Limited role for methane in the mid-Proterozoic greenhouse

Potential Constraints for Proterozoic Methane

In the modern ocean, SO₄²⁻, not O₂, is the primary oxidant for CH₄ because of the coupling between SO₄ and CH₄ during anaerobic oxidation of methane (AOM) by microbes. Consequently, freshwater, terrestrial ecosystems (e.g., wetlands) are the most important source of CH₄ to the modern atmosphere (excluding anthropogenic sources) because the vast quantity of CH₄ produced deep within reducing marine sediments does not readily evade oxidation in overlying SO₄²⁻-rich pore waters (20). Quantitative isotopic constraints on seawater SO₄²⁻ concentrations limit Proterozoic [SO₄²⁻] to only a few millimoles per liter (21), or possibly several hundred micromoles per liter (22)—a small fraction of the 28 mM characteristic of the modern ocean. Although less than ∼10% of the modern marine SO₄²⁻ inventory is certainly low in relative terms, a seawater SO₄²⁻ concentration of 1 mM still carries a far greater electron accepting capacity than the entire O₂ reservoir of the generally well-oxygenated modern ocean. Thus, even at comparatively low concentrations, SO₄²⁻ in pore waters and in an anoxic Proterozoic water column could have been a major obstacle preventing biogenic CH₄ from escaping the ocean environment and entering the atmosphere.

Early arguments for elevated pCH₄ during the Proterozoic were predicated on the notion that surface CH₄ fluxes could have been >10× to 20× modern values due to enhanced methanogenesis and complete inhibition of methanotrophy under low [SO₄²⁻] conditions (9, 10), but this assumption has been challenged both experimentally and through the recognition of extensive anaerobic CH₄ recycling in modern analog environments. For example, efficient microbial oxidation of CH₄ coupled to SO₄²⁻ reduction has been documented at SO₄²⁻ concentrations as low as 100 μM (23), which is potentially much lower than the [SO₄²⁻] conditions that typified the Proterozoic ocean (21). Recent work has also shown that CH₄ oxidation can be metabolically coupled to a range of alternate electron acceptors [e.g., Mn(III/IV) and Fe(III)] in freshwater lake systems (24) and in SO₄²⁻-deficient marine sediments (25). These metabolic pathways are theoretically more energetically favorable than SO₄²⁻-based AOM and are likely to have been important for Precambrian CH₄ cycling given the apparent abundance of Fe and Mn in the Precambrian ocean (23, 26); however, there have been few attempts to include these metabolic sinks for CH₄ in quantitative biogeochemical models, with attempts, to date, focusing on photochemical models that do not explicitly resolve redox cycling within the ocean (23, 27).

Somewhat counterintuitively, low pO₂ in the Proterozoic may have also been an obstacle to CH₄ accumulation in the atmosphere due to strongly diminished UV shielding by O₃ at low pO₂—which would act to greatly decrease the photochemical lifetime of atmospheric CH₄. For this reason, atmospheres with low pO₂ can be more oxidizing toward CH₄ than higher pO₂ atmospheres (16, 17). Thus, there is an optimization of the lifetime of atmospheric CH₄ associated with having sufficient O₂ to promote UV shielding by O₃ but at O₂ levels still low enough to minimize oxidative CH₄ destruction. Previous models that have suggested the existence of a CH₄ greenhouse during the Proterozoic have assumed that pO₂ stabilized at, or above, ∼10% present atmospheric levels (PAL) following the GOE (9), but several diverse proxy records collectively indicate that pO₂ may have eventually stabilized at much lower levels than those seen during the GOE and associated Lomagundi event (28). Recently, the absence of Cr isotope fractionation in mid-Proterozoic marine sediments suggests that atmospheric pO₂ may have been as low as 0.1% PAL, an order of magnitude lower than envisioned by some authors (29). Consequently, it is possible that previous attempts to characterize the Proterozoic greenhouse have overestimated the lifetime for atmospheric CH₄.

Ocean-Resolving Methane Cycle Model

In response to recent revelations regarding the high efficiency of anaerobic CH₄ oxidation under low [SO₄²⁻] conditions (23⇓–25) and refined constraints on atmospheric O₂ levels during the mid-Proterozoic (29), we have revisited the role of CH₄ in mid-Proterozoic climate stabilization and Neoproterozoic climate perturbation. Our calculations were performed using the grid-enabled integrated Earth system model (GENIE), an Earth system model of intermediate complexity. GENIE considers a 3D marine biosphere, including nutrient-limited export production, aerobic respiration, sulfate reduction, and methanogenesis. The marine biogeochemical module also includes aerobic methanotrophy and sulfide oxidation (see ref. 30 for a detailed description). The ocean is divided as a 36 × 36 equal-area grid with 16 depth layers, and the ocean system is coupled to a 2D energy and moisture balance model, a sea ice model, and a 2D model for the calculation of atmospheric chemistry and spatially resolved, bidirectional sea−air exchange fluxes (31).

For our Proterozoic simulations, we have expanded GENIE’s representation of the CH₄ cycle compared with previously published versions of the model. In particular, we have (i) parameterized dynamic calculation of CH₄ photooxidation (incorporated after ref. 17); (ii) refined the competition between methanogens and sulfate reducers (improved from ref. 32); and (iii) added AOM coupled to SO₄²⁻ reduction, the rate law for which we derived from radio-labeled CH₄ oxidation rates and chemical profiles from the anoxic water column of the Black Sea (Eq. S1) (33). In combination, these upgrades add strong nonlinearity to both CH₄ fluxes and CH₄ accumulation as a function of oxidant availability, allowing for an oceanographically realistic calculation of CH₄ cycling during oxidant-deficient intervals of Earth history.

We calculated steady-state biogenic CH₄ fluxes and resulting atmospheric pCH₄ for >75 model configurations, differing primarily in the prescribed sizes of their surface oxidant reservoirs (e.g., O₂ and SO₄²⁻). We systematically quantified the influence of pO₂ (from 10⁻⁵ PAL to 10⁻¹ PAL) and oceanic [SO₄²⁻] (from 0 to 2.8 mM). We then performed a wide range of sensitivity analyses (Supporting Information), starting from our baseline oxidant levels (10⁻³ PAL pO₂, 280 μM SO₄²⁻) and examining the effects of varying (i) prescribed CH₄ fluxes from terrestrial environments (0 Tmol CH₄⋅y⁻¹ to 22 Tmol CH₄⋅y⁻¹), (ii) the oceanic PO₄³⁻ inventory (i.e., export production, from 0.25× to 2× modern), (iii) whether or not N limitation is considered, (iv) the depth distribution of organic carbon (C_org) remineralization, (v) the rate constants for both aerobic and anaerobic microbial CH₄ oxidation, and (vi) the half-saturation constant for SO₄²⁻ reduction. We also explored model sensitivity to the parameterization of CH₄ oxidation within the atmosphere.

Importantly, our modeling experiments were designed to conservatively err in favor of overestimating pCH₄ by calculating generous CH₄ production fluxes while minimizing CH₄ destruction. In our baseline model, we assume complete remineralization of exclusively P-limited export production in a closed ocean−atmosphere system; because this base model stipulates a modern PO₄³⁻ inventory while neglecting potential N stress and C removal through organic burial, these calculations tend to overestimate the amount of organic matter remineralized via methanogenesis. This potential for overestimating CH₄ production is further amplified by our use of a relatively high half-saturation constant for SO₄²⁻ reduction (500 μM SO₄²⁻), which effectively reduces the competitive advantage of SO₄²⁻ reducers over methanogens. Furthermore, because we neglect anaerobic oxidation of CH₄ coupled to electron acceptors other than SO₄²⁻ (e.g., oxidized Fe and Mn), it is unlikely that we overestimate CH₄ oxidation by the marine biosphere. Consequently, the sea-to-air CH₄ fluxes and subsequent pCH₄ calculations presented here should be considered conservative upper limits (see Supporting Information for further details).

Oxidant Controls on Methane Cycling

We find that, during the Proterozoic, SO₄²⁻ would have been the primary control on biogenic CH₄ fluxes from the ocean, not unlike today. Methane production by methanogenesis rapidly declines as SO₄²⁻ reduction, the more energetically favorable metabolism, becomes increasingly competitive at higher [SO₄²⁻] and remineralizes a correspondingly greater fraction of organic material exported from the surface ocean. Meanwhile, CH₄ destruction via AOM increases with increasing [SO₄²⁻], until low CH₄ availability ultimately limits AOM kinetics and slows CH₄ oxidation. The combined result of these effects is that net biogenic CH₄ fluxes to the atmosphere plummet as oceanic [SO₄²⁻] increases (Fig. 1A), and the extreme sea-to-air CH₄ fluxes suggested by Pavlov et al. (9) are not achievable for any of our model configurations. Thus, relatively modest oceanic SO₄²⁻levels can, in principle, preclude significant warming by CH₄ on early Earth, regardless of atmospheric pO₂.

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Fig. 1.

Oxidant controls on CH₄ cycling. (A) CH₄ production flux (methanogenesis; blue squares) and CH₄ consumption flux (the sum of aerobic methanotrophy and AOM; red circles) as a function of marine SO₄²⁻. The difference between methanogenesis and methanotrophy is equal to the net biogenic CH₄ flux (i.e., the flux of CH₄ that enters the atmosphere through sea−air exchange). Note that the SO₄²⁻ experiments shown here assume pO₂ = 10⁻³ PAL (29). (B) Net biogenic CH₄ flux to the atmosphere (red squares) and atmospheric pCH₄ (blue circles) as a function of atmospheric pO₂. Net biogenic CH₄ production declines with increasing O₂ due to enhanced methanotrophic oxidation of CH₄, but CH₄ accumulation in the atmosphere increases despite reduced CH₄ supply due to the establishment of an increasingly effective O₃ layer (see figure 7 from ref. 34). Note that pO₂ experiments shown here assume [SO₄²⁻] = 280 μM.

At higher atmospheric O₂ levels, net biogenic CH₄ fluxes also decline as a consequence of higher rates of methanotrophy—both aerobic (i.e., through direct metabolic consumption of CH₄ with O₂) and anaerobic (i.e., through more effective regeneration of SO₄²⁻ via S²⁻ reoxidation). Globally, however, aerobic methanotrophy is quantitatively much less significant than AOM because, when pO₂ is low, oxygenation and aerobic methanotrophy are spatially restricted to the photic zone (Fig. 2A), which represents less than a few percent of the ocean by volume, whereas AOM occurs throughout the ocean interior in closer association with CH₄ production (Fig. 2C). Consequently, CH₄ oxidation linked to SO₄²⁻ reduction in a broadly anoxic Proterozoic ocean greatly exceeds CH₄ oxidation by O₂ globally. The influence of atmospheric pO₂ on aqueous CH₄ oxidation is further muted when atmospheric pO₂ is low because the surface ocean—the site of O₂ production—is widely supersaturated with O₂ with respect to the atmosphere when pO₂ is less than ∼2.5% PAL. Under these conditions, sea−air exchange of O₂ is effectively unidirectional, which partially decouples atmospheric pO₂ from dissolved O₂ and aerobic metabolism in the surface ocean.

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Fig. 2.

(A) Zonally and annually averaged dissolved O₂ profile. The distribution of dissolved O₂ for our standard Proterozoic ocean is strikingly similar to the O₂ landscape of the Archean ocean (32). Although the photic zone is widely supersaturated with respect to O₂, the subsurface ocean is pervasively anoxic under Proterozoic conditions, and aerobic methanotrophy is restricted to the surface ocean. (B) Zonally and annually averaged SO₄²⁻ profile. SO₄²⁻ is strongly heterogeneous in the subsurface ocean. Where organic fluxes are elevated and water masses are older, pronounced SO₄²⁻ minimum zones are established. These environments are directly analogous to modern O₂ minimum zones (Fig. S1). Note that, where [SO₄²⁻] is plotted elsewhere, globally uniform concentrations are not implied; instead, [SO₄²⁻], as in Fig. 1 and Fig. 3, represents the global average concentration of SO₄²⁻. (C) Zonally and annually averaged CH₄ oxidation rates. Contours represent the combined rates of aerobic and anaerobic methanotrophy. Generally, CH₄ oxidation rates mirror CH₄ concentrations (Fig. S1). For reference, the modern rate of CH₄ oxidation in the anoxic water column of the Black Sea is ∼600 nmol kg⁻¹⋅y⁻¹ (33). Like the Black Sea, CH₄ oxidation is dominated by anaerobic metabolism.

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Fig. S1.

Zonally and annually averaged CH₄ profile. Biogenic CH₄ production is greatest in the shallow subsurface where organic fluxes are high and SO₄²⁻ concentrations are suppressed by extensive anaerobic respiration. CH₄ accumulation in oxygenated surface waters is extremely limited, and CH₄ is not effectively mixed into deep ocean. Unlike atmospheric CO₂, the atmospheric CH₄ reservoir is not buffered by the deep ocean (42).

Despite limited influence on net biogenic CH₄ fluxes, atmospheric pO₂ exerts strong control on the accumulation of CH₄ in the atmosphere when [SO₄²⁻] is sufficiently low to allow significant CH₄ escape to the atmosphere. For the range of atmospheric pO₂ that is consistent with the absence of mass-independent fractionation of S isotopes and unfractionated Cr in Proterozoic marine sediments (10⁻⁵ to 10⁻³ PAL) (29, 34), the atmospheric lifetime of CH₄ increases with increasing pO₂ as the result of an increasingly effective O₃ UV shield (Fig. S2). This effect ultimately dominates at steady state over enhanced methanotrophic oxidation of CH₄ when pO₂ is very low, with the result that atmospheric pCH₄ is greater despite reduced production of CH₄ in an increasingly oxygenated Earth system. For example, as pO₂ increases from 0.1 to 1% PAL, net biogenic CH₄ production declines more than 20% whereas atmospheric pCH₄ increases more than 200% as a consequence of muted photochemistry and a correspondingly greater lifetime for atmospheric CH₄ (Fig. 1B).

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Fig. S2.

Atmospheric CH₄ oxidation vs. pO₂. For the range of Proterozoic pO₂, highlighted in red, CH₄ oxidation rates in the atmosphere decline with increasing pO₂ as a consequence of enhanced UV shielding by O₃; therefore, on Proterozoic Earth, higher pO₂ allows higher pCH₄ for the same net biogenic CH₄ flux. Note that the y axis lacks labels because the actual CH₄ oxidation rate also scales as pCH₄^0.7 (after ref. 17).

Combining the opposing effects of [SO₄²⁻] and pO₂, we find that a robust steady-state CH₄ greenhouse is plausible for only a very narrow window of pO₂−[SO₄²⁻] parameter space (Fig. 3A). Atmospheric pCH₄ values greater than ∼10 ppm by volume (ppmv) do not occur for [SO₄²⁻] > 1 mM, whereas pCH₄ values greater than ∼100 ppmv are restricted to pO₂ between 1 and 10% PAL. Thus, substantial accumulation of CH₄ in the atmosphere requires the unique combination of exceptionally low [SO₄²⁻](<<1 mM), which allows sufficiently large CH₄ fluxes from the marine biosphere, and relatively high pO₂ (1 to 10% PAL), which optimizes the atmospheric lifetime of CH₄. Although uncertainties remain regarding the coevolution of pO₂ and marine [SO₄²⁻], this requirement for modest pO₂ and very low SO₄²⁻ is not consistent with the most recent [SO₄²⁻] and pO₂ proxy records and likely would have been difficult to maintain over geologic timescales considering the coupling between atmospheric pO₂, crustal sulfide oxidation, and SO₄²⁻ fluxes. These conclusions regarding oxidant controls on pCH₄ are insensitive to uncertainty in other model parameters (Fig. 3 B–E and Table S1), indicating that our salient results are robust, and these constraints challenge the paradigm of persistently elevated pCH₄ throughout the Precambrian.

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Fig. 3.

(A) The pCH₄ contours for possible Proterozoic oxidant conditions. For reference, 100 ppmv CH₄ to 300 ppmv CH₄ is required to reconcile the absence of glaciation with the faint young Sun (9), but this value is almost certainly an underestimate (37), unless warming by pCO₂ was higher than that calculated by the current generation of 1D atmospheric models. A CH₄ greenhouse is favored by pO₂ ≈ 1 to 10% PAL, whereas pCH₄ > 10 ppmv is precluded for [SO₄²⁻] > 1 mM. (B−E) The pCH₄ sensitivity to select parameters. The red bar in each panel represents our standard model configuration, and the gray bars represent pCH₄ results for simulations that were run under identical oxidant conditions and differ with respect to (B) the flux of CH₄ from terrestrial environments, (C) the half-saturation constant for SO₄²⁻ reduction, (D) the oceanic PO₄³⁻ inventory, and (E) the depth distribution of organic matter remineralization. Model sensitivity to each of these parameters is small compared with the influence of surface oxidant conditions (A). Unless otherwise specified, the values on the x axis of each panel are presented as times standard value (red bar). See Table S1 for full sensitivity results.

Consequences for Proterozoic Climate

Elevated pCH₄ on the order of 100 ppmv could provide ∼6 K of greenhouse warming, which would substantially, although not completely, compensate for reduced solar luminosity if Proterozoic Earth was persistently warmer than today (10); thus we consider ∼6 K to be the minimum radiative deficit implied by our model if the Proterozoic [SO₄²⁻] approached 1 mM or if pO₂ was lower than previously appreciated. This deficit is a substantial one that is difficult to reconcile with the absence of glaciation and the suggestion that Proterozoic pCO₂ was <10× modern (35).

Other reduced greenhouse gases that are not explicitly considered by our model may have bolstered the mid-Proterozoic greenhouse, helping to alleviate the radiative deficit produced by low atmospheric pCH₄. Most notably, atmospheric levels of nitrous oxide (N₂O) during the Proterozoic may have been much higher if widespread euxinia limited Cu availability and inhibited effective conversion of N₂O to N₂ during the final stage of denitrification (36), potentially yielding ∼5 K of supplemental warming (10). Although SO₄²⁻ would not severely throttle N₂O fluxes, N₂O—like CH₄—is rapidly photolyzed at low pO₂. Nitrous oxide, therefore, is only a viable greenhouse contributor for a partial range of the proposed pO₂ conditions for the Proterozoic (10). Higher hydrocarbons, particularly ethane (C₂H₆), can be photochemically produced from CH₄ and may have been important for regulating the climate of Archean Earth (37), but our model results generally preclude the elevated pCH₄:pCO₂ conditions required for substantial hydrocarbon polymerization. Thus, greenhouse contributions from other hydrocarbons are unlikely during Proterozoic time.

In combination with evidence for relatively low pCO₂ during much of the Proterozoic (35), we conclude that the faint young Sun paradox may still be a problem for the Proterozoic if relatively high CH₄ levels are precluded by current estimates of oceanic SO₄²⁻ and/or atmospheric pO₂. Continued exploration of the metabolic limitations of CH₄ oxidizing consortia (11), revised pCO₂ calculations, or more realistic consideration of heat transport in increasingly sophisticated 3D coupled climate models may eventually reconcile our suggestion of low pCH₄ with clement conditions during the mid-Proterozoic, but, given current constraints on the warming potential of CO₂ during Proterozoic time, the paradox remains unresolved.

Given the possibility that baseline Proterozoic pCH₄ was much lower than previously envisioned, the role of CH₄ in triggering Neoproterozoic climate collapse also requires revisiting. Several authors have previously suggested that the oxidative collapse of a CH₄ greenhouse was involved in the initiation of Neoproterozoic snowball glaciations (1, 8, 9). Indeed, emerging trace metal proxies suggest atmospheric oxygenation in advance of the glaciation (29, 38), despite unclear evidence for oxygenation within the marine realm (39⇓–41). The climatic consequences of the oxygenation of a low pCH₄ atmosphere, however, have not been previously considered.

We suggest that a Neoproterozoic oxygenation event might have triggered a CH₄-based climate collapse, despite low steady-state pCH₄, as a result of the vastly differing timescales over which photochemical, biological, and weathering processes modulate Earth’s greenhouse. For example, the immediate consequence of an atmospheric oxygenation event would be an increase in the photochemical lifetime of CH₄ via UV shielding by O₃, and thus an initial increase in atmospheric pCH₄. Subsequent destabilization of CH₄ hydrates as the result of a warming climate may amplify this initial accumulation of CH₄ in the atmosphere (12, 42). The resulting greenhouse, however, would be inherently unstable because CH₄-induced warming would promote the drawdown of CO₂ on weathering timescales (12, 42); thus, the climate system would then be vulnerable to any disruption of the CH₄ source to the atmosphere (42), including the gradual accumulation of oceanic SO₄²⁻ that would accompany oxygenation (9). A CH₄ greenhouse would also be sensitive to CH₄ instability arising as the result of either further oxygenation or even deoxygenation.

Although conceptually similar to the “methane shotgun” scenario advanced by Schrag et al. (42), our snowball initiation scenario has the advantage of eliminating the need for a large, sustained external source of CH₄ to the ocean−atmosphere system and does not rely on either enhanced biological CH₄ production or high steady-state pCH₄ before climate destabilization. Glacial initiation through the buildup and collapse of a CH₄ greenhouse in response to an oxygenation event may also provide an explanation for the enigmatic C isotope excursions of the late Neoproterozoic (12, 42). Future geochemical and model analyses that clarify the relationship between oxygenation, glaciation, and isotopic perturbation in the Neoproterozoic will be required to validate our proposed snowball scenario and could provide implicit support for our conclusion that baseline pCH₄ was very low during much of the Proterozoic.

Summary

We have shown that long-term stabilization of Earth’s climate system by CH₄ is challenging for much of Earth history, despite the likelihood of greatly enhanced CH₄ cycling within the broadly anoxic Proterozoic ocean. Even at very low SO₄²⁻ concentrations, anaerobic oxidation of CH₄ in the ocean is sufficient to severely throttle CH₄ fluxes to the atmosphere, precluding the stability of an atmosphere with >100 ppmv CH₄ as previously suggested for Proterozoic Earth. If recent pO₂ constraints from Cr isotopes are correct, reduced UV shielding by O₃ may exacerbate this issue, further limiting the potential accumulation of CH₄ in the atmosphere—with the implication that mid-Proterozoic pCH₄ may not have been markedly different from the low levels present today. In this scenario, our results suggest that elevated CH₄ should not be invoked to reconcile existing constraints on the composition of the mid-Proterozoic greenhouse with the absence of glaciation in the mid-Proterozoic. Thus, the faint young Sun paradox remains unresolved.

Although our model results imply relatively low steady-state pCH₄ (<10 PAL) during the second half of Earth history, they do not preclude a significant role for O₂ and CH₄ in destabilizing the Neoproterozoic climate system. Rather than occurring as the result of the demise of a steady-state CH₄ greenhouse, we suggest that the extreme low-latitude glaciations during the Neoproterozoic are the consequence of non–steady-state CH₄–oxidant dynamics produced by the interplay of the marine biosphere, photochemistry, weathering, and climate.