The onset of widespread marine red beds and the evolution of ferruginous oceans

Banded iron formations were a prevalent feature of marine sedimentation ~3.8–1.8 billion years ago and they provide key evidence for ferruginous oceans. The disappearance of banded iron formations at ~1.8 billion years ago was traditionally taken as evidence for the demise of ferruginous oceans, but recent geochemical studies show that ferruginous conditions persisted throughout the later Precambrian, and were even a feature of Phanerozoic ocean anoxic events. Here, to reconcile these observations, we track the evolution of oceanic Fe-concentrations by considering the temporal record of banded iron formations and marine red beds. We find that marine red beds are a prominent feature of the sedimentary record since the middle Ediacaran (~580 million years ago). Geochemical analyses and thermodynamic modelling reveal that marine red beds formed when deep-ocean Fe-concentrations were > 4 nM. By contrast, banded iron formations formed when Fe-concentrations were much higher (> 50 μM). Thus, the first widespread development of marine red beds constrains the timing of deep-ocean oxygenation.

distributed around uplifted "old lands" and consist predominately of red sandstone, siltstone and shales without carbonates. The iron source of these Silurian red beds are thought to be of detrital origin 146 . Therefore, we think that these red beds may have been formed differently from the five intervals we have described. However, if future studies indicate that they were formed by similar processes like the ones in Cretaceous and Triassic, the Telychian red beds could be another representative Phanerozoic MRB.
Most of the Phanerozoic MRBs slightly postdate oceanic anoxic events (OAEs), but in a few cases red beds are also found within the interval of oceanic anoxia. One of the examples is the thin red beds within OAE2 in New Zealand 147 . Further study may reveal if this type of red beds records episodic oxidation within a broad anoxic event or a local phenomenon.
Red-pink carbonates of MRBs all have δ 13 C values that are lower than temporally adjacent strata, creating "negative" δ 13 C excursions (Figs. 2c and 3; Supplementary Figs. [3][4][5][6][7]. This is conceivable because oxidation of reduced iron from anoxic waters would inevitably involve oxidation of organic carbon and incorporation of 13 C-depleted HCO3during carbonate precipitation, adding 13 C-depleted carbon to carbonate. This process may have resulted in negative δ 13 C shifts in the range of -0.5‰ to -2‰, as seen in the Phanerozoic MRBs (Figs. 2c and 3; Supplementary Figs. [3][4][5]. The negative δ 13 C excursion associated with the middle Ediacaran MRB, or the Shuram δ 13 C excursion, however, has a magnitude of ≥ 12‰. While 13 C-depleted carbon from oxidation of organic carbon and 13 C-depleted HCO3certainly made contributions to the Shuram excursion, the amount of oxidants 148 and reduced carbon source 149 required for the Shuram excursion is enormous and has been highly debated.

Supplementary Note 3-Debates on the origin of the Shuram δ 13 C excursion
The negative δ 13 C excursion associated with the middle Ediacaran MRB, or the Shuram excursion, has a magnitude of ≥ 12‰ (ranging from ≥ 4‰ to ≤ -8‰) and a duration of ≥ 5 million years (Myr) 150,151,152,153,154 . The large magnitude and long duration of this δ 13 C excursion make it difficult to interpret using the Phanerozoic carbon cycle models. Early interpretations invoked the upwelling of 13 C-depleted deep water 155,156 , but the enormous amount of 13 C-depleted carbon required for accommodating a > 5 Myr δ 13 C excursion with a magnitude of ≥ 12‰ is difficult to reconcile. This led to the proposal of a large oceanic dissolved organic carbon (DOC) pool (100-1000 times that of the modern ocean DOC) and perhaps a relatively smaller (than modern) dissolved inorganic carbon pool that was more susceptible to carbon isotope changes 149 . Evidence supporting a large DOC pool came from the decoupled carbonate and organic carbon isotopes prior to and across the Shuram excursion 152,157 . This hypothesis, however, is challenged by the equally large amount of oxidants required for remineralizing the large DOC pool 148 . Even with the oxidant budget available in the modern surface environments (including atmosphere and ocean) and with an unlimited organic carbon source, it is difficult to support a 12‰ negative δ 13 C excursion for more than 3 Myr 148 . In addition, more recent paired carbonate-organic carbon isotope analyses documented decoupled-coupled δ 13 Ccarb-δ 13 Corg patterns from multiple intervals of Ediacaran-Cambrian strata 158,159,160 , suggesting that even if a large DOC existed in the Precambrian ocean, it was not large enough to buffer the organic carbon isotopes and the evolution of the DOC reservoir was not unidirectional 159,161 .
The shortage of 13 C-depleted carbon source or oxidants required for the Shuram excursion led to alternative meteoric 162,163 and burial 164  A more recent hypothesis invokes authigenic carbonate precipitation in porewater as a possible origin of the Shuram excursion 165 . Due to anoxic bottom waters, authigenic carbonate precipitation in porewaters in Precambrian oceans may have been much more pervasive than in the modern ocean and might be a major 13 C-depleted carbon flux. This has two implications: (1) the cutoff or decline in the global flux of authigenic carbonate would result in a negative δ 13 C excursion and (2) the addition of authigenic carbonate into primary marine carbonate would result in localized/regional δ 13 C shift. This hypothesis explains some of the spatial variations of the Shuram excursion such as the large isotope gradients 141,161 and local isotope extremes 166 documented from the Doushantuo Formation in South China, but it cannot explain a global δ 13 C excursion with minimum values down to ≤ -12‰ because even a complete cutoff of the authigenic carbonate flux would not result in ocean seawater δ 13 C values lower than the riverine (or average crust) δ 13 C value of ca. -5‰, unless additional evidence confirm that the Shuram excursion is not globally synchronous.
The debate on the origin of the Shuram excursion (and its correlatives) will continue until better constraints on its magnitude, duration, and spatial variations can be achieved, and our findings by no means solve this debate. However, the coincidence of the Phanerozoic-like, middle Ediacaran MRB and the Shuram excursion does confirm that (1) similar to the negative δ 13 C shift associated with the Phanerozoic MRBs, oxidation of organic carbon and incorporation of 13 C-depleted HCO3from anoxic waters during carbonate precipitation likely contributed to the Shuram excursion, (2) the larger magnitude of the Shuram excursion may be related to the longer period of anoxia prior to the middle Ediacaran MRB, during which more 13 C-depleted carbon may have accumulated through remineralization of organic matter, and (3) iron reduction (using iron oxides as electron acceptors) may have contributed, at least locally, to the heterogeneity of the Shuram excursion.