Reverse weathering and the silicon–carbon cycle

As a context for exploring the evolution of reverse weathering, it is useful to consider past key contributions to the field and several aspects of the modern silicon–carbon cycle. To our knowledge, the idea that clay minerals could act as a buffer of seawater pH stems from a seminal paper by Sillén in 19616. Here and in subsequent articles, Sillén challenged the traditional framework of a carbonate system operating as the sole buffer of seawater pH6,34. Several supporting articles followed within the decade, highlighting the importance of the reverse-weathering process for the budgets of not only H+ but also major cations4,5,7,35,36. Perhaps most notably, Mackenzie and Garrels in 19664,7 constructed a mass balance for river water inputs, such that the major constituents were precipitated from seawater as mineral phases commonly found in marine sediments—balancing the budgets required for Na, Mg, K and Si removal through clay authigenesis. Since then, both field and laboratory work have provided mounting evidence for the operation of this process in nature today17,18,22,25,26,27,28,37,38,39,40,41,42, prompting greater acceptance within the broader community. However, when placed in the context of the global carbon27 and silica budgets29,43,44,45, it is generally agreed that reverse-weathering fluxes in modern systems account for a non-negligible, yet small, fraction of either budget. Consequently, this has led to the general exclusion of reverse weathering from the current generation of global carbon cycle models designed to reconstruct past climate states and the composition of seawater (for example, refs 20,46,47). However, because reverse weathering can lead to the long-term sequestration of various elements (for example, Fe, Mg, K, Li, Si and O), secular variations in rates of early diagenetic clay synthesis throughout Earth’s history can alter global mass and isotope balance budgets.

As there have been multiple reviews of the modern global silica budget29,43,44, we provide only a brief overview here (Table 2). The total marine influx of dissolved silica every year is about 10.9 Tmol. Rivers contribute about 7.3 Tmol yr−1 of this source flux, and the remainder are attributed to hydrothermal, aeolian, groundwater and marine weathering sources. In terms of outputs, biogenic silica is responsible for >90% of the flux and the rest is sequestered through reverse-weathering reactions (<10%)29. Strikingly, the magnitude of the recycled flux within modern oceans is far greater than that of net marine input and outputs. The residence time of silica in the oceans is estimated to be 15,000 yr to 20,000 yr, with dissolved silica cycling through the biological system approximately 40 times before sequestration29,43. The uptake of dissolved silica by organisms in surface waters, estimated at 240 Tmol yr−1, is around 30 times larger than the river influx. These rates of biological fixation are so rapid that dissolved silica levels (typically <0.1 mM) are forced into extreme disequilibrium with respect to its amorphous phase (opal-A; equilibrium at about 2.21 mM). This essentially sets up an autocannibalistic system, such that about 56% and about 10% of biogenic silica is remineralized in surface and deep waters respectively and reincorporated into siliceous tests or spicules. Of the remaining 34% or so transported to the sediment–water interface (in highly localized zones), only about 2.5% escapes dissolution and the rest is diffused back into seawater25,29,48. Critically, silicate mineral undersaturation is not limited to biogenic silica—dissolution extends to the majority of common silicate minerals at the seafloor (fed by riverine and aeolian systems) and primary phases within basaltic ocean crust—a process referred to as marine weathering27,31,49. Exceptions include certain select species within the montmorillonite, nontronite and glauconite group of clay minerals that remain saturated at the seafloor despite low dissolved silica levels (Extended Data Fig. 2). In other words, modern seawater conditions are less suitable for reverse-weathering reactions and instead favour marine weathering reactions.

Modern porewater systems, on the other hand, can facilitate reverse weathering. Silica derived from dissolving mineral phases is sometimes re-precipitated as authigenic clays25,27,28, in ‘pockets’ of marine porewater systems with favourable conditions (for example, elevated pH, dissolved silica or clay-forming cations). However, such conditions are rarely observed in modern systems. As illustrated in a compilation of porewater dissolved silica concentrations (n = 6,245) from 453 sediment cores globally, the occurrence of elevated porewater-dissolved silica levels (of >0.5 mM) are globally uncommon in modern systems even where biogenic silica content is high. Estimated Precambrian-like conditions are rarely observed (Extended Data Fig. 3)48. Several key factors contribute to this: (1) about 70% of biogenic silica is remineralized in the water column and does not make it to the sediment–water interface29; (2) the localizing nature of biogenic silica production—meaning that substantial export occurs in specific bands or regions (for example, the Southern Ocean), leaving the rest of the ocean floor largely devoid of biogenic silica50,51; (3) the kinetics of diffusional exchange (between porewater and depleted seawater) is much more rapid than the dissolution of solid particulate phases (horizontal inputs), rendering it difficult to sustain elevated porewater concentrations48. In sum, biotic regulation of the global silicon cycle essentially ‘locks up’ silica in solid phases8,9, forcing a decoupling of the available silica from clay-forming cations that renders reverse-weathering reactions both spatially and kinetically limited.

In the absence of silica-secreting biota, geologic evidence suggest that dissolved silica was sequestered, via saturation and precipitation of inorganic phases, under pervasively enriched Precambrian ocean and porewater systems8,9,52,53,54,55. The stratigraphic, mineralogical and textural features of Precambrian and Phanerozoic siliceous sediments are distinct8,9,55,56,57. Early diagenetic peritidal and subtidal chert deposits that form authigenically through direct precipitation within pore-fluids, are common features of Precambrian successions, and disappear from the record roughly coincident with the Paleozoic ecological rise of radiolarians8,9. These shallow-water authigenic sediments often facilitate the exceptional preservation of non-skeletal Precambrian microfossils8,58,59. The inference is that Precambrian porewater conditions fostered rapid silica ‘entombment’, allowing organic-walled microfossils to escape degradation8,60. Today, such extreme rates of microbial silicification are only observed in silica-rich hot springs61 or reproduced experimentally under elevated silica conditions62. Perhaps more notably, authigenic silicate minerals that form only under silica-enriched conditions are also ubiquitous during this time, as predicted by thermodynamic and experimental guidelines (Fig. 3; Extended Data Fig. 5). For instance the Fe-clay minerals greenalite and minnesotaite are abundant, occurring not only in distinct units but also as impurities within ‘cloudy’ cherts as clay nanoparticles63,64,65,66,67,68,69. The inclusion of authigenic clays within early precipitating diagenetic cherts highlights extremely early formation within porewaters63,64,65,66,67,68,69. These authigenic clays are often regarded as precursor phases of other Fe- and Si-bearing minerals such as hematite, magnetite, siderite, dolomite-ankerite and chert that form at the expense of these early diagenetic clays during later diagenesis, indicating that these authigenic clays were originally even more widespread than current mineralogical records suggest63,64,65,66,67,68,69. Early diagenetic Mg-bearing minerals have also been noted in Proterozoic successions70,71. During this time, dissolved silica would have been maintained between 1.0 mM to 2.21 mM through sorption equilibrium with detrital particles (for example, clays and zeolites), organic matter9, authigenic Fe oxyhydroxides23,72 or saturation with Si-bearing phases. With sorption controlling marine Si levels, dissolved Si concentrations are likely to have been buffered, with any transient shift in dissolved Si inputs or outputs from the marine realm compensated for by exchange (sorption–desorption) from the sorbed pool9,23,72,73. Saturation with amorphous silica9,74 provides an upper, but often invoked, bound on dissolved silica at 2.21 mM8,9,75,76.

The evolution of continental exposure and composition through Earth’s history are likely to have altered the global silicon cycle. The extent of continental exposure can influence weathering intensities, and potentially the balance between marine and terrestrial silicate weathering (see, for example, ref. 77). With lower continental exposure, as has been proposed for the Archean78, two end-member outcomes are possible: either (1) the ratio of terrestrial to marine silicate weathering would have been lowered or (2) terrestrial weathering intensities would have been elevated for a given silicate weathering flux to match an outgassing flux. Nonetheless, recent modelling efforts indicate substantial continental exposure since the Mesoarchean77, which is consistent with recent empirical evidence suggesting extensive terrestrial weathering in the Mesoarchean79. In other words, this work suggests that continental exposure was modern-like for the majority of Earth’s history. The transition from a more mafic to a more felsic upper continental crust would have also changed the ratio of Si released per mole of CO 2 consumed. For instance, at sustained outgassing rate, a more mafic upper continental crust would release less Si during weathering. Under these conditions, the cation to Si ratio released from weathering would be much higher. Traditionally, the transition to a more felsic upper continental crust was proposed to have occurred near the Archean–Proterozoic boundary80,81. However, more recent Ti isotope work suggests a Mesoarchean transition82. Additionally, any increase in Si fluxes linked to changes in continental composition could have been directly offset by decreasing outgassing rates through time32,83,84. Therefore, it is very likely that both of these factors were limited to early to mid-Archean times. As such, we do not explicitly model the effects of continental area and continental composition on global Si input flux. However, we note that it is the f rw ratio (the fraction of alkalinity and the fraction of silica that is sequestered through the reverse-weathering process, as opposed to carbonate and silica precipitation) that governs the impact of reverse weathering on atmospheric \({p}_{{{\rm{CO}}}_{{\rm{2}}}}\) levels and climate (not total fluxes). Further, it is possible that altering the total input flux of dissolved Si may not dramatically affect the total flux of reverse weathering, given that dissolved Si concentrations are set at constant levels (for example, about 2.2 mM) through (non-reverse weathering) abiotic equilibrium (or pseudo-equilibrium) Si reactions within the oceans9.

In a transformative paper, Wallmann and others27 highlighted that the weathering of primary silicate minerals in marine sediments is likely to be a very large CO 2 sink that is not traditionally included in global carbon budgets. In addition, this study also concluded that while reverse-weathering processes can dominate within the upper portion (typically 1–2 m) of the sediment pile, marine weathering processes become more prominent at depth27. Specifically, it was revealed from two sediment cores from the northern slope of Sakhalin Island, Sea of Okhotsk, (1) that limited consumption of cations (Mg2+, Na+, K+ and Li+) through reverse weathering in the upper portion of the sediment pile and (2) that the abundances of primary silicate minerals such as plagioclase feldspar, olivine, pyroxene and volcanic ash decrease down-core into the methanogenic zone and are transformed into smectite. Here, the elevated rates of marine weathering were attributed to the elevated concentrations of dissolved organic humic and fulvic acid anions that complex cations, initiating more rapid dissolution through further undersaturation of primary silicate phases, generating alkalinity in the process. Critically, such elevated total alkalinity levels have also been observed within the methanogenic zones of numerous other sites globally27.

This work also opens the possibility that the by-products of reverse weathering could be subjected to marine weathering deeper in the sediment pile. It is therefore critical to consider the effects of marine weathering in Earth’s geologic past. A general increase in porewater and seawater dissolved silica levels back in time would reduce the degree of undersaturation of primary silicate minerals relative to the modern value (Extended Data Fig. 2). While the degree of saturation of these primary minerals would have still been reduced passing through the methanogenic zone, the starting baseline saturation state in the upper portion of the sediment pile would have been elevated relative to the modern saturation state. Critically, it is unlikely that reverse-weathering products will reach equilibrium in the sediment pile as reverse weathering proceeds, given the strong kinetic limitation on clay formation24,70. Stated in other words, porewaters are likely to remain supersaturated with respect to the clay phases forming in the upper portion of the sediment pile. Thus the journey of authigenic clay minerals would have begun as being supersaturated both in seawater and the upper portion of the sedimentary pile and although there would be a decrease in saturation state within the methanogenic zone, this drop is not likely to drive undersaturation. In contrast, modern systems typically begin with phases in the upper portion of the sediment pile being strongly undersaturated or close to equilibrium (such as albite, anorthite, olivine, pyroxene), and achieve an even greater degree of undersaturation when passing though the methanogenic zone. Whether there is undersaturation for clay phases will depend on the degree of kinetic inhibition for the mineral species and the degree of complexation of the clay-forming elements. In sum, regardless of complexities in this process, under elevated Precambrian-like silica conditions, a much larger degree of cation complexation would be required to facilitate the same degree of marine weathering as in the modern era under silica-depleted conditions. The preservation of these reverse-weathering products in the Precambrian sedimentary record (Fig. 3, Extended Data Fig. 5), is consistent with super-saturation through the sediment pile.

To explore further the importance of marine weathering, we incorporate this process into LOSCAR-RW, parameterized as a production flux of alkalinity at the sediment–water interface (that is, CO 2 consumption). We explore the full proposed range of global CO 2 consumption through marine weathering of around 2–20 Tmol per year27,31. Specifically, we conducted analysis for both modern and Precambrian-like (elevated reverse-weathering buffer) systems (Extended Data Fig. 4). In both systems, the overall trend with increasing marine weathering is towards lowered CO 2 and increased marine pH. A more potent Precambrian reverse-weathering buffer acts to dampen this effect, resulting in a smaller net stable state shift given the same change in marine weathering (that is, a more stable system). Similarly, the Precambrian system also achieves a lower pH stable state and a higher CO 2 at the same marine weathering rate compared to the modern system. In other words, it is possible for Earth to achieve a high-CO 2 state with elevated marine weathering and reverse-weathering rates.

Authigenic clays through time

Clay authigenesis has long been recognized as a critical process that governs the mineralogical and textural make-up of marine sediments21. A wide range of different clay species has been reported to form authigenically in the marine environment (such as palygorskite, montmorillonite, glauconite, saponite, berthierine, greenalite and minnesotaite)18,19,21,68,69,85,86,87. Of these species, many (for example, palygorskite, saponite and montmorillonite) also form authigenically in lacustrine and soil environments, thereby rendering it a challenge to distinguish marine and terrestrial sources85,88. As such, for a representative marine framework, we turn to the minerals greenalite and minnesotaite, which are both common in the record and are generally regarded to form exclusively in marine environments13,63,64,65,66,67,68,69,86,87. Stilpnomelane is another phyllosilicate mineral that has recently been shown to form under early-diagenetic conditions23, despite being traditionally viewed to be of metamorphic origin. Stilpnomelane has also been observed to replace early-diagenetic minerals such as greenalite89,90.

We present an extensive compilation of the reported occurrences of greenalite, minnesotaite and stilpnomelane through time (Extended Data Fig. 5, Supplementary Table 2). Our compilation is based on an extensive literature review, with a survey of over 3,500 individual papers and reports mentioning greenalite, minnesotaite or stilpnomelane, for unique occurrences of these minerals. Members, formations, and groups were all considered unique occurrences. Given the nature of our findings, and that the stratigraphy of Phanerozoic units is typically much more finely divided than the Precambrian, this scheme for what merits consideration of an individual occurrence is conservative. As highlighted by ref. 91, the geologic record becomes less voluminous the further back in time it reaches. Accounting for increasing rock quantity through time is thus necessary for any meaningful result. As such, we normalize our data to the proportion of marine sediment coverage (for a given geologic age) after ref. 92, per time bin (500 Myr; see also the Macrostrat database at https://macrostrat.org). Greenalite and minnesotaite data only are shown in Fig. 3, given the possible metamorphic origins of stilpnomelane (Extended Data Fig. 5). Full raw and normalized data are presented in Extended Data Fig. 5. We find the results of our compilation to be consistent with several previous studies that have noted that authigenic clay minerals are more common in Precambrian than Phanerozoic aged sediments24,57,63,68,69,76,87,93,94.

The composition of the globally integrated authigenic clay assemblage plays a part in determining the absolute flux of acidity generated through reverse weathering. To maintain charge balance during the formation of authigenic clays, the total charge (of dissolved cations) must be balanced by the consumption of an equal amount of alkalinity (herein represented as HCO 3 −) in the reaction. For our modelling purposes, this can be expressed as the ratio of the total amount of alkalinity to the Si consumed during reverse weathering (Alk:Si). We find that this ratio varies between 1 and 4, for common authigenic clay minerals found within Earth’s sedimentary archive21 (Table 1). The common Precambrian clay greenalite, for instance, has an Alk:Si of 3.0. Our model centres on linking the global carbon and silicon cycles because silica forms the basic framework of all clay minerals, and helps in controlling rates of reverse weathering22,25,26,28,37. Although aluminium (Al) can limit the formation of some clay minerals17,39, Al is not required for the formation of non-aluminous clays. Thus, any attempt to track reverse weathering through the global Al (or any other cation such as Mg, Fe and K) mass budget will yield an incomplete estimate of both the extent of reverse weathering and its effect on the global carbon cycle and marine pH.

Modifications to LOSCAR

For this work, we modified the Long-term Ocean-atmosphere-Sediment CArbon Cycle Reservoir Model version 2.0.4 (LOSCAR), which was designed to compute carbon partitioning between the ocean–atmosphere and sedimentary reservoirs on timescales ranging from centuries to millions of years20. The preindustrial ocean geometry was adopted for all model runs. Modifications to LOSCAR for the modelling purposes of this contribution are as follows. In LOSCAR, the terrestrial weathering of silicate rock is simply related to atmospheric \({p}_{{{\rm{CO}}}_{{\rm{2}}}}\) levels via the expression:

$${F}_{{\rm{s}}}={F}_{{\rm{s}}}^{{\rm{o}}}{\left(\frac{{p}_{{{\rm{CO}}}_{2}}}{{p}_{{}_{{{\rm{CO}}}_{2}}}^{{\rm{o}}}}\right)}^{{n}_{{\rm{Si}}}}$$ (6)

where F s refers to the weathering flux of silicate rock from continents, the superscript ‘o’ denotes the initial steady-state value, and n Si controls the strength of the silicate weathering feedback. We calibrate the model according to the most recent global silica budget compilation after ref. 29 (Table 2). In addition to sources of silica that are sensitive to atmospheric \({p}_{{{\rm{CO}}}_{{\rm{2}}}}\) levels (riverine, groundwater and aeolian sources). We also include marine silica sources (seafloor weathering and hydrothermal; F n ). The total marine silica input (F in ) is therefore:

$${F}_{{\rm{in}}}={F}_{{\rm{s}}}^{{\rm{o}}}{\left(\frac{{p}_{{{\rm{CO}}}_{2}}}{{p}_{{}_{{{\rm{CO}}}_{2}}}^{{\rm{o}}}}\right)}^{{n}_{{\rm{Si}}}}+{F}_{n}$$ (7)

At steady state, this input value equals outputs through reverse weathering and opal (F out = F rw + F opal ).

Carbonate precipitation was switched from a biotic (Phanerozoic) to an abiotically controlled Precambrian ocean system, to account for the onset of carbonate biomineralization at the Cambrian–Precambrian boundary95 and the Mesozoic rise of planktic calcifiers96,97,98. In our Precambrian mode carbonate precipitation occurs only at high degrees of supersaturation, consistent with geologic evidence for mineralogically ‘anomalous’ carbonates including marine cements, calcified cyanobacteria and thick precipitated beds with fibrous textures that are extremely rare in Phanerozoic aged sediments98,99. More specifically, in contrast to the standard LOSCAR parameterization where calcite supply to the sediment water interface is tied to the biological pump, we modified the carbonate supply term within all ocean basins to be driven by the degree of saturation, using the experimental kinetic data presented in ref. 100. Net accumulation (preservation) in each ocean basin is achieved only if formation exceeds chemical erosion (dissolution) in the model.

Modelling reverse weathering using LOSCAR-RW

To parameterize reverse weathering we coupled an early diagenetic model (for example, see Extended Data Figs. 6 and 7) with our modified version of LOSCAR. Building from the accepted view of anoxic and iron-rich Precambrian oceans, the upwelling of dissolved phosphorus (P) and ferrous iron governs the export of organic matter and Fe-oxyhydroxides from surface waters respectively57,76,101. With deep iron-rich seawater there would be inorganic scavenging of bioavailable P from surface oceans (a deep-sea P trap), for which, following ref. 101, we have modelled a percentage scavenging efficiency. We used a modified anoxic organic remineralization power law102,103 to determine the export efficiency of organic matter to the sediment–water interface, with water column remineralization coupled directly to the reduction of sinking Fe-oxyhydroxide particulates. Tracers in the diagenetic model include solid (organic matter, Fe-oxyhydroxide and calcite) and dissolved phases (CO 2 HCO 3 −, CO 3 2−, pH, alkalinity, Ca and Si) within the topmost metre of the sediment column. We assume that anaerobic respiration governs organic matter remineralization (tied to the reduction of Fe-oxyhydroxide) and reverse weathering governs the consumption of dissolved Si.

$$\frac{\partial \left[{\rm{X}}\right]}{\partial t}=D\frac{{\partial }^{2}\left[{\rm{X}}\right]}{\partial {z}^{2}}-\omega \frac{\partial \left[{\rm{X}}\right]}{\partial z}-{\rm{rxn}}$$ (8)

The sum reaction of the time rate of change of any dissolved constituent (X) in porewaters can be described as the sum of the diffusion term (\(D\frac{{\partial }^{2}\left[{\rm{X}}\right]}{\partial {z}^{2}}\)), the advection term (\(\omega \frac{\partial \left[{\rm{X}}\right]}{\partial z}\)), and the reaction term (rxn), where D is the diffusion coefficient, ω is the sedimentation rate and z the depth in the sediment104,105. Steady state (\(\frac{\partial \left[{\rm{X}}\right]}{\partial {\rm{t}}}=0\)) is achieved in all model runs. The reverse-weathering reaction term is described as \({k}_{{\rm{rw}}}\left[{\rm{Si}}-{{\rm{Si}}}_{{\rm{o}}}\right]\), where k rw is the reverse-weathering reaction rate constant and Si o is the lower boundary dissolved silica condition (equilibrium concentration). From first principles there should be a pH control on authigenic clay formation (for example, refs 23,24,70,106,107,108,109,110), which has been observed in experimental work (for example, refs 19,24,70,111,112,113,114). The degree (order) of dependence on pH scales with the Alk:Si consumption ratio of the clay species. As such, we relate k rw and Si o to porewater pH conditions according to the expressions k rw = 1.01 × 10−19 × pH22.4 yr−1 and Si o = 2.02−5.57 × pH mol l−1, based on previously presented experimental data24, which explicitly relate pH and the extent of reverse weathering. This formulation is the basis for a stabilizing feedback on \({p}_{{{\rm{CO}}}_{{\rm{2}}}}\). These expressions determine the strength of the reverse-weathering pH buffering effect within marine porewater systems (which are in diffusional exchange with bottom seawater; for instance, Extended Data Fig. 6) and thus controls the buffering capacity of the whole ocean system. The reaction term for calcite is described as \({k}_{{\rm{calc}}}\left[{\Omega }_{{\rm{calc}}}-1\right]\), where Ω calc is the saturation state. Organic matter remineralization is parameterized as \({k}_{{\rm{org}}}\times \left[{\rm{org}}\right]\times \frac{\left[{\rm{Fe}}{\left({\rm{OH}}\right)}_{3}\right]}{\left[{\rm{Fe}}{\left({\rm{OH}}\right)}_{3}\right]+{\left[{\rm{Fe}}{\left({\rm{OH}}\right)}_{3}\right]}_{lim}}\), where \({[{\rm{Fe}}{\left({\rm{OH}}\right)}_{3}]}_{{\rm{lim}}}\) is the limiting concentration115 and ‘[org]’ is the concentration of organic matter. The total flux of dissolved silica consumed in sedimentary porewater systems can be described according to Fick’s first law at the sediment–water interface such that:

$${F}_{{\rm{rw}}}=-\phi D\frac{\partial \left[{\rm{Si}}\right]}{\partial z}\Lambda $$ (9)

where φ is sediment porosity and Λ is the authigenically ‘active’ area of the seafloor. The total flux of alkalinity consumed (or acidity generated) is then simply:

$${{\rm{Alk}}}_{{\rm{rw}}}=-\left({\rm{Alk:Si}}\right){F}_{{\rm{rw}}}$$ (10)

Alk:Si relates total alkalinity to the flux of dissolved silica consumed through reverse weathering in porewaters. In LOSCAR-RW, reverse weathering is parameterized as the removal flux of alkalinity (Alk rw ) across the sediment–water interface within each environment. Specifically, the deep sea, slope and shelf sedimentary environments are linked to the deep, intermediate and surface ocean boxes respectively. Steady-state outputs of marine pH, total dissolved inorganic carbon (DIC), total alkalinity and calcite saturation state that correspond to the results presented in Fig. 1 are presented in Extended Data Fig. 8. Sensitivity analyses to changes in the strength of the silicate weathering feedback (n Si ) between 0.1–0.511,20,46,116,117,118,119, are presented in Extended Data Fig. 9. We base all our preferred model runs on experimental data and previous constraints on Precambrian marine conditions (see Supplementary Table 1 for full list of parameter values). We find that rates of reverse weathering over the range of conditions previously predicted for the Precambrian are most sensitive to changes in marine pH and the concentration of dissolved silica in seawater (Extended Data Figs. 4, 6–10). We note, however, that any changes to the globally integrated authigenic clay assemblage through time have the potential to alter the pH, dissolved Si, reaction rate law relationships, and hence the extent and pH buffering potential of the reverse-weathering process that was adopted in this study. For instance, from a thermodynamic viewpoint, the sensitivity of clay authigenesis to ambient pH conditions ought to increase with a global assemblage with a higher Alk:Si consumption ratio, potentially altering the buffering strength of the entire ocean system. Reconstructing the evolution of the Alk:Si consumption ratio throughout Earth’s history is therefore a critical topic for future work. Nonetheless, these sensitivity tests suggest that our central conclusion—that reverse weathering rates would have been elevated in Si-rich oceans leading to more stable and elevated atmospheric CO 2 conditions—is robust over a broad range of parameter space.

Data availability

All data collected for this study are included in this published article (and the Supplementary Information).