sample and ensure all U compounds were in the same chemical form prior to ion chromatographic separation. Chemical purification of U and its isotopic analysis were conducted following the procedures outlined in Connelly et al. (2012) using a Neptune Plus Multi-Collector ICPMS. The spike/sample ratio was kept constant for all samples and standards during each analytical session. We estimate the long-term reproducibility and accuracy of our ²³⁸U/²³⁵U data to be ±0.03 ‰ (2SD). For example, BCR-2 (whole rock) and IAPSO seawater are found to be δ²³⁸U = –0.271 ± 0.017 ‰ and –0.372 ± 0.022 ‰ (2SE, n =5), respectively.

δ³⁴S_CAS. The sulphur isotope composition of carbonate-associated sulphate (CAS) was obtained using a traditional approach (Wotte et al., 2012). In summary, 3–8 g of the fine sample powder were treated with a NaCl solution, followed by several distilled water rinses to remove any soluble sulphates that might have been present in the samples. The samples were then dissolved using 4 N HCl. The leachates were centrifuged and vacuum filtered (45 µm) to remove any insoluble portion of the sample, before ~100 mL saturated BaCl₂ solution were added to allow sulphate to precipitate as BaSO₄ from the solution. We let samples sit for at least three days to ensure complete precipitation. The BaSO₄ was separated from the remaining solution via filtration and allowed to dry. The ³⁴S/³²S isotope ratios were analysed through on-line combustion using an Elementar Isotope Cube elemental analyser coupled to Isoprime 100 continuous-flow stable isotope ratio mass spectrometer at the Virginia Polytechnic Institute and State University. Samples are reported in standard delta notation (δ³⁴S) as per mille deviations from Vienna Canon Diablo Trolite (VCDT) and were calibrated to this scale with international isotope standards NBS-127, IAEA SO-5 and IAEA SO-6. Long-term reproducibility of the δ³⁴S measurement to be <0.2 ‰ (1 SD) through replicates of international isotope standards listed above. We estimate the external reproducibility of our δ³⁴S data to be better than 1 ‰ based on with replicate dissolutions of two samples (DM4-12, DM4-36).

S2. Additional Geochemical and Mineralogical Characterisation

The concentration of Mg, Al, Fe and Ca in the samples was measured in the same diluted acetic leachates that carbonate-associated uranium was obtained (Extended Data 1). The low Mg, Al and Fe concentration relative to Ca indicates that dolomite and clay are not dissolved. This is fully consistent with mineralogical characterisations showing that samples contain mostly calcite and only subordinate proportions of other mineral phases, i.e. <5 wt. % clay and dolomite (Extended Data 2).

The total organic carbon (TOC) content was measured from the residue after the 4 N HCl leach in CAS extractions using an FLASH 2000 element analyser coupled to Thermo Finnigan Delta V Plus continuous-flow mass spectrometer. All samples contain low TOC (<0.16 wt. %, median 0.03 wt. %) (Extended Data 3).

Full size image | Download in Powerpoint

Correlations
There is no systematic correlation between the δ²³⁸U excursion and geochemical indicators for dolomitisation, pore water redox conditions, detrital input and primary carbonate mineralogy (Fig. S-5, Table S-1). There is, however, a good correlation between δ²³⁸U and δ¹³C (R = 0.82, p = 0.005, Table S-1).

Full size image | Download in Powerpoint

	R	p-value
δ238U vs. U	0.12	0.71
δ238U vs. U/Th	0.15	0.67
U vs. U/Th	0.96	8.11E-07
δ238U vs. δ13C	0.82	0.005
δ238U vs. Mg/Ca	0.52	0.15
δ238U vs. TOC	-0.31	0.41
δ238U vs. δ18O	0.68	0.04
δ238U vs. Mn/Sr	0.14	0.69
δ238U vs. Al/Ca	#N/D	#N/D
δ238U vs. Sr/Ca	0.5	0.14
δ238U vs. calcite	-0.66	0.07
δ238U vs. quartz	0.69	0.05
δ238U vs. dolomite	0.38	0.34
δ238U vs. ankerite	-0.68	0.16
δ238U vs. chlorite	0.14	0.76
δ238U vs. mica	0.21	0.61
δ238U vs. K-feldspar	0.44	0.27
U/Th vs. δ13C	0.15	0.67
U/Th vs. Mg/Ca	0.49	0.15
U/Th vs. TOC	-0.1	0.78
U/Th vs. δ18O	0.59	0.07
U/Th vs. Mn/Sr	-0.13	0.71
U/Th vs. Al/Ca	#N/D	#N/D
U/Th vs. Sr/Ca	0.16	0.65
U/Th vs. calcite	-0.13	0.75
U/Th vs. quartz	0.04	0.92
U/Th vs. dolomite	0.3	0.44
U/Th vs. ankerite	-0.24	0.63
U/Th vs. chlorite	0.01	0.98
U/Th vs. mica	-0.13	0.73
U/Th vs. K-feldspar	0.06	0.88

Download in Excel

S3. Methods - Calibration of the CAU Extraction Procedure

The extraction of carbonate-associated uranium was calibrated using a four-step sequential leaching procedure. The extraction procedure is designed to minimise extraction of uranium associated with carbonate fluoroapatite, ferric oxyhydroxide-associated P, silicates and organic-bound U, while maximising the yield of calcium carbonate-associated uranium. Phosphate minerals resist dissolution in mild acetic acid at pH > 3.6, where limestone dissolves (Jeppsson et al., 1999). Silicate-bound U and organic-bound U dissolve in hydrofluoric acid and nitric acid, respectively. Existing extraction procedures utilise 0.5 M HCl leaching instead of acetic, so we explored the following leaching procedure:

Sequential extraction
Step 1: 1.7 M (10 %) acetic acid (pH > 3.6, T = 50 °C, overnight)
Step 2: 0.5 M HCl (T = 50 °C, overnight)
Step 3: 2 M HCl (T = 50 °C, overnight)
Step 4: 28 M HF + 7 M HNO₃ (T = 50 °C, overnight)

The U distribution in two end-member reference samples with variable phosphate contents were compared, including a chalk (~100 % CaCO₃, 0 wt. % P₂O₅) from the Cretaceous–Paleogene boundary at Stevns, Denmark, and phosphoritic sedimentary rock (8.5 wt. % P₂O₅) from the Middle Cambrian Alum Shale formation, Sweden.

The samples were crushed to fine powders and analysed in duplicates. The U concentration in the leachates was measured using ICP-MS. Error bars represent 1 SD of replicate analyses.

U (ppb)	KTChalk	MCPhos
	(100 % CaCO3)	(8.5 wt. % P2O5)
10 % acetic acid	70 ± 4	440 ± 5
0.5 M HCl	19 ± 1	872 ± 423
2 M HCl	4 ± 0	133 ± 49
HF + HNO3	14 ± 0	994 ± 8
Total	107 ± 4	2439 ± 426

Download in Excel

Most of the U in the chalk sits in the acetic acid extractable fraction, whereas the majority of U in the phosphorite-containing sediment sits in the HCl extractable fraction and HF–HNO3 extractable portions. This result is expected since the U content in phosphate minerals is substantially higher (10–100 ppm) in the phosphate minerals than in the limestone (0.1–1 ppm).

The relative U proportions extracted with acetic and hydrochloric acid are shown in Table S-3 below.

U	KTChalk	MCPhos
10 % acetic acid	75 ± 4 %	30 ± 0 %
0.5 M HCl	20 ± 1 %	60 ± 29 %
2 M HCl	4 ± 0 %	9 ± 3 %
Total	100 %	100 %

Download in Excel

We favour an extraction based on mild acetic acid, since a substantial portion of phosphorite-bound U is extracted with 0.5 M HCl and it only comes with a loss of ~20 % of the remaining CAU.

Reproducibility
Repeated leaching experiments with acetic acid were performed with three natural limestone samples. Importantly, our acetic acid leaching procedure always yields the same d²³⁸U values even though the extraction yield sometimes varies. Therefore, we can conclude that the leaching procedure does not induce any isotope offset.

Replicate	Reaction time	U (ppb)	δ238U (‰)
SRM-1d (modern argillaceous limestone)
#1	16 hr	70	–0.09 ± 0.02
#2	16 hr	743	–0.08 ± 0.02
#3	16 hr	639	–0.11 ± 0.02
T1-13 (Cambrian limestone)
#3	30 min	298	–0.85 ± 0.02
#4	4 hr	292	–0.83 ± 0.02
#5	16 hr	281	–0.82 ± 0.03
T1-26.5 (Cambrian limestone)
#1	16 hr	102	not determined
#2	16 hr	155	–0.56 ± 0.02
#3	16 hr	154	–0.60 ± 0.03

Download in Excel

Low extraction yields when carbonate grains are protected
To explore the nature of variable extraction yields, we mixed KT chalk and MCPhos in distinct proportions to see if the presence of acid-resistable phases would protect carbonate grains from dissolving. Indeed, we observe that the extraction yields in the acetic fraction are systematically lower than predicted (35 % vs. 40 % in a 1:1 mixture and 24 % vs. 33 % in a 1:2 mixture) in the presence of MCPhos. Accordingly, a systematically higher yield was observed in the subsequent HCl extractable fraction (37 % vs. 27 % in a 1:1 mixture and 34 % vs. 30 % in a 1:2 mixture). Hence, we infer that acetic acid-resistant phases protected carbonate grains against dissolution. By implication, a similar phenomenon may explain the variable extraction yields observed in the repeated analyses of some samples (e.g., SRM-1d #1), if the sample powder is not fully mixed with the acid. Therefore, acid and sample powders are now carefully shaken during the extraction. In any case, the protection of carbonate grains does not induce any isotope fractionation.

Acid	Chalk	1:1†	1:2ß	MCPhos	1:1†	1:2ß
	measured				predicted
10 % acetic	65 %	35 %	24 %	18 %	40 %	33 %
0.5 M HCl	18 %	37 %	34 %	36 %	27 %	30 %
2 M HCl	4 %	5 %	12 %	5 %	5 %	5 %
HF + HNO3	13 %	24 %	3 %	41 %	28 %	33 %

^†Mixing ratio 1:1 corresponds to 1.0187 g KTChalk (108 ng U) + 0.0519 g MCPhos (127 ng U).
^ßMixing ratio 1:2 corresponds to 1.0993 g KTChalk (117 ng U) + 0.1138 g MCPhos (278 ng U).

Download in Excel

S4. Marine Molybdenum and Uranium Cycle Box Model

Uranium
A mass balance budget for the oceanic uranium cycle must include all the major sources and sinks (Tissot and Dauphas, 2015). The dominant U input to the oceans (J_IN) is via rivers. The sink fluxes (J_OUT) can be divided into two groups – a) sediments deposited under anoxic waters (J_ANOX) with large isotopic difference (Δ_ANOX ~ 0.50 ‰) and b) sediments deposited in other settings (J_OTHER) under oxygenated waters with less isotopic difference (Δ_OTHER ~ <0.10 ‰) relative to overlying seawater. At steady state, the following mass balance equations must apply:

Elemental mass balance at steady state:

J_IN = J_OUT

where

J_OUT = J_ANOX + J_OTHER

Isotopic mass balance at steady state can be written as:

δ_IN * J_IN = δ_OUT * J_OUT

The majority of oceanic U enters via rivers (≥80 %) and groundwater (≤20 %) (Dunk et al., 2002; Tissot and Dauphas, 2015). The weighted average riverine input carries uranium with δ²³⁸U = –0.34 ‰ (Andersen et al., 2016). However, two analyses from the Yangtze River (-0.72 and -0.68 ‰), which has a large discharge, significantly shift this average to a lower value. The reason for the low value is not known, and anthropogenic contamination does not appear to be the cause (e.g., phosphate fertiliser; Andersen et al., 2016). Previous global riverine estimates yielded –0.30 ‰ to –0.27 ‰ (Noordmann et al., 2016). Regardless, it appears that the riverine composition is indistinguishable from average continental crust at –0.30 ± 0.04 ‰ and –0.31 ± 0.05 ‰ (2 SE) (Noordmann et al., 2016). We adopt a value of –0.27 ‰, but also explore a range of values, which are listed in Table S-6.

δ_OUTJ_OUT = δ_ANOXJ_ANOX + δ_OTHERJ_OTHER

We define the fraction of total sedimentary U burial occurring in anoxic settings as:

ƒ_U = J_ANOX/J_OUT

Therefore, the isotope mass balance equation can be written as:

δ_OUT = δ_ANOXƒ_U + δ_OTHER (1 − ƒ_U)

The isotope composition of each sink, i, (δ_i) relates to the isotope composition of contemporaneous seawater (δ_SW) via the average isotopic difference (Δ_i) expressed during sediment burial in each sink:

δ_ANOX = δ_SW + Δ_ANOX

δ_OTHER = δ_SW + Δ_OTHER

Therefore, the isotope mass balance equation (Eq. S-6) relates the isotope composition of seawater to the fraction of anoxic U burial (f_U) assuming the other parameters are constant through time (δ_IN, Δ_ANOX, Δ_OTHER).

δ_IN = δ_SW + Δ_OTHER + (Δ_ANOX − Δ_OTHER) * ƒ_U

Eq. S-9 tells us that the U isotope composition of seawater at steady state is offset from input composition by the magnitude of the isotopic difference in other settings (Δ_OTHER) - and decreases by the magnitude of (Δ_ANOX - Δ_OX) as the proportion U burial into anoxic settings increases from 0 to 100 %:

δ_SW = δ_IN − Δ_OTHER − (Δ_ANOX − Δ_OTHER) * ƒ_U

All model parameters carry considerable uncertainty due to limited number of representative isotope analyses for each flux and uncertainty in the magnitude of the burial fluxes associated with each isotopically distinct pathway. However, all parameterisations consistently demonstrate a negative relationship between δ_SW and f_U (Fig. S-6).

Case	1	2	3	4	5	6
	Default	low δIN	hi ΔANOX	hi δIN	low ΔANOX	hi ΔOTHER
δIN	–0.27 ‰	–0.34 ‰	–0.27 ‰	–0.30 ‰	–0.27 ‰	–0.27 ‰
ΔOTHER	0.04 ‰	0.00 ‰	0.10 ‰	0.10 ‰	0.04 ‰	0.15 ‰
ΔANOX	0.50 ‰	0.50 ‰	0.60 ‰	0.60 ‰	0.40 ‰	0.40 ‰

Download in Excel

Full size image | Download in Powerpoint

Sensitivity analyses
The U isotope composition of seawater is particularly sensitive to U burial in anoxic settings due to the large U isotopic difference when sediment is buried in these settings: Δ_ANOX is 0.4–0.6 ‰ (Romaniello, 2012; Andersen et al., 2014; Noordmann et al., 2016).

The J_other flux incorporates U removal via reducing sediments, hydrothermal alteration and biogenic carbonate and silica. The modelled isotopic difference from seawater in these settings, Δ_OTHER, represents a weighted average of these burial fluxes.

Here, the fraction of total U burial in various settings (i) is denoted f_i and the associated isotopic difference from open-ocean seawater is Δ_i. Here, SAD stands for Sulphidic At Depth and includes both siliciclastic and carbonate sediments overlain by a fully oxic water column, but with hydrogen sulphide accumulating in pore waters of the sediment. Reductive U removal induces an U isotopic difference from seawater similar to both reducing siliciclastic (0.04 to 0.29 ‰) and carbonate sediments (0.2–0.4 ‰) that are deposited in high-productivity oxic settings (Noordmann et al., 2016). A similar isotopic difference (0.0–0.2 ‰) is also observed during low temperature basaltic alteration (Andersen et al., 2015), which has been estimated to account for 24 ± 12 % of total U removal from the oceans today (Dunk et al., 2002; Noordmann et al., 2016). These isotopic differences are small, but distinctly higher than those produced during carbonate precipitation in oxygenated environments with negligible isotopic difference (0.0 ± 0.1 ‰; Chen et al., 2016), and during removal with metalliferous oxides in the deep ocean (f_FeMn = 7 ± 5 %), which induces an isotopic offset in the opposite direction (–0.24 ‰; Noordmann et al., 2016).

A sensitivity analysis of Eq. S-10 using the above average values and uncertainty estimates for f_i and Δ_i yields a small value for Δ_OTHER = 0.09 ± 0.08 ‰. This value is indistinguishable from the value, Δ_OTHER = –0.02 ± 0.05 ‰, obtained from Eq. S-9 for the modern oceans at steady state with respect to input and anoxic burial fraction (δ_IN = 0.285 ± 0.015 ‰, δ_SW = 0.39 ± 0.01 ‰ and f_ANOX = 12–25 %). Uncertainty estimates represent propagated errors in Eqs. S-9 and S-10. The anoxic burial fraction is calculated using these values and uncertainties for Δ_ANOX, and δ_IN and Δ_OTHER = –0.02 ± 0.05 ‰.

The marine Mo budget can be written in terms of its major sources (rivers, >90 %) and three major sinks: a) euxinic sediments (here denoted anoxic), b) high-productivity oxic sediments that are sulphidic at depth (SAD), and c) oxygenated sediments containing ferromanganese oxides (OX), each of which has increasing isotope fractionation from seawater. The equations for the marine Mo cycle at steady state are derived in Dahl et al. (2011). Briefly, the elemental mass balance for the ocean implies that the input flux (J_IN) matches the total output flux (J_OUT):

J_IN = J_OUT

where the output flux consists of the three major sinks:

J_OUT = J_ANOX + J_SAD + J_OX

Isotope mass balance can be written as in Eq. S-3. Most Mo enters the ocean via rivers (~ 95 %; Miller et al., 2011). The weighted average riverine input carries molybdenum with δ⁹⁸Mo = 0.7 ‰ based on analyses of waters representing 22 % of global riverine discharge (Archer and Vance, 2008).

The isotope composition of the oceanic output is given by the isotope composition (δ_i) of the major sinks weighted according to the burial fraction of each sink (f_i = J_i/J_OUT). Here, we use f_Mo for the Mo burial fraction in anoxic settings.

δ_OUT = δ_ANOXƒ_Mo + δ_SADƒ_SAD + δ_OXƒ_OX

The magnitude of the isotopic difference associated with burial in each sink is here defined as a positive number (Δ_i = δ_SW - δ_i ), which increases from anoxic (euxinic) to oxic environments ( Δ_OX > Δ_SAD > Δ_EUX). Eqs. S-3 and S-12 combine to:

δ_SW = δ_IN + (1 − ƒ_Mo)Δ_OX − ƒ_SAD(Δ_OX − Δ_SAD)

Eq. S-13 states that seawater δ⁹⁸Mo will always be positively offset from riverine input. The maximum value of 3.5 ‰ (δ_IN + Δ_OX = 3.5 ‰) will be found in a fully oxygenated ocean, and a minimum value of 0.7 ‰ occurs in a fully anoxic (euxinic) ocean (in both cases, f_SAD = 0).

The combined Mo and U box model for the ocean
The anoxic burial fluxes of Mo and U are correlated, and the total anoxic burial fractions (f_U, f_Mo) yield 0 and 1 when all sediment burial occurs in oxic and anoxic (euxinic) settings, respectively. Hence, the correlation may be described by a power law with exponent a.

ƒ_Mo = ƒ^α_U

We constrain the value of a from the anoxic burial fraction obtained in the marine budget of the modern Mo and U cycles. Published estimates are f_Mo = 0.06–0.15 (Scott et al., 2008; Kendall et al., 2009) and f_U = 0.12–0.25 (Morford and Emerson, 1999; Dunk et al., 2002), so that a = ln(f_Mo)/log(f_U) = 1.34 ± 0.38 (errors propagated assuming parameter estimates are uncorrelated). Three examples of the functional relationships between the anoxic Mo and U burial fractions are shown in Figure S-7 and the consequences for seawater composition is shown in Figure S-8.

With parameter values listed in Table S-7, our calculations can simultaneously fit modern and Cambrian ocean δ²³⁸U and δ⁹⁸Mo constraints. We note that a = 1.15–1.24 best fits simultaneously the two ocean states, favouring the higher estimates for the modern anoxic Mo burial fraction (f_Mo = 0.12–0.14, if f_U = 0.18 today).

	Mo	U
δIN	0.7	-0.27
ΔOX	2.8	0.04
ΔANOX	0.0	0.50
ΔSAD	1.4	-
ƒOX	30 %	82 %
ƒSAD	58 %	-
ƒEUX	12 %	18 %
Predicted δSW	2.35	-0.393
Observed δSW	2.34 ± 0.10	-0.392 ± 0.005

(Note that all Δs are here defined positive, despite Mo and U isotope fractionations have opposite signs).

Download in Excel

Full size image | Download in Powerpoint

Full size image | Download in Powerpoint

Full size image | Download in Powerpoint

S5. Marine Carbon and Sulphur Cycle Box Model

We present a simple model to quantify the organic carbon loading onto the seafloor and its influence on seawater δ¹³C and δ³⁴S.

The isotopic and elemental mass balance equations for carbon and sulphur in a 1-box ocean is given as a function of time by the long-term imbalance between sources and sinks (Eqs. 3a and 3b in Kurtz, 2003). The δ¹³C of seawater (δ_C) is a function of the global oceanic C inventory (M), oceanic C input flux (J_IN) and its isotope composition (δ_{C, IN}) as well as the organic proportion of total marine C burial (f_ORG) and average isotopic difference (Δ_BIO) between δ¹³C of buried organic matter and seawater.

δ_C = δ_C,IN − M/J_IN * d(δ_C)/dt + ƒ_ORG * Δ_BIO

The equation reduces to the steady state equation at times when the δ¹³C of seawater is not changing in time: d(δ_C)/dt = 0.

	Modern alue	Explored range	Explanation	Ref
α	0.0072	x (3-30)	Ratio of pyrite formation to organic carbon remineralisation rate	B07
JRAIN	192 · 1012	¥	mol/yr. Modern organic rain rate
JORG	26 · 1012	¥	mol/yr. Modern marine organic C burial rate	B07
JLAND	26 · 1012	ß	mol/yr. Modern terrestrial organic C burial rate
JEVAP	2.1·1012	x (0.2 - 1)	mol/yr. Cenozoic sulphate burial rate	CF09†
JCARB	216·1012	x (0.2 - 1)	mol/yr. Modern carbonate burial rate	BB12
δIN.C	–5 ‰	–5 ± 1 ‰	δ13C of oceanic input
ΔC	26 ‰	24-30 ‰	Average δ13C offset between seawater and buried organic matter
δin.S	8 ‰	4-16 ‰	δ34S of oceanic input
ΔS	35 ‰	35-50 ‰	Average δ34S offset between seawater and buried pyrite
SR	40 · 1012		mol/yr. Modern organic burial rate	B07
SO42-	28	2-3	mM. Modern oceanic sulphate level
JPY	1.2 · 1012	¥	mol/yr. Modern pyrite burial rate	B07

^† The size of the modern oxidising C and S sinks were established from the isotope-derived f-constraints (Burdige, 2007; Saltzman and Thomas, 2012), suggesting ƒ_ORG = 0.192 and ƒ_PY = 0.36 leads to steady state seawater values of δ¹³C = 0 ‰ and δ³⁴S = 21 ‰, respectively. We assume that the modern terrestrial organic carbon burial equals that of marine organic carbon burial (Bergman et al., 2004), and neglect terrestrial pyrite burial.
^¥ Model results are given in Tables S-9 to S-14
^ß Terrestrial organic carbon burial (e.g., coal) is assumed equal to marine organic carbon (Bergman et al., 2004). Ref. B07 = Burdige (2007), Ref CF09 = Canfield and Farquhar (2009), Ref BB12 = Berner and Berner (2012).

Download in Excel

δ_C = δ_C,IN + ƒ_ORG * Δ_BIO

The C and S cycles are instantaneously at steady state before the isotope excursion, at the peak of the isotope excursion, and after the isotope excursion, when d(δ_C)/dt = 0 and d(δ_S)/dt = 0. In the following, we use the constraints on f_ORG and f_PY at these time points to explore how enhanced organic carbon loading on the seafloor could have simultaneously affected the marine C and S cycles.

Carbon isotopes
The C isotope composition of seawater is a function of the fraction of marine organic carbon burial to total C burial (f_ORG = J_ORG / [J_CARB + J_ORG]), where the remainder C is buried as marine carbonate with same isotope composition as average seawater.

ƒ_ORG = J_ORG / (J_CARB + J_ORG)

For example, assuming modern values for δ_in and Δ_C (Table S-8) a change in seawater δ¹³C from 0 to 3 ‰ implies that f_ORG increases from 19 % to 31 % (+60 % relative increase).

Marine organic carbon burial is related to organic carbon rain rate measured at the seafloor (J_RAIN) and the fraction of this organic carbon that is remineralised inside the sediments (f_REMIN):

J_ORG = J_RAIN * (1 - ƒ_REMIN)

Although, global estimates for organic burial rates vary by more than an order of magnitude (13-217 · 10¹² mol/yr), Burdige (2007) revised these estimates and provided a best estimate of J_ORG = 26 · 10¹² mol/yr and depth-integrated rates of organic carbon remineralisation (= f_REMIN · J_ORG) accounted for 166 · 10¹² mol/yr. Thus, the total organic load on ocean sediments and corresponding weighted average of the remineralisation fraction are J_RAIN = 192 · 10¹² mol/yr and f_REMIN = 87 % (Burdige, 2007), respectively. The average burial efficiency (BE = 1-f_REMIN = 13 %) mainly reflects shallow settings with high sedimentation rates where the burial efficiency can be up to 30 % in muddy sediments and only 1 % in sandy sediments (Burdige, 2007).

The fraction of organic carbon that is remineralised in sediments varies as a function of the quality of organic matter. With faster sinking rates for marine organic matter, we expect a higher labile fraction to be delivered to the sediments. Therefore, we also expect less remineralisation inside the sediments, f_REMIN <87 %, prior to the Cambrian Stage 2–3 oxygenation episode, when organic particles were smaller and the overall quality of organic matter delivered to the seafloor was lower. We derive the range of possible values for f_REMIN consistent with the carbon and sulphur isotope data.

Sulphur isotopes
The stable isotope composition of sulphur in the oceanic sulphate pool (δ³⁴S, here abbreviated δ_S) is at quasi-steady state and is related to the ratio of global pyrite to total sulphur burial in the oceans (f_PY), the average isotopic difference between seawater and buried pyrite (Δ_PY), and the isotope composition of the inputs (δ_S,IN):

δ_S = δ_S,IN + ƒ_PY * Δ_PY

In the sulphur cycle, the major sinks are pyrite burial (J_PY) and carbonate plus evaporate deposition (J_EVAP) with the latter displaying same S isotope composition as average seawater.

ƒ_PY = J_PY / (J_PY + J_EVAP)

For example, the observed change in seawater δ³⁴S from 28 to 40 ‰ during the Cambrian Stage 2–3 oxygenation episode suggests f_PY increases from 40 % to 70 % (+~75 % relative increase) if we assume δ_S,in and Δ_PY were constant and at modern values (Table S-8).

	Before the event	At the peak	Behaviour	Modern value
δC	0 %	3 %	increase	0 ‰
δS	28 %	40 %	increase	21 ‰
this leads to:
ƒORG	17 ± 4 %	28 ± 5 %	increase	20 ‰
ƒPY	40 ± 17 %	70 ± 22 %	increase	36 ‰

Download in Excel

Relationship between organic carbon loading and seawater δ¹³C and δ³⁴S
Enhanced organic carbon export to the seafloor will promote both marine organic carbon burial and pyrite burial. Below we explore the solution space consistent with the isotope trajectories during the Cambrian Stage 2–3 oxygenation event.

Pyrite burial is a function of the remineralised fraction of the organic carbon supplied to the sediments (e.g., J_RAIN·f_REMIN):

J_PY = α * J_RAIN * ƒ_REMIN

Here, a constant value α = 0.0072 can be derived empirically from the modern-day global ocean state (using parameter values listed in Table S-8). The α value is determined by numerous physical and chemical parameters. We assume that the sulphate reducing microorganisms account for a constant fraction of anaerobic remineralisation of organic carbon and that the stoichiometric factors relating reactive sulphur species produced (∑S^2- and S⁰) per C respired, as well as FeS₂ trapped per reactive S produced, are constant. The ratio between sulphide trapped as pyrite and total sulphide produced (β) will, however, change as a function of the degree of sediment mixing by bioturbation (Canfield and Farquhar, 2009).

Generally, the value of α need not have remained constant over Earth history, since both sediment mixing by bioturbation and oxygen availability in sediments have certainly changed over time. Cambrian oxygen levels were lower and bioturbation was less widespread and less efficient than today (this work; Mangano and Buatois, 2014; Tarhan and Droser, 2014). Therefore, we expect higher reactive Fe availability in more anoxic oceans, less sulphide oxidation in sediments and more pyrite sulphur to be trapped per organic carbon loaded into the sediment. This would increase the value of α, so that α > 0.007. Conversely, lower marine sulphate levels might decrease α if anaerobic processes other than sulphate reduction contributed to a larger part of the remineralisation. Canfield and Farquhar (2009) estimated the change of β in response to the onset of bioturbation. They found that a decrease of β from 0.22–1.00 to 0.033 would lead to a –30 ‰ drop of δ³⁴S by ‘turning on’ bioturbation to modern levels. This would also lead to a 7–30 fold drop in α. Bioturbation was, however, still not as efficient as in modern oceans during the Cambrian Stage 2–3 event. Indeed, the maximum depth of bioturbation was high (up to 1 metre; Mangano and Buatois, 2014), but the degree of mixing was limited and the lateral extent of mixing were likely restricted to only the inner part of the continental shelves (Mangano and Buatois, 2014; Tarhan et al., 2015). Our model is designed to understand the effect of enhanced organic rain rate on the coupled C and S cycles, so we keep α constant and high during the event (3-30 fold higher than today, Table S-8 and Table S-18).

Eqs. S-18 and S-21 can be combined into one equation that relates organic carbon burial and pyrite sulphur burial via the proportionality constant, α, and remineralisation:

J_PY = α * ƒ_REMIN / (1 - ƒ_REMIN) * J_ORG

This equation states that for a given set of the fluxes J_ORG and J_PY (consistent with isotopic constraints), there will be a unique relationship between α and f_REMIN.

In summary, there are five equations (S-16, S-17, S-19, S-20 and S-22) with a total of ten unknown parameters (δ_IN,C, Δ_C, δ_IN,S, Δ_S, J_CARB, J_EVAP, α, f_REMIN, J_ORG, J_PY) that describe the consequences for the global marine C and S cycles when organic carbon export to the seafloor is subject to change.

Model estimates for increases in organic carbon export during the Cambrian Stage 2–3 oxygenation episode
We use the carbon and sulphur isotope compositions of the ocean to calculate how increases in the export organic matter to sediments would drive changes in f_REMIN, J_ORG and J_PY during the event.

In our model simulations we enhance organic carbon export to the seafloor (increasing J_RAIN) while keeping seven parameters constant in each model run (δ_IN,C, Δ_C, δ_IN,S, Δ_S, J_CARB, J_EVAP, α). The oceans are assumed in quasi-steady state with respect to C and S sinks and sources before the event and at the peak of the event (as explained above). As a first step, we utilise Eqs. S-17 and S-20 to get estimates for f_ORG and f_PY prior to the event and at the peak of the event using the isotope compositions of the ocean derived from our data. Therefore, the f-ratios for the marine C and S cycle during the Cambrian Stage 2–3 event changed as shown in Table S-10. The propagated uncertainties are derived from the range of parameter values from Table S-8.

	Before the event	At the peak	Behaviour	Modern value
δ13C	0 ‰	3 ‰	increase	0 ‰
δ34S	28 ‰	40 ‰	increase	21 ‰
ƒORG /(1-ƒORG)	0.21 ± 0.06	0.40 ± 0.10	increase	0.25
ƒPY/(1-ƒPY)	0.81 ± 0.51	6.21 ± 5.29	increase	0.56

Download in Excel

Table S-10 shows that while there is an increase in the reducing sink for carbon (organic C burial) it was still subsidiary to the oxidising sink (carbonate burial), whereas there was a remarkable increase of the reducing sink for S (pyrite sulphur burial flux) which is ~6-fold greater than S burial in evaporites during the event.

For a given set of constraints on the three parameters—J_CARB, J_EVAP, and α—the three variables (f_REMIN, J_ORG, J_PY) can be adjusted to fit the observed isotope stratigraphy. First, we assume J_CARB, J_EVAP, and α were at their modern values (Table S-11). The model implies that the organic export (J_RAIN) to the seafloor exceeded modern values by 2–15 times, and that the fraction of organic carbon remineralised (f_REMIN) was even larger than today (>0.87). With slower organic sinking rates and less labile organic matter exported available for remineralisation in the sediments (Logan et al., 1995), such a scenario appears unrealistic.

	Before the event	At the peak	Behaviour	Modern value
JRAIN	440·1012	3190·1012	increase	192 ·1012
JORG	52·1012	96·1012	increase	26 ·1012
JPY	2.8·1012	22·1012	increase	1.2 ·1012
ƒREMIN	0.88	0.97	increase	0.87

Download in Excel

However, previous models suggest a larger proportion of total sulphide produced by sulphate reduction (a 7–30 fold increase over the modern) formed pyrite in the sediments during the early Cambrian (Canfield and Farquhar, 2009). If we take this into consideration within our model, this leads to a proportional increase in the value of α and more realistic estimates for the remineralisation in global sediments. The consequence of a 10-fold increase in α is shown in Table S-12. In these estimates, we have kept all oxidative burial fluxes (J_CARB, J_EVAP) at modern values.

	Before the event	At the peak	Behaviour	Modern value
JRAIN	90·1012	406·1012	increase	192 ·1012
JORG	50·1012	93·1012	increase	26 ·1012
JPY	2.8·1012	22·1012	increase	1.2 ·1012
ƒREMIN	0.43	0.76	increase	0.87

Download in Excel

Results from this modelling exercise indicate plausible scenarios with high organic rain flux, high organic carbon and pyrite burial fluxes produce acceptable solutions. However the solution space also includes significantly smaller estimates for the organic export fluxes to the sediments. Due to limited vegetative cover coverage on land during the Cambrian, weathering fluxes were likely lower than today; the expansion of land plants is thought to increase Ca weathering ~4 fold in the Mid-Palaeozoic (Moulton and Berner, 1989; Lenton et al., 2012; Quirk et al., 2015). This would imply that carbonate and gypsum burial fluxes (J_CARB, J_EVAP) may also have been smaller than today, due to the reduced Ca fluxes to the oceans. The consequence on our calculations is shown in Table S-13.

	Before the event	At the peak	Behaviour	Modern value
JRAIN	23·1012	100·1012	increase	192 ·1012
JORG	13·1012	24·1012	increase	26 ·1012
JPY	0.70·1012	5.6·1012	increase	1.2 ·1012
ƒREMIN	0.43	0.76	increase	0.87

Download in Excel

Generally, the estimated f_REMIN values are not constant when J_CARB and J_EVAP are changed. This is the case in the example, Table S-13 vs Table S-12, simply because J_CARB/J_EVAP is constant (both J_CARB and J_EVAP change by a factor 1/5), implying J_ORG/J_PY remains the same, and therefore the same f_REMIN is derived in Eq. S-23.

Alternatively, lower peak pyrite S burial fluxes are calculated, if the average sulphur isotopic difference associated with pyrite burial was larger than 35 ‰. The consequences of Δ_S = 50 ‰ are shown in Table S-14 (compare to Table S-11).

	Before the event	At the peak	Behaviour	Modern value
JRAIN	69·1012	145·1012	increase	192 ·1012
JORG	50·1012	93·1012	increase	26 ·1012
JPY	1.4·1012	3.7·1012	increase	1.2 ·1012
ƒREMIN	0.27	0.35	increase	0.87

Download in Excel

Also lower pyrite S burial fluxes are calculated, if the isotope composition of the oceanic S input was greater than 8 ‰ (e.g., Halevy et al., 2012). For example, the consequences of δ_IN,S = 16 ‰ are shown in Table S-15 (compare to Table S-11).

	Before the event	At the peak	Behaviour	Modern value
JRAIN	67·1012	159 ·1012	increase	192 ·1012
JORG	50·1012	93 ·1012	increase	26 ·1012
JPY	1.1·1012	4.6 ·1012	increase	1.2 ·1012
ƒREMIN	0.23	0.40	increase	0.87

Download in Excel

J_ORG and J_PY decline with higher δ_IN,C, but the change is small for changes in δ_IN,C (+1 ‰). The 1 % shift could occur with additional ~20 % increase in the carbonate weathering flux relative to average carbon input flux to the oceans (compare Table S-16 vs. Table S-11).

	Before the event	At the peak	Behaviour	Modern value
JRAIN	71·1012	373·1012	increase	192 ·1012
JORG	31·1012	61·1012	increase	26 ·1012
JPY	2.8·1012	22·1012	increase	1.2 ·1012
ƒREMIN	0.55	0.83	increase	0.87

Download in Excel

J_ORG and J_PY estimates decline with higher Δ_BIO, but the magnitude of change is small for any plausible variation of the average C isotopic difference expressed in marine sediments Δ_BIO (+ 4‰, assumed). Compare Table S-17 vs. Table S-11.

	Before the event	At the peak	Behaviour	Modern value
JRAIN	82·1012	388·1012	increase	192 ·1012
JORG	42 ·1012	76 ·1012	increase	26 ·1012
JPY	2.8 ·1012	22 ·1012	increase	1.2 ·1012
ƒREMIN	0.47	0.80	increase	0.87

Download in Excel

In summary, several scenarios of enhanced organic export to the seafloor would simultaneously lead to the observed magnitudes of the δ¹³C and δ³⁴S excursions. Importantly, the various model examples presented and discussed above show that organic carbon export to the seafloor could have increased by a factor of 2–6 during the event (Table S-18). Favoured models include higher values for α consistent with less sediment mixing by bioturbation suggestive of ~10-fold decrease in sulphide loss via oxidation. In addition, a higher S isotopic difference Δ_S = 50 ‰, a higher δ_IN,S = 16 ‰, and/or smaller oxidative C and S burial fluxes than today are scenarios that also produce acceptable solutions.

Case	Example	Initial JRAIN	Increase of JRAIN	ƒREMIN
	Modern oceans	219·1012		87 %
1	Higher α (10-fold, α = 0.072) (Table S-12)	90·1012	x 4.5	43 to >76 %
1b	Higher α (3-fold, α = 0.022)	181·1012	x 6.3	71 to >91 %
1c	Higher α (30-fold, α = 0.22)	64·1012	x 3.1	20 to >52 %
2	Case 1 + smaller JCARB, JEVAP (1/5) (Table S-13)	23·1012	x 4.5	43 to >76 %
3	Case 1 + higher ΔS (50 ‰) (Table S-14)	71·1012	x 2.1	27 to >35 %
4	Case 1 + higher δIN,S (16 ‰) (Table S-15)	67·1012	x 2.4	23 to >40 %

Download in Excel

S6. Model: Oxygenation dynamics during the Cambrian Stage 2–3 event

In our conceptual model for the oxygen dynamics during the event (Fig. 3), the appearance of large animals with guts expands the oxygenated zone in the upper ocean while bioturbating animals respond slowly to cause deoxygenation of the atmosphere and oceans as a whole. The feedbacks at play are illustrated in Figure S-10.

Onset of event
The emergence of larger animals with guts in the pelagic zone enhances the average sinking rate of organic particulates in the oceans. Once marine organic matter begins to sink faster through the water column, it enhances the organic load onto the seafloor and fuels more anaerobic remineralisation in (primarily unbioturbated) sediments. With less aerobic consumption in the oceans, oceans become more oxygenated. Accordingly, this evolutionary event would trigger ocean oxygenation at a characteristic time scale, τ_EVOL, dictated by the radiation of the relevant pelagic fauna (Fig. S-10, red arrow).

Our model assumes sinking rates of organic matter have forever increased, and that two stabilising feedback loops slowly acted to dampen oceanic oxygen levels via the effect on atmospheric pO₂.

Stabilisation
With more food and O₂ at the benthos over greater areas of the continental shelves, this reorganisation of the marine carbon (and sulphur) cycle would promote bioturbation (e.g., the proportion of organic carbon burial that occurs in bioturbated sediments, f_biot (Boyle et al., 2014). Bioturbation both enhances aerobic remineralisation in sediments (Fig. S-10, green arrow) and limits the availability of the ultimate limiting nutrient for new photosynthetic production: phosphate (Boyle et al., 2014) (Fig. S-10, blue arrow). Both effects lead to stabilising feedbacks that limit organic carbon burial, which is the principal source of atmospheric pO₂ in all Earth system models (Bergman et al., 2004; Berner, 2006).

The slowest process involved in the feedback loops ultimately dictates the rate of the response to ocean oxygenation at this time. Here, a) τ_biot is the characteristic time scale for bioturbating fauna to exploit the new habitats; b) τ_PO4 is the residence time of P in the Cambrian oceans; c) τ_PP and d) τ_SINK are the characteristic time scales for marine primary producers to respond to changing P availability (awaiting the response of the proximate limiting nutrients, e.g., N₂-fixation) and the characteristic travel time for organic matter in the oceans, respectively. The characteristic time scales for remineralisation of organic matter in the sediments and in the water column are denoted τ_{REMIN, SED}, and τ_REMIN,WC, respectively. Lastly, the characteristic time scale associated with changing atmospheric O₂ levels is given by the residence time of atmospheric O₂ with respect to organic carbon burial:

τ_pO2 = M_O2 / J_ORG

Although, the Cambrian rates are generally not known, we can assume that rates of organic production and remineralisation (τ_PP, τ_SINK, τ_{REMIN, SED}, and τ_REMIN,WC) and evolution (τ_biot) are fast compared to the response time of the biogeochemical cycles (τ_PO4 and τ_pO2). Today, τ_PO4 ~15–50 kyr (Sarmiento and Gruber, 2006) and τ_pO2 ~ 1 Myrs (Bergman et al., 2004; Berner, 2006), which means the feedback loop responds on a time scale dictated by the size of the atmospheric O₂ inventory and marine organic carbon burial flux (Eq. S-23).

Therefore, in our model the time scale of the Cambrian Stage 2–3 oxygenation episode (1.3 ± 0.8 Myrs Maloof et al., 2010a), see calculation in Methods, Samples and Geological Settings) constrain the atmospheric O₂ inventory relative to the organic carbon burial flux, because the slow deoxygenation is dictated by the rate at which slower organic burial leads to decreasing atmospheric O₂. If we assume that it took between 1/2 and 1 times the falling limb of the deoxygenation trend, that is between 325 ± 200 and 650 ± 400 kyr, we obtain a ballpark estimate for the Cambrian Stage 2 atmospheric pO₂ to be 26 and 52 % PAL (present atmospheric level). However, the lack of a terrestrial biosphere at this time, accounting for ~33 % of modern day organic burial (Burdige, 2007), means that a corrected estimate for Cambrian Stage 2 atmospheric pO₂ is 17–35 % PAL. This estimate is, however, proportional to the actual global organic carbon burial flux during the falling limb of the C and U excursion.

Full size image | Download in Powerpoint

Extended Data

Extended data 1

Sample	Strat. height (m)	Age* (Ma)	Th (ppb)	U (ppb)	U/Th	δ238U (‰)	± 2SE (‰)	n	U EFTh ƒ	Mg/Ca	Al/Ca	Fe/Ca	Mn/Sr	Sr/Ca
Sukharikha River Section (A), Krasnoporog Fm
A393	624.6	519.81	301	93	0.31	-0.667	0.032	5	0.2	0	0	0	0.45	0.0015
A388	619.6	519.98	249	83	0.33	-0.686	0.021	5	0.3	0	0	0	0.65	0.0013
A384	615.6	520.12	408	39	0.09	-0.6	0.02	1	-0.7	0	0	0	0.82	0.0012
A379	610.6	520.29	155	60	0.39	-0.433	0.043	5	0.5	0.02	0	0.01	0.89	0.0014
A376	607.6	520.39	221	169	0.77				1.9					0.0018
- rpt	607.6	520.39	228	214	0.94	-0.63	0.015	5	2.6				0.89
A372	603.6	520.53	449	54	0.12	-0.517	0.026	3	-0.5					0.0014
- rpt	603.6	520.53	333	44	0.13				-0.5	0.01	0	0.02	0.76
A368	599.6	520.66	317	196	0.62	-0.432	0.027	5	1.4	0.01	0	0.01	0.56	0.0016
A366	597.6	520.73	275	445	1.62	-0.542	0.023	5	5.2	0.02	0	0.01	0.52	0.0014
A364	595.6	520.8	231	90	0.39	-0.625	0.015	5	0.5	0.01	0	0.01	0.52	0.0013
A361	592.6	520.9	352	82	0.23	-0.569	0.015	3	-0.1	0.01	0	0.01	0.46	0.0016
A358	589.6	521	487	142	0.29	-0.643	0.025	5	0.1	0.02	0	0.02	0.71	0.0014
Malaya Kuonamka River Section (K3), Emyaksin Fm
K3-18P	21.5	521.4	268	120	0.45				0.7	0.03	0	0.02	1.52	0.0009
K3-22P	26	521.13	246	239	0.97				2.7	0.06	0	0.04	1.01	0.0014
K3-26P	30	520.89	456	54	0.12				-0.5	0.03	0	0.02	2.12	0.0008
K3-28P	32	520.77	183	31	0.17				-0.4	0.02	0	0.02	1.25	0.0011
K3-30P	34	520.65	310	54	0.17				-0.3	0.05	0	0.03	1.38	0.0010
K3-32P	36	520.53	276	74	0.27				0.0	0.07	0	0.23	1.29	0.0014
K3-35P	39	520.35	418	49	0.12				-0.6	0.02	0	0.02	0.74	0.0012
K3-37P	41	520.23	943	87	0.09				-0.6	0.02	0	0.01	1.18	0.0009
K3-39P	43	520.11	338	34	0.10				-0.6	0.02	0	0.02	1.29	0.0009
K3-41P	45	519.99	348	39	0.11				-0.6	0.03	0	0.02	1.65	0.0010
Bol'shaya Kuonamka (K6), Emyaksin Fm
K6-24	23	521.4	569	151	0.26				0.0	0.05	0	0.02	1.03	0.0012
K6-27	26	521.15	363	67	0.19				-0.3	0.04	0	0.02	1.44	0.0012
K6-30	29	520.89	568	65	0.12				-0.6	0.06	0.01	0.02	1.62	0.0011
K6-31A	30	520.81	615	137	0.22				-0.1	0.02	0	0.02	0.98	0.0015
K6-34	33	520.55	815	111	0.14				-0.5	0.07	0	0.01	1.44	0.0013
K6-35	34	520.47	475	72	0.15				-0.4	0.02	0.01	0.01	0.97	0.0047
K6-38	37	520.21	380	66	0.17				-0.3	0.01	0	0.01	0.76	0.0009
K6-39	38	520.13	312	46	0.15				-0.4	0.01	0	0.02	1.35	0.0008
K6-40	39	520.04	251	178	0.71				1.7	0.01	0	0	1.00	0.0009
Reference materials
IAPSO seawater¬						-0.342	0.022	5
SRM-1d			77	743	9.64	-0.083	0.017	7	36.8	0.01	0.01	0.09
- rpt			70	478	6.87	-0.089	0.024	2	26.2	0	0	0
SRM-1d †			5475	1475	0.27	-0.105	0.017	7	1	0	0	0.01
COQ-1 †			112	727	6.51	-0.337	0.024	5	24.9	0.08	0.03	0.2
BCR-2 †						-0.274	0.017	5
BCR-2 †						-0.255	0.013	4
BCR-2 †						-0.24	0.009	5

*Age model based on δ¹³C chemostratigraphy after Maloof et al. (2010a).
Samples marked with † are whole-rock analysis, all others are mild acetic leaches (see methods for details).
^¬) The processing of IAPSO seawater followed the procedure of Weyer et al. (2008).
^ƒ) Uranium Enrichment factor is calculated as follows: U EF_Th = (U/Th) /(U/Th)_UCC, where the UCC stands for upper continental crust with a U/Th = 0.26

Download in Excel

Extended data 2

Sample	Calcite	Quartz	Dolomite	Ankerite	Chlorite (Clinochlore)	Mica	K-Feldspar	Haematite	sum
	wt. %	wt. %	wt. %	wt. %	wt. %	wt. %	wt. %	wt. %
Sukharikha River Section. Igarka River. NW Siberia
A393	93	2	0		1	1	3		100
A388
A384	88	4	0		1.5	1.75	4.5	0.25	92
A379	56	25.5	8	0		2	8.5		66
A376	85	11	<5	0.5	1	0.5	2.5		97.5
A372
A368	76	7	4	2.5	4	2	4.5		95.5
A366	79	5	6.5	1.5	1	1	6		100
A364	90	5	0		1	0.5	3.5		96
A361	90	2	0.5	0.5	0.5	1	5.5		94
A358	71	4	9.5	9.5	6	3	7.5		103
Bolshaya Kuonamka. Anabar Uplift. Section 96-6
K6-24	88	4	1.5		2	2.5	2		97.5
K6-27	85	4	4		2.5	2	2	0.5	97
K6-30	80	4	8		3	2	2.5	0.5	97
K6-31A	90	4	0.5	1	2	1.5	1.5	< 0.5	97.5
K6-31B	86	4.5	0.5	1.5	2	2.5	2.5	0.5	90
K6-34	76	5	3	6.5	3	2.5	3	1	97.5
K6-35	82	5.5	3.5	2	2.5	2	2.5		89.5
K6-38	91	3	0.5	0.5	1	2	2		100
K6-39	87	4.5	1	1.5	2	2	2	< 0.5	95.5
K6-40	84	2		0.5	1.5	1.5	0.5		86.5

Muscovite and Apatite-OH were looked for, but not detected, in the samples.

Download in Excel

Extended data 3

Sample	Age assigned	TOC	δ34S	δ13C	δ18O
	Ma	wt. %	‰ (V-CDT)	‰ (V-PDB)	‰ (V-PDB)
Sukharikha River Section, Krasnoporog Fm
A350	521.27	0.04	30.48	-1.48	-6.85
A358	521	0.04	32.58	-1.27	-6.9
A364	520.8	0.04	39.53	0.09	-7.24
A366	520.73	0.03	38.27	0.67	-6.6
A368	520.66	0.03	37.41	1.29	-6.58
A372	520.53	0.03	39.37	1.56	-7.12
A376	520.39	0.02		2.63	-6.79
A379	520.29	0.02	37.5	2.03	-6.59
A384	520.12	0.03		1.09	-6.98
A388	519.98	0.03	35.02	0.07	-6.94
A393	519.81	0.02		-0.5	-7.15
Malaya Kuonamka River Section (K3), Emyaksin Fm
K3-18P	521.4	0.033	27.93	-1.29	-6.24
K3-22P	521.13	0.159		-1.1	-5.78
K3-26P	520.89	0.025	27.93	-0.72	-5.99
K3-28P	520.77	0.032	33.27	0.22	-5.96
K3-30P	520.65	0.028	36.16	0.97	-5.57
K3-32P	520.53	0.024	35.2	1.76	-5.62
K3-35P	520.35	0.026	39.03	2.43	-5.8
K3-37P	520.23	0.018		1.72	-5.92
K3-39	520.11	0.027		1.03	-5.88
K3-41	519.99	0.131		0.23	-6.32
Bol'shaya Kuonamka (K6), Emyaksin Fm
K6-24	521.4	0.046		-1.22	-6.01
K6-27	521.15	0.05	28.03	-1.02	-6.26
K6-30	520.89	0.097	28.81	-0.54	-6.07
K6-31	520.81	0.035	32.35	0.21	-6.18
K6-32	520.72	0.062		0.99	-5.91
K6-34	520.55	0.041	36.86	1.8	-5.42
K6-35	520.47	0.1		2.63	-5.49
K6-38	520.21	0.024		1.81	-5.65
K6-39	520.13	0.095	34.93	1.27	-5.84
K6-40	520.04			0.11	-5.95

Download in Excel

Extended data 4

Formation

Locality

Biozone

Lithology

Age (max)

Age (min)

Mo

U

Mo/TOC

U/TOC

δ98Mo

δ98Mo ref ‰ offset from NIST 3136 at 0.25‰

References

Dictyonema Shales

Sweden

Shale

482

√

Quinby-Hunt et al.(1989)

Alum Shale

Albjära, Sweden

Cambrian Series 4

Shale

485

√

√

√

0.08

Dahl et al. (2010)

Alum Shale

Gislövhammar, Sweden

Cambrian Series 4

Shale

485

√

√

√

0.08

Dahl et al. (2010)

Alum Shale

Cambrian Series 4

Shale

499

√

√

Lewan and Buchardt (1989)

Alum Shale

Cambrian Series 4

Shale

499

√

√

Partin et al. (2013)

Alum Shale

Andrarum-3, Sweden

Cambrian Series 3-4

Shale

500

√

√

√

0.08

Dahl et al. (2010)

Alum Shale

Andrarum-3, Sweden

Cambrian Series 3-4

Shale

500

√

√

√

0.08

Gill et al. (2011) (except δ98Mo)

Alum Shale

Närke Area, Sweden

Shale

505

499

√

√

√

Leventhal (1991)

Burgess Shale

Canada

Cambrian Series 3

Shale

505

√

√

√

0.08

Dahl et al. (2010)

Hay River

Georgina Basin, Australia

Shale

505

√

√

Donnelly et al. (1988)

Niutitang

Zunyi, South China

Cambrian Stage 3

Shale

520

√

√

Jiang et al. (2006)

Niutitang

Zhangjiajie

Cambrian Stage 3

Shale

520

520

√

√

Jiang et al. (2006)

Niutitang

Dingtai

Cambrian Stage 3

Shale

521

517

√

√

√

Xu et al. (2012)

Niutitang

Maluhe

Cambrian Stage 3

Shale

521

√

√

√

Xu et al. (2012)

Niutitang

Dazhuliushui

Cambrian Stage 3

Shale

521

√

√

√

Xu et al. (2012)

Niutitang

Sancha

Cambrian Stage 3

Shale

521

√

√

√

Xu et al. (2012)

Niutitang

Ganziping

Cambrian Stage 3

Shale

521

520

√

√

√

√

Lehmann et al. (2007)

Niutitang

Yuanling

Cambrian Stage 3

Shale

521

520

√

√

√

√

Lehmann et al. (2007)

Yuertushi

Xiaoerbulaki, Tarim Basin, NW China

Cambrian Stage 2

Shale

521

√

√

√

Yu et al. (2009)

Yuertushi

Sugaitebulaki, Tarim Basin, NW China

Cambrian Stage 2

Shale

521

√

√

Yu et al. (2009)

Niutitang

Maluhe

Cambrian Stage 3

Shale

521

√

Lehmann et al. (2007)

Yu’anshan

Chengjiang, South China

Cambrian Stage 3

Shale

521

√

√

√

0.08

Dahl et al. (2010)

Niutitang

Ganziping

Cambrian Stage 3

Shale

521

520

√

Wille et al. (2008)

Niutitang

Yuanling

Cambrian Stage 3

Shale

521

520

√

Wille et al. (2008)

Yuanshan

Xiaotan, China

Cambrian Stage 3

Shale

521

518

√

√

Och et al. (2013)

Yuanshan

Dapotuo, China

Cambrian Stage 3

Shale

521

518

√

√

Och et al. (2013)

Niutitang

Zhongnan, China

Cambrian Stage 3

Shale

521

520

√

√

Och et al. (2013)

Niutitang

Dazhuliushui

Cambrian Stage 3

Shale

521

√

√

Och et al. (2013)

Niutitang

Maluhe

Cambrian Stage 3

Shale

521

√

√

Och et al. (2013)

Niutitang

Cili

Cambrian Stage 3

Shale

521

√

√

Och et al. (2013)

Guojiaba

Songtao section, South China

Cambrian Stage 2-3

Shale

524

508

√

√

√

√

Guo et al. (2007)

Jiumenchong

Songtao section, South China

Cambrian Stage 2-3

Shale

524

520

√

√

√

√

Guo et al. (2007)

Shiyantou

Meischucun, China

Cambrian Stage 2

Shale

524

522

√

√

Och et al. (2013)

Shiyantou

Dapotuo, China

Cambrian Stage 2

Shale

524

522

√

√

Och et al. (2013)

Shiyantou

Xiaotan, China

Cambrian Stage 2

Shale

529

521

√

√

Och et al. (2013)

Shiyantou

Meischucun, SYT

Cambrian Stage 2

Shale

529

521

√

√

√

Wen et al. (2011)

Hetang

Cambrian Stage 3

Shale

531

√

Zhou and Jiang (2009)

Ara Group

ALNR-1, Oman

Ediacaran-Cambrian

Shale

540

√

√

Wille et al. (2008)

Ara Group

MM NW-1, Oman

Ediacaran-Cambrian

Shale

540

√

√

Wille et al. (2008)

Ara Group

MM NW-1, Oman

Ediacaran-Cambrian

Shale

542

√

Schröder and Grotzinger (2007)

Dengying

Ediacaran Series 2

Shale

545

√

Partin et al. (2013)

Dengying

Ediacaran Series 2

Shale

545

√

√

Guo et al. (2007)

Liuchapo

Ediacaran Series 2

Shale

545

√

√

Guo et al. (2007)

Dengying

Songtao section, South China

Shale

545

√

Guo et al. (2007)

Nama Group

Namibia

Ediacaran Series 2

Shale

548

√

McLennan et al. (1983)

Doushantuo

Songtao section,South China

Shale

551

√

Partin et al. (2013)

Doushantuo

Songtao section,South China

Shale

551

√

Guo et al. (2007)

Isaac

Castle Creek + Cariboo Mountains, Windermere, Canada

Shale

565

√

0.08

Dahl et al. (2010)

Drook

New Foundland, Canada

Shale

565

√

√

√

0.08

Dahl et al. (2010)

Doushantuo, Cycle 3

Jiulongwan, Zhongling, Minle, Longe Nanhua Basin, South China

Shale

565

555

√

√

Li et al. (2010)

Gaskiers

New Foundland, Canada

Shale

580

√

√

√

0.08

Dahl et al. (2010)

Upper Kaza

Windermere, Canada

Shale

580

√

√

√

0.08

Dahl et al. (2010)

Download in Excel

Supplementary Information References