Exploring reverse silicate weathering across geological time: a review

Abstract

Marine clay mineral authigenesis, referred to as reverse (silicate) weathering, is one of the first-order controls on seawater pH through the generation of acidity and thus plays a significant role in controlling carbon cycling between marine sediments, oceans and the atmosphere over geological timescales. Reverse weathering is mainly regulated by the rates of silicate and carbonate weathering on the continents, the reactivity of detritus supplied to the oceans and the rates of seafloor weathering. These processes provide essential dissolved components (e.g. K⁺, Mg²⁺, Ca²⁺, Si(OH)₄, Al³⁺, Fe^2+/3+) to the marine porewater inventory that cause authigenic clay minerals, such as odinite, glauconite, celadonite and greenalite, to form close to the sediment–seawater interface. Such clay mineral reactions impact the sedimentary cycling versus sequestration of chemical elements, importantly Si, Fe, Mg and K, and consequently contribute to the fluctuations in climate and seawater composition recorded in marine archives over geological time. This review explores the links between reverse silicate weathering and the climate system across geological timescales and provides estimates of the elemental uptake fluxes associated with modern-day clay mineral authigenesis. Novel isotope proxies (e.g. δ⁴¹K and δ³⁰Si) and promising new dating techniques (e.g. in situ Rb/Sr geochronology) provide improved constraints on the timing, kinetics and environmental significance of clay mineral reactions on the ocean floor. We also consider recent geoengineering developments linked to reverse weathering reactions, such as ongoing attempts to reduce atmospheric CO₂ concentrations via marine alkalinity enhancement and the application of marine clay mineral-based slow-release fertilizers to soils to optimize nutrient availability.

Retrograde clay mineral formation on the ocean floor, also called reverse weathering, has strongly influenced the mineralogical and biogeochemical composition of Earth’s marine sediments, the chemistry of its seawater and the composition of its atmosphere across geological time (Mackenzie & Garrels, 1966a, 1966b; Mackin & Aller, 1984; Siever, 1992; Mackenzie & Kump, 1995; Michalopoulos & Aller, 1995; Misra & Froehlich, 2012; Isson & Planavsky, 2018; Li et al., 2021a; Baldermann et al., 2022; Rauzi et al., 2024). In this process, porewater- or seawater-derived Si(OH)₄ reacts with HCO₃^– ions, dissolved cations and metal aquo-complexes, such as Fe^2+/3+, Mg²⁺, K⁺, Na⁺, Ca²⁺ and Al(OH)₃⁰, to form a suite of authigenic clay minerals, as well as variable amounts of acidity (H⁺) during early marine diagenesis (e.g. Siever, 1992; Michalopoulos et al., 2000; Hazen et al., 2013; Tosca et al., 2016; Cuadros et al., 2017; Frings, 2017; Cao et al., 2022; Ma et al., 2022; Steiner et al., 2022; Wallmann et al., 2023). As a result, near-surface clay mineral authigenesis impacts the evolution of seawater pH and the marine element cycles, and potentially serves as a long-term regulator of Earth’s climate by influencing atmospheric CO₂ concentrations (e.g. Kump et al., 2000; Isson & Planavsky, 2018; Andrews et al., 2020; Isson et al., 2020; Krissansen-Totton & Catling, 2020; Li et al., 2021a; Warr, 2022; Yin et al., 2023). Reverse weathering reactions have been documented in a diverse range of marine settings, including ‘low’-temperature shallow shelf and deep-marine environments, as well as ‘high’-temperature (up to a few hundred °C) settings associated with mid-oceanic ridge sites, or transform faults. Common examples of clay mineral precipitation reactions are given in Equation 1a–e (modified after Isson & Planavsky, 2018):

(1a)\begin{align} &{\text{Celadonite:}}\;{\text{1}}{\text{.0}}{{\text{K}}^{\text{\! + }}}{\text{ + 1}}{\text{.2F}}{{\text{e}}^{{\text{3\! +\! /2\! + }}}}{\text{ + 0}}{\text{.8M}}{{\text{g}}^{{\text{2\! + }}}}{\text{ + 3}}{\text{.8Si}}{\left( {{\text{OH}}} \right)_{\text{4}}} \nonumber\\ &\quad {\text{ + 0}}{\text{.2Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + 6}}{\text{.2HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \to {\text{K}}\left( {{\text{F}}{{\text{e}}^{{\text{3\! +\! /2 + }}}}_{{\text\!{1}}{\text{.2}}}{\text{M}}{{\text{g}}_{{\text{0}}{\text{.8}}}}} \right) \nonumber\\ &\quad \left[ {{\text{A}}{{\text{l}}_{{\text{0}}{\text{.2}}}}{\text{S}}{{\text{i}}_{{\text{3}}{\text{.8}}}}} \right]{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{2}}}{\text{ + 6}}{\text{.2C}}{{\text{O}}_{\text{2}}}{\text{ + 10}}{{\text{H}}_{\text{2}}}{\text{O}} \end{align}

(1b)\begin{align} &{\text{Glauconite:}}\;{\text{0}}{\text{.6}}{{\text{K}}^{\text{\! + }}}{\text{ + 0}}{\text{.05N}}{{\text{a}}^{\text{\! + }}}{\text{ + 1}}{\text{.6F}}{{\text{e}}^{{\text{3\! +\! /2\! + }}}}{\text{ + 0}}{\text{.4M}}{{\text{g}}^{{\text{2\! + }}}} \nonumber\\ &\quad {\text{ + 3}}{\text{.8Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 0}}{\text{.2Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + 6}}{\text{.2HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \to \left( {{{\text{K}}_{{\text{0}}{\text{.6}}}}{\text{N}}{{\text{a}}_{{\text{0}}{\text{.05}}}}} \right) \nonumber\\ &\quad \left( {{\text{F}}{{\text{e}}^{{\text{3\! +\! /2\! + }}}}_{{\!\text{1}}{\text{.6}}}{\text{M}}{{\text{g}}_{{\text{0}}{\text{.4}}}}} \right)\left[ {{\text{A}}{{\text{l}}_{{\text{0}}{\text{.2}}}}{\text{S}}{{\text{i}}_{{\text{3}}{\text{.8}}}}} \right]{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{2}}}{\text{ + 6}}{\text{.2C}}{{\text{O}}_{\text{2}}}{\text{ + 10}}{{\text{H}}_{\text{2}}}{\text{O}} \end{align}

(1c)\begin{align}&{\text{Greenalite:}}\;{\text{3F}}{{\text{e}}^{{\text{3\! + }}}}{\text{ + 2Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 6HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \nonumber\\ &\quad \to {\text{ }}\left( {{\text{F}}{{\text{e}}^{{\text{3\! + }}}}_{\text{3}}} \right)\left[ {{\text{S}}{{\text{i}}_{\text{2}}}} \right]{{\text{O}}_{\text{5}}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 6C}}{{\text{O}}_{\text{2}}}{\text{ + 5}}{{\text{H}}_{\text{2}}}{\text{O}}\end{align}

(1d)\begin{align} &{\text{Odinite: 1}}{\text{.3F}}{{\text{e}}^{{\text{2\! +\! /3\! + }}}}{\text{ + 0}}{\text{.7M}}{{\text{g}}^{{\text{2\! + }}}}{\text{ + 1}}{\text{.8Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 0}}{\text{.7Al}}{\left( {{\text{OH}}} \right)_{\text{3}}} \nonumber\\ &\quad {\text{ + 4}}{\text{.3HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \to \left( {{\text{F}}{{\text{e}}^{{\text{2\! +\! /3\! + }}}}_{{\!\text{1}}{\text{.1}}}{\text{M}}{{\text{g}}_{{\text{0}}{\text{.7}}}}{\text{A}}{{\text{l}}_{{\text{0}}{\text{.7}}}}} \right) \nonumber\\ &\quad \left[ {{\text{F}}{{\text{e}}^{{\text{3\! + }}}}_{{\!\text{0}}{\text{.2}}}{\text{S}}{{\text{i}}_{{\text{1}}{\text{.8}}}}} \right]{{\text{O}}_{\text{5}}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 4}}{\text{.3C}}{{\text{O}}_{\text{2}}}{\text{ + 5}}{{\text{H}}_{\text{2}}}{\text{O}} \end{align}

(1e)\begin{align} &{\text{Saponite:}}\;{\text{0}}{\text{.15C}}{{\text{a}}^{{\text{2\! + }}}}{\text{ + 0}}{\text{.1N}}{{\text{a}}^{\text{\! + }}}{\text{ + 2}}{\text{.5M}}{{\text{g}}^{{\!\text{2 + }}}}{\text{ + 0}}{\text{.8F}}{{\text{e}}^{{\text{2\! + }}}} \nonumber\\ &\quad {\text{ + 3Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + 7HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \to \left( {{\text{C}}{{\text{a}}_{{\text{0}}{\text{.15}}}}{\text{N}}{{\text{a}}_{{\text{0}}{\text{.1}}}}} \right) \nonumber\\ &\quad \left( {{\text{M}}{{\text{g}}_{{\text{2}}{\text{.5}}}}{\text{F}}{{\text{e}}^{{\text{2\! + }}}}_{{\!\text{0}}{\text{.8}}}} \right)\left[ {{\text{AlS}}{{\text{i}}_{\text{3}}}} \right]{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{2}}}{\text{ + 7C}}{{\text{O}}_{\text{2}}}{\text{ + 10}}{{\text{H}}_{\text{2}}}{\text{O}} \end{align}

Many of these authigenic marine clay minerals are Fe-rich, have a greenish colour and show a high but temporally restricted abundance in the rock record. Some of them even form distinct facies, such as the well-known glaucony and verdine facies (e.g. Rubio & López-Pérez, 2024). These unique Fe-rich clay minerals preferentially precipitate in semi-confined geochemical micro-environments, such as in faecal pellets or foraminifera chambers, within marine sedimentary sequences that cover large areas of modern shelf and deep-sea areas and characterized equivalent ancient environments (e.g. Baldermann et al., 2012, 2022; Banerjee et al., 2015, 2016).

In the geological record, the occurrence of large deposits of authigenic marine clay minerals, such as ironstones and greensands, is often found to be associated with high rates of bio-opal remineralization (Michalopoulos & Aller, 2004). This relationship suggests a close linkage between the global silica and other element cycles – importantly, C, K, Mg and Fe (e.g. Pogge von Strandmann et al., 2013; Dunlea et al., 2017; Baldermann et al., 2022; Isson & Rauzi, 2024). However, quantifying the role of reverse weathering in global biogeochemical cycles and the climate system remains difficult because the mechanisms and rates of individual reactions are often not known (e.g. Ehlert et al., 2016; Rahman et al., 2016; Krissansen-Totton & Catling, 2020). Nevertheless, there is ample evidence that marine clay mineral authigenesis has had a strong effect on Earth’s climatic evolution and the marine sedimentary element cycles – relationships that are discussed in this article. Moreover, this article reviews the temporal distribution and abundance of authigenic clay minerals in marine sedimentary archives from the Precambrian to recent epochs by focusing on the formation of beidellite, berthierine, chamosite, celadonite, glauconite, greenalite, kaolinite, minnesotaite, odinite, palygorskite, saponite, sepiolite, stevensite, stilpnomelane, talc and vermiculite (Fig. 1). Recent advances in trace element and isotope geochemistry of authigenic clay minerals are also presented, and examples are provided highlighting the importance of clay mineral formation versus weathering in marine geoengineering and in slow-release fertilizer technology.

Figure 1. Illustration showing selected sedimentary environments characterized by authigenic clay mineral formation. Common clay mineral assemblages forming in each depocentre are indicated at the top right. The yellow meandering strands indicate deltas and subaquatic channels. International Mineralogical Association (IMA)-approved clay minerals abbreviations after Warr (2020): Alp = allophane; Bei = beidellite; Brh = berthierine; Cel = celadonite; Chl = chlorite; Chl-Sme = chlorite-smectite; Chm = chamosite; Glt = glauconite; Gre = greenalite; Hsg = hisingerite; Ilt = illite; Ilt-Sme = illite-smectite; Imo = imogolite; Kln = kaolinite; Mns = minnesotaite; Non = nontronite; Odn = odinite; Plg = palygorskite; Sap = saponite; Sep = sepiolite; Sme = smectite; Stp = stilpnomelane; Stv = stevensite; Tlc = talc; Vrm = vermiculite.

Reverse weathering scenarios and climate evolution across geological time

Weathering, erosion and biodegradation are the main mechanisms that cause the breakdown of rocks and minerals on Earth’s surface. These processes trigger congruent and incongruent dissolution of primary siliceous, calcareous and other minerals, providing dissolved elements, colloids, degraded minerals and solid matter to rivers, groundwaters and surface waters, which ultimately enter the oceans (Fig. 2). Thus, by regulating supply, weathering reactions on the continents play a central role in controlling ocean (bio)geochemical reactions at the sediment–seawater interface and in the marine element cycle(s) through time (e.g. Singer, 1980, 1984; Michalopoulos & Aller, 1995; Isson & Planavsky, 2018; Isson et al., 2020; Kalderon-Asal et al., 2021).

Figure 2. Illustration showing the complex interplay between terrestrial (silicate) weathering, marine bioproductivity, reverse weathering reactions, ocean biogeochemistry and atmospheric circulation, which all contribute to Earth’s climatic evolution (after Isson & Planavsky, 2018). The yellow meandering strands indicate deltas and subaquatic channels.

Silicate minerals compose over 90% of the rocks found in Earth’s near-surface environments. As most of them are unstable under ambient pressure and temperature conditions, they are susceptible to various forms of physical alteration and chemical degradation (e.g. Wilson, 1999, 2004). The incongruent dissolution of primary igneous silicate minerals, such as olivine, pyroxene, amphibole, feldspar and mica, and the resulting supply of cations are therefore often viewed as prerequisites for the formation of secondary clay minerals in the critical zone (i.e. soils) and in the widespread marine realm (e.g. Banfield et al., 1991; Wilson, 1999, 2004; Gillis & Coogan, 2011; Rafiei et al., 2020). These reactions are ultimately linked to regulating Earth’s climate and its carbon cycle over the long term. There are various numerical models available (reviewed by Penman et al., 2020) that aim at understanding and quantifying (at least parts of) the complex feedback mechanisms and rates between silicate weathering, reverse weathering and climate change. These consider changes in parameters through geological time, such as the partial pressure of CO₂ in the atmosphere (pCO₂), the rates of continental and seafloor weathering and erosion (e.g. Coogan & Gillis, 2013, 2018, 2020), vegetational cover, topography and temperature, among other factors (e.g. Penman et al., 2020). The key parameters controlling climate change and the palaeoclimatic evolution of Earth remain, however, focuses of continued debate. This reflects the complexity of quantifying silicate dissolution (CO₂ sink) versus clay mineral formation processes (CO₂ source) occurring in Earth’s broad spectrum of near-surface environments and the resulting limitations in modelling and predicting global elemental fluxes and their sensitivity to climate (e.g. Keller & Wood, 1993; Isson et al., 2020; Krissansen-Totton & Catling, 2020). On geological timescales, silicate weathering and clay mineral cycles are also linked to plate tectonic processes, with faster subduction cycles occurring along active plate margins and slower cycles taking place on passive continental margins (reviewed by Warr, 2022). This means that there is an intimate relationship between silicate weathering, clay mineral precipitation, climate and the tectono-thermal behaviour of the lithosphere.

It is well-established that Earth’s early climate was characterized by warmth and stability, with ice-free conditions prevailing during most of the Proterozoic eon (2,500–541 Myr ago), except for the extreme glaciation events in the Cryogenian (e.g. Lipp et al., 2021; Isson & Rauzi, 2024). This occurred despite the solar luminosity being ∼5–20% lower than in the present day (e.g. Gough, 1981). The long-lasting climate stability changed dramatically in the Phanerozoic eon (541 Myr–present), with the appearance of frequent and rapid icehouse–greenhouse cycles (e.g. Royer et al., 2004). The reasons for this climate transition have been controversially discussed in the literature. Proposed explanations include temporal changes in Earth’s crustal composition (e.g. Gaillardet et al., 1999; Coogan & Dosso, 2015), increased uplift and erosional rates, as well as changes in ocean currents (e.g. Foster & Rohling, 2013), which had maintained elevated pCO₂ during the Precambrian (e.g. Kump et al., 2000; Isson & Planavsky, 2018; Sheldon et al., 2021). Others have suggested that large variations in seawater pH and certain processes, such as carbonate mineral precipitation and reverse weathering reactions (Equation 1a–e), helped to regulate the global carbon–silica cycle in the Precambrian and thus the input and output fluxes of C and Si between marine sediment sinks, the ocean and the atmosphere (Fig. 2; e.g. Sun & Turchyn, 2014; Isson & Planavsky, 2018; Kasting, 2019; Isson et al., 2020; Krissansen-Totton & Catling, 2020; Torres et al., 2022). All of the mechanisms discussed above are expected to impact atmospheric CO₂ concentrations and climate stability (e.g. Royer et al., 2004).

CO₂ fluxes related to silicate weathering

Five main mechanisms control clay mineral-related fluxes of CO₂ to the atmosphere, which are described in detail by Warr (2022). These are: (1) silicate weathering on the continents, which creates clay mineral covers and crusts that can slow down the rate of rock alteration; (2) organic–clay mineral complexation, which helps to pass dissolved carbon species into the oceans; (3) clay mineral-based pH buffering, which aids carbonate precipitation in the oceanic environment; (4) reverse weathering reactions, causing clay mineral precipitation onto the seafloor; and (5) crustal diagenesis, metamorphism, melting and volcanism in the upper lithosphere, where (4) and (5) release water and other volatiles (e.g. CO₂) from transforming clay minerals (Fig. 3).

Figure 3. Illustration of the interaction of clay minerals, silicate weathering and atmospheric CO₂ levels. (1) Formation of continental surface clay mineral deposits and weathered crusts consumes CO₂ and slows down rock alteration; (2) organic–clay minerals complexation fixes carbon and passes it into the oceans; (3) clay mineral buffers aid carbonate mineral precipitation in the marine environment; (4) reverse weathering and clay mineral precipitation occur on the seafloor, releasing CO₂; and (5) the release of water and volatiles from clay minerals during subduction enhances crustal melting and volcanism (Warr, 2022, with minor modifications).

Within this context, Earth’s climate today is controlled mainly by the (im)balance between CO₂ emissions to the atmosphere and carbon burial by silicate weathering and carbonate formation (e.g. Friedlingstein et al., 2010; Colbourn et al., 2015; Coogan & Gillis, 2020). Carbon removal by marine biogenic and diagenetic–abiogenic carbonate mineral formation consumes the majority of the atmospheric CO₂ inventory (e.g. Torres et al., 2022), as well as some of the ocean-bound carbonate alkalinity sourced from silicate and carbonate weathering on the continents and the deep ocean floor (e.g. Walker et al., 1981; Colbourn et al., 2015). It is believed that the long-term carbon cycle and thus Earth’s climate are stabilized by a negative-feedback mechanism known as the silicate weathering thermostat (e.g. Berner et al., 1983; Brantley et al., 2023), which involves temperature- or climate-sensitive atmospheric CO₂ consumption via the chemical weathering of susceptible silicate minerals (Figs. 2 & 3; e.g. Walker et al., 1981; Kasting, 2019; Jin et al., 2023).

In this widely accepted negative-feedback climate model, silicate weathering is enhanced when pCO₂ and temperature increase, leading to increased consumption of atmospheric CO₂ and hence a cooling effect. As pCO₂ falls, the silicate weathering rate decreases, permitting the build-up of atmospheric CO₂ and consequent warming (e.g. Isson & Planavsky, 2018; Kasting, 2019; Penman et al., 2020). The timescale of the climate response related to changes in silicate weathering is considered to lie in the range of 170–380 kyr, with an average response time of ∼240 kyr (Colbourn et al., 2015). As a direct consequence of the described negative-feedback loop, continental and seafloor weathering of primary igneous silicate minerals, such as pyroxene, olivine, amphibole, feldspar and phyllosilicates, is currently considered as a key reaction step for regulating Earth’s climate and seawater composition (e.g. Coogan & Gillis, 2013, 2018). As an example, this can be represented by the reactions show in Equation 2a–c using wollastonite (CaSiO₃; e.g. Walker et al., 1981; Berner et al., 1983; Schott et al., 2012; Farkaš et al., 2025):

(2a)\begin{equation}{\text{2C}}{{\text{O}}_{\text{2}}}{\text{ + 3}}{{\text{H}}_{\text{2}}}{\text{O + CaSi}}{{\text{O}}_{\text{3}}} \to {\text{ C}}{{\text{a}}^{{\text{2\! + }}}}{\text{ + 2HC}}{{\text{O}}_{\text{3}}}^{\text{--}}{\text{ + Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}\end{equation}

(2b)\begin{equation}{\text{C}}{{\text{a}}^{{\text{2\! + }}}}{\text{ + 2HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \to {\text{ CaC}}{{\text{O}}_{\text{3}}}{\text{ + C}}{{\text{O}}_{\text{2}}}{\text{ + }}{{\text{H}}_{\text{2}}}{\text{O}}\end{equation}

(2c)\begin{equation}{\text{Net: C}}{{\text{O}}_{\text{2}}}{\text{ + 2}}{{\text{H}}_{\text{2}}}{\text{O + CaSi}}{{\text{O}}_{\text{3}}} \to {\text{ CaC}}{{\text{O}}_{\text{3}}}{\text{ + Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}\end{equation}

These reactions illustrate how released bivalent cations (mainly Ca²⁺, but also Ba²⁺, Mg²⁺, Mn²⁺, Fe²⁺, Sr²⁺ and Zn²⁺) and carbonate alkalinity are finally deposited as CaCO₃ minerals in marine sediments, causing a net reduction of 1 mol of atmospheric CO₂ per 1 mol of CaCO₃ formed. In addition, dissolved cations (including also Al³⁺, Na⁺, K⁺ and Li⁺, among others) and Si(OH)₄, if transported to the ocean, can trigger the formation of biogenic opal and authigenic clay minerals in various shallow and deep marine settings (Fig. 1; e.g. Michalopoulos & Aller, 1995; Isson & Planavsky, 2018; Baldermann et al., 2022). However, we note here that clay mineral formation also occurs in various continental settings, such as in soils and lakes (e.g. Huggett & Cuadros, 2010; Bristow et al., 2012; Pozo & Calvo, 2018), although none of these depositional environments belong (by definition) to clay minerals formed through reverse weathering reactions. In any case, the rates of silicate weathering, CO₂ uptake and carbonate mineral precipitation are temperature dependent and therefore can be viewed as representing a planetary thermostat or climate regulator. This is indicated, for example, by the climate-sensitive weathering rates of north-eastern Iceland river catchments (e.g. Gislason et al., 2009) or the terrestrial climate archive from the Valley of Lakes (Mongolia, Central Asia; see Fig. 4; e.g. Baldermann et al., 2021). There, variations in temperature and climatic conditions during the Cenozoic were directly related to pulses of intensified physical versus chemical weathering, as is evident by the correlation between the Chemical Index of Alteration (CIA) and excursions of oxygen isotopes (δ¹⁸O) recorded in pedogenic carbonates (e.g. Richoz et al., 2017). The Valley of Lakes continental archive thus provides a sensitive record of changing hydroclimatic conditions, as calcrete (i.e. authigenic calcite) deposited in palaeosols resulted in systematically lighter δ¹⁸O values during wetter conditions and relatively heavier δ¹⁸O values during incursions of drier conditions. Analogously, increased chemical weathering indices (here: the CIA) correspond to periods of wetter conditions, and vice versa (Baldermann et al., 2021).

Figure 4. Regional response of pedogenic carbonate oxygen isotope (δ¹⁸O_carbonate) geochemistry and silicate-based CIA associated with Cenozoic climate change in the Valley of Lakes, Mongolia (source: Baldermann et al., 2021). Note the high sensitivity of the study site to regional and global climatic variability during this time, which is expressed as a higher abundance of detrital illite (Ilt) versus authigenic illite-smectite (Ilt-Sme) in periods of increased physical versus chemical weathering (see scanning electron microscopy images on the left). Oi-1a/b = Oi-1a/b glaciation (∼34–33 Myr); EOA = early Oligocene aridification (∼31 Myr); OGM = Oligocene glacial maximum (∼28 Myr); LOW = late Oligocene warming (∼25 Myr); Mi-1 = Mi-1 glaciation (∼23 Myr). Aqu = Aquitanian; Bar = Bartonian; Bur = Burdigalian; Cha = Chattian; Lan = Langhian; Lut = Lutetian; Pri = Priabonian; Rup = Rupelian; Ser = Serravallian; V-PDB = Vienna Peedee Belemnite; Ypr = Ypresian.

The reactivity of Earth’s crust (e.g. Gaillardet et al., 1999; Coogan & Dosso, 2015) is also modulated by biological factors, such as the activities of terrestrial plants or soil microbes. As these factors may accelerate or slow down the rates of physical and chemical weathering (e.g. Cuadros, 2017), they are ultimately linked to the regulation of the carbon cycle. These became more important since the advent of biologically active soils and biogenic sediments approximately in the late Neoproterozoic (e.g. Kennedy et al., 2006; Rafiei et al., 2020), as well as the rise of land plants in the Palaeozoic (e.g. Lenton et al., 2012; Porada et al., 2016; McMahon & Davies, 2018; Davies et al., 2020; Isson & Rauzi, 2024). Furthermore, tectonic uplift and topographic elevation (e.g. the Himalayas and the Tibetan Plateau), combined with high precipitation rates, further increased the rate of silicate chemical weathering (and CO₂ consumption) by up to 40–60% (e.g. Raymo & Ruddiman, 1992; Kump & Arthur, 1997; Xu et al., 2024). In addition, the age of an orogen or a mountain belt impacts the rate of silicate weathering. Young orogens weather faster than older ones due to the rapid exhumation of fresh lithologies that are less resistant to chemical weathering processes (e.g. Maher & Chamberlain, 2014).

To summarize, it is widely accepted that changes to global continental weathering have led to fluctuations in atmospheric CO₂ over geological timescales (e.g. Berner, 1990). However, it remains unclear how the combination of tectonic processes, climate change and rock lithologies, among other factors, has influenced the rate of silicate weathering on a global scale (e.g. Hilley & Porder, 2008). Hilley & Porder (2008) calculated a global annual Si weathering flux of 19–46 Tmol, obtained by considering local erosion rates, dust fluxes, temperature and the water balance of various regions worldwide. These authors concluded that the uplift of the Himalayas since the late Cenozoic has had a significant impact on Earth’s climate. Similarly, (Caves Rugenstein et al., 2018) have argued that the Cenozoic global cooling (∼52 Myr to the present day) was mainly triggered by the increasing exposure of highly reactive lithologies exposed in the Himalayas and the Tibetan Plateau. This underlines the importance of the global weathering feedback to Earth’s climate evolution, and vice versa (Fig. 4; e.g. Caves et al., 2016; Baldermann et al., 2021; Jin et al., 2023). Recently, Lipp et al. (2021) have estimated global weathering intensities based on the evolution of the continents’ bulk composition over geological time and proposed that silicate weathering was intensified in the Archaean (∼25% more CO₂ sequestration between ∼4000 and 2500 Myr). This was followed by reduced but relatively constant rates of silicate weathering during the Proterozoic (∼2500–541 Myr) and the Phanerozoic (<541 Myr to present). However, these authors also identified peaks in the weathering intensities on shorter timescales within the Phanerozoic, such as during the Carboniferous, Triassic and Cretaceous periods (see Fig. 3 in Lipp et al., 2021), illustrating how silicate weathering acts in response to changes in atmospheric CO₂. Similar conclusions were recently drawn by Isson & Rauzi (2024), who have argued that the evolution of Earth’s global climate, seawater pH and changes in silicate weathering triggered distinct pulses or declines in early clay mineral diagenesis (i.e. continental and reverse weathering), which are detailed further in later sections of this article.

Early clay mineral diagenesis and climate feedback

Reverse weathering involves the formation of authigenic clay minerals in near-surface marine sediments, a process that consumes carbonate alkalinity and produces acidity and leads to a net release of CO₂ (cf. Equation 1a–e & Fig. 2; e.g. Farkaš et al., 2025). It can proceed via entirely de novo precipitation of clay minerals, but it can also proceed via the partial dissolution of detrital precursor phases, followed by cation enrichment, substitution and restructuring as a ‘new phase’ (e.g. Garrels, 1965; Mackenzie & Garrels, 1966b; Mackin & Aller, 1984). The latter pathway is illustrated by the transformation of kaolinite to berthierine with no additional Si(OH)₄ being consumed in this reaction (see Equation 3; Bhattacharyya, 1983):

(3)\begin{align} &{\text{A}}{{\text{l}}_{\text{4}}}{\text{S}}{{\text{i}}_{\text{4}}}{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{8}}} {\text{ + 3F}}{{\text{e}}^{{\text{2\! + }}}}{\text{ + 6HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \nonumber\\ &\quad \to {\text{F}}{{\text{e}}_{\text{3}}}{\text{A}}{{\text{l}}_{\text{2}}}{\text{S}}{{\text{i}}_{\text{4}}}{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{8}}}{\text{ + 2Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + 6C}}{{\text{O}}_{\text{2}}}\end{align}

Reverse weathering is thought to have exerted a significant influence on global climate and marine elemental cycling in the past and in the present (e.g. Berner et al., 1983; Mackenzie & Kump, 1995; Baldermann et al., 2022; Isson & Rauzi, 2024), given that the formation of authigenic clay minerals releases CO₂ and impacts on Earth’s climate and seawater pH (e.g. Mackenzie & Garrels, 1966b; Kump et al., 2000; Isson & Planavsky, 2018). In simple terms, reverse weathering reactions remobilize carbon species in the ocean–atmosphere system (e.g. Isson & Planavsky, 2018; Farkaš et al., 2025) and act to counterbalance (reverse) the sequestration of carbon linked to silicate weathering (e.g. West et al., 2005). This process can be illustrated via the example of kaolinite formation by incongruous weathering of silicates followed by kaolinite alteration to authigenic illite. A standard example of silicate weathering occurring on the continents is the reaction of anorthite (CaAl₂Si₂O₈) with dissolved CO₂ (i.e. carbonic acid) at pH 4–6 and at high leaching rates to form kaolinite and alkalinity in (sub)tropical soils (e.g. West et al., 2005; Galán, 2006), according to the reaction in Equation 4:

(4)\begin{align}&{\text{2C}}{{\text{O}}_{\text{2}}}{\text{ + 3}}{{\text{H}}_{\text{2}}}{\text{O + CaA}}{{\text{l}}_{\text{2}}}{\text{S}}{{\text{i}}_{\text{2}}}{{\text{O}}_{\text{8}}} \nonumber\\ &\quad \to {\text{ C}}{{\text{a}}^{{\text{2 + }}}}{\text{ + 2HC}}{{\text{O}}_{\text{3}}}^{\text{--}}{\text{ + A}}{{\text{l}}_{\text{2}}}{\text{S}}{{\text{i}}_{\text{2}}}{{\text{O}}_{\text{5}}}{\left( {{\text{OH}}} \right)_{\text{4}}}\end{align}

When delivered to marginal marine settings, such as mangrove forests, either 1 or 2 moles of soil-derived, detrital kaolinite (depending on the Si flux) can react further with dissolved cations, provided from saprolite weathering and marine porewater ingress, to form 1 mole of Fe(III)-illite and other components through kaolinite-illitization according to Equation 5a–b (Cuadros et al., 2017):

(5a)\begin{align} &{\text{A}}{{\text{l}}_{\text{2}}}{\text{S}}{{\text{i}}_{\text{2}}}{{\text{O}}_{\text{5}}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 1}}{\text{.43Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + 0}}{\text{.57M}}{{\text{g}}^{{\text{2\! + }}}}{\text{ + 1}}{\text{.16F}}{{\text{e}}^{{\text{3\! +\! /2\! + }}}} \nonumber\\ &\quad {\text{ + 0}}{\text{.68}}{{\text{K}}^{\text{\! + }}} \to \left( {{{\text{K}}_{{\text{0}}{\text{.68}}}}{\text{M}}{{\text{g}}_{{\text{0}}{\text{.06}}}}} \right)\left( {{\text{F}}{{\text{e}}_{{\text{1}}{\text{.16}}}}{\text{A}}{{\text{l}}_{{\text{0}}{\text{.66}}}}{\text{M}}{{\text{g}}_{{\text{0}}{\text{.51}}}}} \right) \nonumber\\ &\quad \left[ {{\text{A}}{{\text{l}}_{{\text{0}}{\text{.57}}}}{\text{S}}{{\text{i}}_{{\text{3}}{\text{.43}}}}} \right]{{\text{O}}_{{\text{10}}}}{\left( {{\text{OH}}} \right)_{\text{2}}}{\text{ + 0}}{\text{.77Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + 4}}{\text{.59}}{{\text{H}}^{\text{\! + }}}{\text{ + 0}}{\text{.41}}{{\text{H}}_{\text{2}}}{\text{O}} \end{align}

or

(5b)\begin{align} &2\text{Al}_2Si_2\text{O}_5(\text{OH})_4\! + 0.57\text{Mg}^{2+}\! + 1.16\text{Fe}^{3+/2+}\! + 0.68\text{K}^{+}\! + 4.59\text{H}_2 \text{O} \nonumber\\ &\quad\to \left(\text{K}_{0.68}\text{Mg}_{0.06}\right)\left(\text{Fe}_{1.16}\text{Al}_{0.66}\text{Mg}_{0.51}\right) [\text{Al}_{0.57}\text{Si}_{3.43}]\text{O}_{10}(\text{OH})_2 \nonumber\\ &\quad+ 2.77\text{Al}(\text{OH})_3 + 0.57\text{Si}(\text{OH})_4 + 4.59\text{H}^{+} \end{align}

It has been proposed that reverse weathering was more efficient under the pervasively silica-rich conditions found in Earth’s early oceans, helping to maintain warm, ice-free conditions throughout the Precambrian period (e.g. Isson & Planavsky, 2018). This probably occurred via the establishment of a strong relationship between global climate, continental silicate weathering rates, seawater pH and authigenic clay mineral formation, according to the empirical reactions show in Equations 6 & 7:

(6)\begin{align}&{\text{Silicate weathering: Mafic minerals and glass + C}}{{\text{O}}_{\text{2}}}{\text{ + }}{{\text{H}}_{\text{2}}}{\text{O }} \nonumber\\ &\quad \to {\text{ HC}}{{\text{O}}_{\text{3}}}^{\text{--}}{\text{ + Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + cations}}\end{align}

(7)\begin{align} &{\text{Reverse weathering: HC}}{{\text{O}}_{\text{3}}}^{\text{--}}{\text{ + Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}{\text{ + cations}} \nonumber\\ &\quad \to {\text{C}}{{\text{O}}_{\text{2}}}{\text{ + phyllosilicates}}\end{align}

Banerjee et al. (2016) and Baldermann et al. (2022) have also argued that glauconite – the most abundant authigenic clay mineral in the Phanerozoic – preferentially formed during warm climate periods with sea-level high stands, such as in the late Palaeozoic, early Mesozoic and Cenozoic. However, the influence of marine authigenic clay mineral formation on the carbon cycle and seawater pH remains poorly quantified, as the reaction rates are unknown and as the amount of CO₂ released during reverse weathering is probably much lower than the cooling produced by increased silicate weathering reactions that consume CO₂. Further research should assess the kinetics of clay mineral reactions and their influence on the pH sensitivity of aquatic systems, which is required to model the impact of reverse weathering on Earth’s CO₂ inventory and climate change (e.g. Isson & Rauzi, 2024).

Carbon cycle models (e.g. Isson & Planavsky, 2018) can incorporate silica (re)cycling and alkalinity consumption through reverse weathering reactions (Equation 1a–e) to account for the formation of berthierine, chamosite, clinochlore, corrensite, glauconite, greenalite, minnesotaite, odinite, saponite, sepiolite and sudoite. These models attempt to provide quantitative estimates on long-term fluctuations in Earth’s climate versus seawater pH and pCO₂ (e.g. Isson & Planavsky, 2018). It should be noted that the rates of reverse weathering and clay mineral authigenesis in marine sediments are very often considered to be slow (e.g. Ehlert et al., 2016; Baldermann et al., 2018). In contrast, faster reverse weathering rates may occur in proximal and distal marine settings, such as in deltas, estuaries and in some deep-sea environments (e.g. Michalopoulos & Aller, 1995, 2004; Wallmann et al., 2008, 2023; Rahman et al., 2017; Steiner et al., 2022; Geilert et al., 2023). In proximal marine settings, it has been estimated that reverse weathering may sequester 10–50% of the total dissolved silica exported to the modern oceans (e.g. Tréguer & De La Rocha, 2013; Rahman et al., 2016; Tréguer et al., 2021).

The marine silica cycle: biogenic silica versus authigenic clay mineral precipitation

Reverse weathering strongly influences the global marine Si cycle via the transformation of biogenic and amorphous silica (SiO₂), reactive silicates (e.g. feldspar) or dissolved Si(OH)₄ into authigenic clay minerals (e.g. Rahman et al., 2016; Geilert et al., 2023; Wallmann et al., 2023). The individual transformation rates are believed to be relatively slow in open marine settings but may be significantly faster in estuaries, deltas and hydrothermal (deep-sea) environments. Importantly, it is the chemical composition (and temperature) of local marine porewaters that controls the rate of clay mineral authigenesis, with continental margin and shelf settings being particularly favourable. This is probably because the faster and quantitatively more important precipitation of biogenic opal by marine silicifying organisms in the upper water column of the oceans consumes dissolved Si(OH)₄, resulting in low Si(OH)₄ concentrations in open, deep marine porewaters, according to the reaction shown in Equation 8:

(8)\begin{equation}{\text{Si}}{\left( {{\text{OH}}} \right)_{{\text{4(aq)}}}} \to {\text{ Si}}{{\text{O}}_{{\text{2(s)}}}}{\text{ + 2}}{{\text{H}}_{\text{2}}}{\text{O}}\end{equation}

Throughout the Phanerozoic, the main silicifying organisms have changed, with siliceous sponges expanding in the Cambrian, followed by radiolarians in the Ordovician and diatoms in the Jurassic–Cretaceous (Fig. 5a). Today, such marine biomineralizers account for ∼70–90% of the global silica sink (e.g. Tréguer & De La Rocha, 2013; Tréguer et al., 2021). Consequently, modern seawater has a Si(OH)₄ concentration (<0.1 mM) that is at least an order of magnitude lower than that inferred for the Precambrian oceans (∼1.00–2.21 mM; cf. Fig. 5a; e.g. Siever, 1992; Conley et al., 2017). The modern ‘Si-depleted’ seawater is therefore unsaturated with respect to the conditions required to precipitate most marine clay minerals (e.g. Baldermann et al., 2018; Isson & Planavsky, 2018), contributing to the reduced number of authigenic clay mineral species forming in the modern ocean compared to the wider array documented in Precambrian marine sedimentary rocks. For example, Han et al. (2024) have proposed that widespread evaporitic conditions on the Yangtze Gorges shelf (south China) provided favourable conditions during the Ediacaran for authigenic saponite formation and the development of the Doushantuo Biota.

Isson & Planavsky (2018) have also argued that the rise of silica-biomineralizing eukaryotes may have caused a substantial reduction of the reverse weathering rate in the ancient oceans, resulting in a drop in pCO₂ and a change to a less stable climate system by reducing the pH buffer capacity of the global oceans. Such a scenario is consistent with the high diversity of authigenic clay minerals, such as greenalite (Fig. 5b), saponite (Fig. 5c), glauconite and minnesotaite, in the Precambrian geological record, but it is incompatible with the significant increase of glauconite in the marine sediments since the Phanerozoic (Fig. 5a, d). The negative correlation between the Si(OH)₄ concentration in Phanerozoic seawater and the abundance of glauconite in the rock record further suggests that seawater Si was not a limiting factor for this type of marine clay mineral authigenesis. Consequently, the advent of biomineralization and the resulting drop in the seawater Si concentration did not result in the cessation of authigenic clay mineral formation during the Phanerozoic (cf. Equation 8, reverse reaction, & Fig. 5a). Future research should explore the abundance and type of authigenic clay minerals in Precambrian versus Phanerozoic marine sedimentary rocks.

Since the emergence and the proliferation of diatoms in the Jurassic–Cretaceous, a drop in the seawater Si(OH)₄ concentration occurred (cf. Fig. 5, ∼200 Myr; e.g. Conley et al., 2017), which came along with a (time-displaced) increase in the abundance of glauconite in the marine rock record (Fig. 5a). Such a scenario appears to be consistent with the comparatively high abundance of glauconite and celadonite observed in modern coastal and oceanic sediments (e.g. Banerjee et al., 2016; Cuadros et al., 2017; López-Quirós et al., 2019, 2023; Baldermann et al., 2022; Geilert et al., 2023). There, re-mineralization of reactive biogenic opal in the sedimentary pile provides an early diagenetic (secondary) pool of dissolved Si(OH)₄, which is required for clay mineral formation (e.g. Vorhies & Gaines, 2009; Isson & Planavsky, 2018; Woltz et al., 2023). This process illustrates the linkage between biogenic opal formation and early diagenetic re-mineralization and clay mineral authigenesis, which is expressed by Equation 9:

(9)\begin{equation}\begin{gathered} {\text{Si}}{{\text{O}}_{{\text{2(s)}}}}{\text{ + Al}}{\left( {{\text{OH}}} \right)_{\text{3}}}{\text{ + }}\left( {{\text{M}}{{\text{g}}^{{\text{2\! + }}}}{\text{, N}}{{\text{a}}^{\text{\! + }}}{\text{, }}{{\text{K}}^{\text{\! + }}}{\text{, F}}{{\text{e}}^{{\text{2\! +\! /3\! + }}}}} \right){\text{ + HC}}{{\text{O}}_{\text{3}}}^{\text{--}} \hfill \\ \to {\text{Authigenic clays + C}}{{\text{O}}_{\text{2}}}{\text{ + }}{{\text{H}}_{\text{2}}}{\text{O}} \hfill \\ \end{gathered} \end{equation}

Despite the above relationships, porewater data collected from shelf, slope and deep-sea sites across modern oceans indicate that the dissolved Si(OH)₄ concentration is generally too low and not favourable for more extensive marine clay mineral authigenesis (e.g. Isson & Planavsky, 2018). This is mainly due to the high degree of bio-opal remineralization occurring in the upper water column of the oceans (∼70%), the regionally uneven distribution of bio-opal production and the rapid vertical diffusion of porewater-related Si(OH)₄ to the Si-depleted bottom seawater. All of these factors prevent the establishment of the high concentrations of Si(OH)₄ at the sediment–seawater interface that are required for authigenic clay mineral formation (e.g. Tréguer & De La Rocha, 2013). The notable exceptions are glauconite, nontronite, saponite and sepiolite-palygorskite group clay minerals with the right compositions to form in marine environments. According to hydrochemical modelling of modern seawater and porewater compositions, these require comparatively low amounts of dissolved Si(OH)₄ to precipitate (e.g. Baldermann et al., 2018; Isson & Planavsky, 2018). However, it has been frequently documented that authigenic clay minerals can precipitate directly on dissolving biogenic silica particles, which could imply that at least the reactive parts of siliceous skeletons react in situ to form phyllosilicates (e.g. Michalopoulos et al., 2000; Presti & Michalopoulos, 2008; Loucaides et al., 2010; Baldermann et al., 2013; López-Quirós et al., 2019; Geilert et al., 2023). This is presumably because the local pore fluid Si(OH)₄ concentration at this reactive interface is much higher than that measured in the bulk fluid, generating geochemical microenvironments that favour clay mineral authigenesis, as confirmed by the hydrochemical modelling of modern porewater compositions (e.g. Wallmann et al., 2023).

Consequently, it is the variation of the local dissolved Si(OH)₄ concentration in porewaters across geological time that controlled the rate and magnitude of authigenic clay mineral formation. Whereas biogenic opal mineralization may have decreased the average seawater Si concentration, the massive sedimentation of siliceous components may have favoured large local porewater or early diagenetic micromilieu Si(OH)₄ concentrations that helped to boost the authigenesis of diverse clay mineral species. In the Precambrian, the global ocean was likely to have been anoxic (Fe-rich) and rich in Si(OH)₄ (e.g. Siever, 1992; Poulton & Canfield, 2011; Conley et al., 2017; Tosca et al., 2019), favouring the deposition of, for example, extensive silica-rich banded iron formations (BIFs), as well as greenalite and other marine clay minerals (e.g. Rasmussen et al., 1998, 2013, 2017, 2021; Klein, 2005; Johnson et al., 2018; Tosca et al., 2019; Muhling & Rasmussen, 2020). The latter Fe-phyllosilicates not only occur in distinct and separate sedimentary units; they also appear as encapsulated pockets and inclusions in early diagenetic chert (Fig. 5b). This implies that authigenic clay minerals, despite their tendency to dissolve, dehydrate and/or transform during diagenesis and metamorphism, could have been more abundant in the Precambrian record than recently thought or preserved in the rock record (e.g. Morris, 1993; Klein, 2005; Rasmussen et al., 2017; Isson & Planavsky, 2018; Huang et al., 2024; Isson & Rauzi, 2024). This abundance of clay minerals formed by reverse weathering would have potentially impacted the global climate at that time.

In contrast, Dunlea et al. (2017) suggested that during the Cenozoic (∼66 Myr until today) stagnating rates of marine authigenic clay mineral formation (Fig. 5a) caused seawater Mg/Ca to rise and atmospheric CO₂ to decline over the past ∼50 Myr. This was linked to the remarkable Cenozoic global cooling that continued until the beginning of the temperature rise in the late Pleistocene. This latter relationship highlights the complex links between the global marine element cycles and Earth’s climate response – features that are evident in the diversity of authigenic clay minerals found in the geological record.

Figure 5. (a) Relationship between the evolution of marine siliceous organisms, the dissolved Si concentration in the ocean (data source: Conley et al., 2017) and the abundance of glauconite in the geological record (data source: Banerjee et al., 2016) from the early Precambrian to the present. The glauconite abundance (in %) is expressed by the logarithm of the actual number of occurrences in different intervals divided by the time duration. The sedimentation and early diagenetic re-mineralization of biogenic opal probably provided porewater Si(OH)₄ favouring clay mineral authigenesis. (b) Photomicrograph showing authigenic greenalite (Gre; marked with white arrow) embedded in a greyish chert and reddish hematite matrix from the Late Archean Kushtagi-Hungund Schist Belt (India). (c) Artificially coloured scanning electron microscopy with energy-dispersive X-ray spectroscopy mineral map displaying authigenic saponite (Sap) hosted in a blueish dolomite matrix from the Ediacaran Doushantuo Formation (China). (d) Photomicrograph showing glauconite (Glt; marked with white arrow) growing in foraminiferous ooze in Pleistocene–Holocene-aged Ivory Coast–Ghana Marginal Ridge sediments. Images (b)–(d) are provided by AB.

Stratigraphic distribution of authigenic clay minerals in the marine rock record

In this section, an overview is given of the stratigraphic and temporal distributions of authigenic clay minerals that typically form in the marine realm during early diagenesis (cf. Fig. 1). Secondary clay mineral assemblages that may precipitate in the continental critical zone and during late diagenesis are not considered, as these minerals do not represent reverse weathering products. However, we emphasize that recognizing and quantifying the frequently fine-grained products of reverse weathering reactions in the marine rock record are challenging (e.g. Hazen et al., 2013; Han et al., 2022a), given that authigenic clay minerals are commonly intermingled with clay minerals of detrital origin or have been partly overprinted by deep burial diagenetic reactions. Therefore, marine clay minerals are typically of diverse origin. However, most of the truly authigenic clay minerals found in marine sediment archives are Fe-rich (here: 12 out of the 18 clay minerals considered; Hazen et al., 2013). The reasons for the dominance of authigenic clay minerals enriched in Fe(III) and/or Fe(II) over Fe-poor or Fe-depleted clay minerals in marine sediments remain disputed, but controlling factors may include the early diagenetic Fe(III) reduction coupled to complex inorganic and biological processes and the progressive decay of organic matter. These reactions provide (partly organo-complexed) reactive Fe²⁺ and also Fe³⁺ ions to the marine porewater inventory, from which authigenic Fe-rich clay minerals tend to precipitate (see earlier).

Beidellite and chlorite-smectite

Beidellite is a rather uncommon mineral of the dioctahedral smectite group that has been documented mainly in Cretaceous sequences (Figs. 6 & 7). It is thought to have formed by the alteration of montmorillonite in the presence of dissolved Fe^3+/2+, typically under the elevated temperatures of mid-oceanic ridge sites (e.g. Post et al., 1997; Bouna et al., 2012). Orange-coloured beidellite has been found in macropores and intergranular pores of altered mafic host rocks, together with intergrowths of zeolites, goethite and greenish-brown chlorite-smectite mixed-layered clay minerals (e.g. Hu et al., 2023). Corrensite, an ordered chlorite-smectite group mineral, is locally found in hydrothermally altered marine clastic sediments and ophiolitic rocks. For example, Buatier et al. (1995) reported that authigenic corrensite formed ∼32 m below the sediment–seawater interface at 300°C ± 30°C during active hydrothermal venting at the axial rift valley of the northern Juan de Fuca Ridge.

Figure 6. Temporal distribution of authigenic clay minerals found in different sedimentary environments across geological time. The occurrence of authigenic clay minerals is provided in Table S1; however, the nature of kaolinite is disputed – it can be of detrital or authigenic origin. Note that glauconite is the most abundant authigenic clay mineral across time, whereas chamosite is particularly abundant in the Palaeozoic.

Figure 7. Stratigraphic distribution of common clay minerals found in various sedimentary environments. Kaolinite is not shown because the great majority of kaolinite occurrences are considered to be of detrital rather than marine authigenic origin. The numbers at the top represent the numbers of published occurrences of specific authigenic clay minerals provided in Table S1.

Berthierine

Berthierine, the Fe member of the serpentine mineral group, formed preferentially as oolitic grain infills (together with chamosite) during Palaeozoic and Mesozoic times (Figs. 6 & 7). In the Jurassic and Cretaceous, it formed by the alteration of volcanogenic particles in oxygen-depleted seawater (e.g. Maynard, 1986; Rudmin et al., 2022a; Kalinina et al., 2024; Roy Choudhury et al., 2024). Reactions involved the transformation of siderite, kaolinite, odinite, chamosite and glauconite precursor grains under reducing conditions (e.g. Iijima & Matsumoto, 1982; Bhattacharyya, 1983; Hornibrook & Longstaffe, 1996; Tang et al., 2017) or the alteration of transported laterite-derived materials under reducing, marine porewater conditions (e.g. Fritz & Toth, 1997; Toth & Fritz, 1997). The availability of dissolved Fe²⁺ is key to the formation of yellowish-greenish-brownish berthierine in shallow marine, estuarine and fresh to brackish settings under (sub)tropical conditions (Fig. 8a). In some cases, berthierine has been documented in near-coastal coal-forming environments (Iijima & Matsumoto, 1982) and marine hydrothermal surroundings (Taylor, 1990; Roy Choudhury et al., 2024).

Figure 8. Transmission electron microscopy and scanning electron microscopy images of common authigenic clay minerals formed during reverse weathering across geological time. (a) Fibrous berthierine (Bth) crystals (∼7 Å) infilling ooids (inset image) from the Oligocene channel ironstones of the Lisakovsk deposit in Kazakhstan (Rudmin et al., 2021). (b) Lath-like celadonite (Cel) particles occurring in interstitial and void spaces in green bole beds (inset image) within the Late Cretaceous Deccan Volcanic Province in India. This Cel formed in a non-marine environment, but it otherwise precipitated under submarine conditions (Singh et al., 2022). (c) Microcrystalline chamosite (Chm) and siderite (Sd) spar from the Middle Triassic Kremikovtsi Sedimentary Exhalative Siderite Iron Formation in Bulgaria (Damyanov & Vassileva, 2001). (d) Lath-like glauconite (Glt) infilling foraminifera tests (inset image) at the Ivory Coast basin close to the Ghana continental margin (Baldermann et al., 2013). (e) Fibrous greenalite (Gre) crystals in polygon-shaped quartz (Qz) grains from the Proterozoic-aged Brockman Iron Formation in Australia (Rasmussen et al., 2021). (f) Honeycomb-like saponite (Sap) aggregates covering the surface of altered basalt grains (Voigt et al., 2020). Coloured stars in the low-resolution images indicate the positions of regions of interest used for high-resolution imaging.

Celadonite

This mineral of the Fe- and K-rich mica group preferentially forms as vesicle fillings and as massive aggregates in altered basaltic rocks (Fig. 8b). It is common in the Proterozoic, often in association with BIFs (e.g. Savko, 2006), but it is most abundant in the well-preserved Mesozoic rock record (Figs. 6 & 7), probably because of the peak mid-oceanic ridge length and fast seafloor spreading rates at that time (e.g. Müller et al., 2013, 2022). The most favourable conditions for the formation of bluish-green celadonite developed in the Cretaceous, when this mineral was primarily formed by the in situ alteration of mafic and intermediate volcanic rocks in the presence of either marine or meteoric fluids rich in K⁺, Fe^3+/2+ and Mg²⁺ ions (e.g. Baker et al., 2012). Even though celadonite and glauconite can precipitate under similar physicochemical conditions, celadonite is rarely found in Palaeozoic successions and modern oceanic sediments and volcanic rocks (e.g. Singh et al., 2023).

Chamosite

Even though chamosite, the Fe²⁺ member of the chlorite group, has formed in marine sediments since the Precambrian, the best conditions for its crystallization prevailed during the Ordovician, Silurian, Devonian and Jurassic to Neogene (Figs. 6, 7 & 8c). Favourable conditions for the precipitation of authigenic chamosite often occurred in close association with the formation of ooidal ironstones, where the supply of dissolved Fe²⁺ and Mg²⁺ ions from detrital sources or reducing marine fluids and the leaching of Si(OH)₄ and K⁺ ions from reactive precursor clays, such as berthierine or glauconite, triggered the formation of this mineral (e.g. Mücke, 2006; Tang et al., 2017). However, the burial diagenetic or hydrothermal formation of chamosite from smectite is also possible. The stratigraphic distribution and the abundance of authigenic chamosite, a relatively common marine clay mineral, which often appears with a greyish-greenish-brown colour, commonly overlap to some degree with those of glauconite (e.g. Van Houten & Bhattacharyya, 1982; Rudmin et al., 2019, 2021). Chamosite has, however, also been reported in Proterozoic sequences, where Fe²⁺- and Si(OH)₄-rich seawater facilitated its formation by ion exchange between the sediment, porewater and seawater (e.g. Tang et al., 2017).

Glauconite

This Fe-rich mica group mineral was commonly precipitated as light to dark green rounded concretions (pellets) during the Precambrian, Mesozoic and Cenozoic (Figs. 6 & 7). The Late Cretaceous and Palaeogene marked notable periods of glauconite formation on the continental shelves of that time. Factors that facilitated the formation of glauconite (and of its precursor, Fe(III)-smectite) during greenhouse climates included the availability of dissolved Si(OH)₄, Fe^2+/3+, K⁺ and Mg²⁺ in seawater or sediment porewater, the presence of microenvironments, such as faecal pellets, bioclasts and corroded feldspar, and the development of oxygen-depleted conditions at the sediment–seawater interface (e.g. Meunier & El Albani, 2007; Charpentier et al., 2011; Baldermann et al., 2012; Banerjee et al., 2016; Zhang et al., 2022b). Glauconite formation was also favoured by global sea-level fluctuations and low sedimentation rates (e.g. Roy Choudhury et al., 2022). The early diagenetic precipitation of glauconite and other ferric illites contributed significantly to the marine sinks of Fe, Mg, K, Si and Al (e.g. Cuadros et al., 2017; Baldermann et al., 2022). In contrast, during the Holocene the authigenic glauconite grains formed almost exclusively in deep marine sediments (Fig. 8d).

Greenalite

This uncommon Fe-rich mineral of the serpentine group primarily formed in the Precambrian in close association with BIF deposits (Figs. 6, 7 & 8e). Some appearances have also been documented in the Palaeozoic and Mesozoic rock records (e.g. Rasmussen et al., 1998, 2021; Grenne & Slack, 2019; Ghosh, 2020). Greenalite formation from amorphous Si- and Fe-bearing precursor phases in the Archaean ocean (e.g. Tosca et al., 2016, 2019; Rasmussen et al., 2017, 2021) was likely favoured by the enrichment of Fe²⁺, as well as Al³⁺, Mg²⁺ and Si(OH)₄ in near-surface sediments (e.g. Johnson et al., 2018; Mohanty & Mishra, 2023). Hydrothermal sources have been suggested to have contributed to the source of Fe in greenalites (e.g. Tosca et al., 2019; Muhling & Rasmussen, 2020; Tosca & Tutolo, 2023; Tostevin & Ahmed, 2023).

Kaolinite

Kaolinite, the dioctahedral member of the kaolinite-serpentine group, occurs as a chemical weathering product in all rock types (e.g. Stoch & Sikora, 1976). Its formation and enrichment in soils and shelf sediments were favoured during the Jurassic and Cretaceous (Fig. 7). Although most of the kaolinite found in continental shelf and deep-sea sediments is detrital in origin, some occurrences have been attributed to reverse weathering (e.g. Maliva et al., 1999; Pe-Piper et al., 2005). Hydrothermal alteration also induces kaolinization by the removal of Si(OH)₄, K⁺, Mg²⁺ and Fe^2+/3+ from volcanic glasses and mafic minerals in both terrestrial and marine environments (e.g. Altschuler et al., 1963; Erkoyun & Kadіr, 2011; Kadіr et al., 2011).

Minnesotaite

This Fe-silicate mineral belonging to the pyrophyllite-talc group mainly formed in the Precambrian, but minor occurrences have also been documented in Lower Carboniferous and Permian sequences (Figs. 6 & 7). Minnesotaite deposits are often associated with BIFs, where greenish-grey to olive-green aggregates of fine needles and platelets probably formed via the alteration of greenalite under reducing conditions and in the presence of fluids with a high Mg/Fe ratio (e.g. Blake, 1965; Rasmussen et al., 1998; Lempart-Drozd et al., 2022).

Nontronite

This yellowish-green (brownish-orange when weathered) Fe³⁺ member of the smectite group formed dominantly in the Precambrian, Cambrian, Cretaceous, Palaeogene and Neogene as an alteration product of mafic minerals at mid-oceanic ridge sites or in close association with ironstone deposits, as well as in continental soils (e.g. Köhler et al., 1994; Dekov et al., 2007; Afify et al., 2015, 2018; Rudmin et al., 2022b). However, nontronite (or Fe(III)-smectite in general) may also precipitate in (semi)confined marine microenvironments at low temperatures and subsequently act as a precursor phase to glauconite formation (e.g. Baldermann et al., 2013).

Odinite

Odinite is a rare, Fe(III)-enriched clay mineral of the kaolinite group. It formed in ancient and modern shallow marine settings, lagoons, pro-deltas and estuaries characterized by warm waters enriched in dissolved Fe³⁺, Mg²⁺ and Si(OH)₄, such as in the Cretaceous (Figs. 6 & 7; e.g. Bailey, 1988). Due to its high reactivity, odinite rapidly alters to berthierine and chlorite (e.g. Hornibrook & Longstaffe, 1996; Šegvić et al., 2020).

Palygorskite

Palygorskite is a relatively uncommon mineral in the marine rock record. It formed either by the transformation of sepiolite, smectite and volcanic ash or as a precipitate from pore fluids rich in Si(OH)₄ and Al³⁺ and Mg²⁺ ions (e.g. Yalçin & Bozkaya, 2011), preferentially in the Cretaceous and Neogene (Figs. 6 & 7; e.g. Akbulut & Kadir, 2003; Galán & Pozo, 2011). Microstructural and mineralogical evidence suggests that the purest palygorskite deposits of Cretaceous to Palaeogene age were formed in situ by reverse weathering on the seafloor (e.g. Thiry & Pletsch, 2011). Palygorskite forms a complete solid solution with sepiolite and primarily occurs in lacustrine, fluvial, (peri)marine and soil environments (e.g. Isphording, 1973; Galán & Pozo, 2011).

Saponite

The trioctahedral (Mg-rich) smectite group mineral saponite mainly formed during the Cenozoic, although rare occurrences of this mineral have been reported in the Precambrian, Devonian and Permian (Figs. 6 & 7). Precambrian saponite preferentially formed in restricted, evaporitic lagoons that were temporally connected to the ocean (e.g. Han et al., 2022b). Bluish-green (reddish-brown when weathered) saponite is also known to form via the alteration of mafic rocks or volcanic matter under the influence of hydrothermal fluids in marine settings (Fig. 8f), as well as in evaporitic and alkaline lakes (e.g. Andrews, 1980; Post, 1984; Akbulut & Kadir, 2003; Rudmin et al., 2023a).

Sepiolite

This hydrous Mg-rich silicate mineral is found in abundance in Cenozoic marine sediments (together with palygorskite), but minor occurrences have also been reported from Permian and Cretaceous strata (Figs. 6 & 7). Fluids rich in Mg²⁺ and Si(OH)₄ facilitated the formation of sepiolite in (peri)marine, transitional and lacustrine environments (e.g. Akbulut & Kadir, 2003; Yalçin & Bozkaya, 2011; Sánchez-Roa et al., 2016). Sepiolite may also form via the hydrothermal alteration of basalts and volcanic ash (e.g. Akbulut & Kadir, 2003; Galán & Pozo, 2011; Pozo & Calvo, 2018).

Stevensite

Stevensite is a trioctahedral (Mg-rich) smectite group mineral that often formed by precipitation in alkaline and playa lakes (e.g. de Oliveira Nardi Leite et al., 2020; Netto et al., 2022). This mineral mainly formed from the Cretaceous to recent (Figs. 6 & 7) within ooids, peloids and intraclasts in Mg²⁺- and Si(OH)₄-rich lacustrine and coastal marine settings, and more rarely by the alteration of oceanic crust under the influence of hydrothermal fluids (e.g. Banfield et al., 1991; Cuevas et al., 2003; Armelenti et al., 2016; Lima & De Ros, 2019).

Stilpnomelane

This blackish-greenish-yellowish smectite group mineral formed mainly during the Precambrian and Mesozoic (Figs. 6 & 7) within marine Fe-rich deposits, such as BIFs, that received upwelling of Fe²⁺- and Si(OH)₄-rich deep-seated (hydrothermal) fluids mixed with shallower and slightly oxygenated seawater (e.g. Garzanti et al., 1989; Pickard, 2003; Wang et al., 2015).

Talc

This Mg member of the talc-pyrophyllite group is a relatively rare authigenic mineral in the sedimentary rock record. It mostly formed in Precambrian, Permian and Cretaceous strata (Figs. 6 & 7) under the influence of Al-poor and Si(OH)₄- and Mg-rich fluids under marine (or lacustrine) conditions (e.g. Noack et al., 1989; Chen et al., 2002; Tosca et al., 2011; Han et al., 2024). Rarely, talc also formed by the alteration of sepiolite under the influence of microbial activity and in the presence of Si(OH)₄-rich marine pore fluids.

Vermiculite

Dioctahedral vermiculite, an expandable and abundant clay mineral in continental soils, typically formed by the chemical weathering or the alteration of amphibole, biotite and phlogopite precursors (e.g. Wilson, 2004). However, vermiculite is a rather uncommon mineral in the marine rock record, with some rare occurrences being documented in the Cretaceous and Pleistocene to recent times (Figs. 6 & 7). In these rocks, vermiculite appears in its trioctahedral form, typically within interstratified illite-vermiculite and chlorite-vermiculite (e.g. Craw et al., 1995; Gutierrez-Mas et al., 1997; Yin et al., 2018).

Reverse weathering and clay mineral authigenesis across geological time

Earth’s greenhouse climatic periods clearly influenced the rates of reverse weathering and the associated distribution and abundance of authigenic clay minerals in marine sediments across geological time. Such periods were characterized by an elevated seawater temperature (on the shelf), a higher continental (dissolved) element export to the oceans, comparatively higher Si(OH)₄ concentrations in marine porewaters and enhanced microbial activity that catalysed reverse weathering reactions (e.g. Schieber et al., 2007; Aubineau et al., 2019, 2020). In this section, some examples are provided of specific periods that favoured clay mineral authigenesis in the oceans and along the submerged continental margins.

Precambrian palaeoclimate

The authigenesis of a suite of Fe(II/III)-phyllosilicates (cf. Fig. 7) in the Precambrian ocean is believed to have released significant amounts of CO₂, which helped maintain the warm and ice-free global climate despite the much lower solar luminosity at that time (e.g. Isson & Planavsky, 2018; Han et al., 2024). Glauconite formed dominantly on marine shelves during the Late to Middle Proterozoic, resulting in thick glauconitic sandstones and glauconitic mudrock deposits in many places worldwide (e.g. Banerjee et al., 2008, 2016; Mandal et al., 2020, 2022). The formation of sedimentary greenalite in the Palaeoproterozoic and of authigenic talc in the Neoproterozoic was facilitated by high Fe²⁺, Al³⁺, Si(OH)₄ and Mg²⁺ concentrations in seawater (e.g. Tosca et al., 2011, 2016; Mohanty & Mishra, 2023; Tostevin & Ahmed, 2023). Moreover, authigenic saponite formed in restricted evaporitic lagoons over large areas of the Yangtze Gorges shelf in south China from Mg²⁺- and Si(OH)₄-rich Ediacaran seawater (e.g. Han et al., 2022b, 2024). A combination of palaeo-environmental factors, such as the widespread presence of suboxic seawater masses, the absence of marine silicifying organisms, adequate seawater chemistry and the overall low sedimentation rates (e.g. Eriksson et al., 2001; Banerjee et al., 2008), favoured the authigenesis of Fe-silicates in the Precambrian (e.g. Banerjee et al., 2008; Mandal et al., 2020). Reverse weathering was also facilitated by microbial activity, as is evident from the Precambrian-aged microbial mat deposits enriched in Fe-illite, ferroan illite-smectite, stilpnomelane and minnesotaite (e.g. Konhauser & Urrutia, 1999; Schieber et al., 2007; Rasmussen et al., 2013; Konhauser et al., 2017; Aubineau et al., 2020). These microbial biofilms trapped K⁺ ions from the seawater and released them into the porewater during the degradation of organic matter at early diagenesis, but they also provided the porewater with other important ions, such as Fe^2+/3+, during respiration. This resulted in the progressive illitization of smectite precursors, as well as in the neoformation of K- and Fe-phyllosilicates, which subsequently acted as mineral archives that allow us to identify biomineralization processes on ancient Earth (e.g. Vorhies & Gaines, 2009; Aubineau et al., 2019).

Late Palaeozoic to early Mesozoic warming

The Ordovician to Devonian timespan, as well as the Jurassic, witnessed the formation of authigenic chamosite and berthierine in massive ooidal ironstone facies (cf. Fig. 7). Mid-Ordovician-aged chamositic ironstones formed on storm-dominated shelves characterized by Fe²⁺-rich sea-bottom conditions (e.g. Dunn et al., 2021). The authigenesis of these Fe(II)-rich clay minerals that occurred during such periods was linked to a sea-level highstand, the presence of oxygen-depleted but cation-rich seawater, microbial mediation and a warm climate (e.g. Van Houten & Arthur, 1989; Young, 1989; Li et al., 2021b).

Mesozoic greenhouse

Glauconite authigenesis was significant during transgressive events in the Cretaceous (cf. Fig. 7), especially from the Aptian to the Maastrichtian (e.g. Bansal et al., 2020, 2021). The higher temperature and suboxic seawater enriched in K⁺, Mg²⁺, Si(OH)₄ and Fe²⁺ ions, coupled with a sea-level highstand during the Cenomanian greenhouse conditions, facilitated the high abundance of glauconite on the marine shelves (e.g. Haq, 2014; Banerjee et al., 2016; Baioumy et al., 2020; Wilmsen & Bansal, 2021; Scheiblhofer et al., 2022). The formation of glauconite probably sequestered significant amounts of K, Fe, Mg, Si and Al from Cretaceous seawater (e.g. Baldermann et al., 2022). Upper Cretaceous glauconite deposits mainly formed along the margins of the Tethys Ocean (e.g. Bansal et al., 2020). In addition, a high abundance of other authigenic clay minerals has been documented for the Cretaceous, such as berthierine, celadonite, palygorskite, talc, stevensite and beidellite (cf. Fig. 7), which was possibly related to a fast sea-level rise, warm seawater and high seafloor spreading rates at mid-oceanic ridges at that time (e.g. Jeans, 2006; Müller et al., 2013, 2022; Singh et al., 2023). In addition, chamositic-siderite ironstone deposits formed during Oceanic Anoxic Event 3, and abundant Jurassic chamositic-ooidal ironstones precipitated at the Neo-Tethyan margin due to upwelling and fluctuating redox conditions against the backdrop of Mesozoic greenhouse conditions (e.g. Van Houten & Arthur, 1989; Han et al., 2023).

Cenozoic greenhouse climate

The Cenozoic era was marked by several hyperthermal events, which were associated with abundant authigenic Fe(II/III)-phyllosilicates preserved in the marine sedimentary record (cf. Fig. 7). Glauconitization was frequent due to a high sea level and warm greenhouse conditions since at least the Cretaceous (e.g. Baldermann et al., 2022), with maximum occurrences documented during hyperthermal events in the Palaeogene, Eocene and Oligocene (e.g. Banerjee et al., 2020). A multi-proxy study by Roy Choudhury et al. (2022), which applied biostratigraphy, carbon isotopes, Sr concentrations and ⁴⁰Ar/³⁹Ar chronostratigraphic data, demarcated the Palaeogene hyperthermal events as being remarkably rich in glauconite deposits, such as those reported for the Palaeocene–Eocene Thermal Maximum. The combination of oceanic hypoxia, warm climate conditions and rapid sea-level changes during the Palaeogene temperature maxima favoured the formation of glauconite deposits (e.g. Banerjee et al., 2020; Roy Choudhury et al., 2022). Before that time, around the Cretaceous–Paleogene boundary, ooidal ironstones were deposited, which contain abundant chamosite, berthierine and glauconite (e.g. Rudmin et al., 2017).

Modern-day geochemical fluxes associated with reverse weathering

Clay mineral reactions, such as smectite-illitization, kaolinization and chloritization, which occur during late burial diagenesis, are often considered to be slow (e.g. Cuadros et al., 2017). In contrast, the formation of authigenic clay minerals in modern marine sediments by reverse weathering reactions is increasingly recognized as a faster process and can therefore influence modern and past ocean chemistry and related element (re)cycling at the sediment–seawater interface (e.g. Baldermann et al., 2022).

The first indications for the substantial removal of Na⁺, Mg²⁺ and K⁺ ions and Si(OH)₄ from present-day seawater by clay mineral authigenesis were provided by Mackenzie & Garrels (1966a). They concluded that element sequestration by reverse weathering balances the riverine influx of the dissolved components from the continents. Since then, extensive field, modelling and laboratory studies have been published that strengthen the evidence for reverse weathering as a key process impacting modern marine element cycles (e.g. Michalopoulos & Aller, 1995; Michalopoulos et al., 2000; Wallmann et al., 2008, 2023; Baldermann et al., 2015; 2022; Ehlert et al., 2016; Cuadros et al., 2017; Geilert et al., 2023). Although the reverse weathering fluxes of individual elements are relatively small (<10%) compared to the global marine carbon and silicon budgets, clay mineral authigenesis can sequester various chemical elements, such as Na, K, Li, Mg, Fe, Al and Si, implying that the varying rates of clay mineral formation during marine diagenesis can impact global element budgets.

Isson & Planavsky (2018) provided quantitative estimates of the significance of reverse weathering on the modern global silica cycle. They calculated that the oceanic Si output fluxes related to chert formation and reverse weathering reactions account for 10.4 and 0.5 Tmol year^–1, respectively. The latter flux is equivalent to 4.6% of the total continental Si input to the oceans by combined riverine, hydrothermal, aeolian and groundwater sources (10.9 Tmol year^–1). However, Tréguer et al. (2021) proposed that the Si output flux due to reverse weathering is much higher than previously calculated, at ∼5 Tmol year^–1. This 10-fold difference in the calculated Si output flux reflects the problems associated with model parameterization of reverse weathering reactions occurring in the modern oceans.

More recently, Wallmann et al. (2023) reported on the modern oceanic element fluxes associated with marine silicate weathering and authigenic clay mineral formation. They calculated net fluxes across the sediment–seawater interface in the order of +1.5 Tmol Na year^–1, –2.5 Tmol K year^–1, –2.0 Tmol Mg year^–1, +2.5 Tmol Ca year^–1 and +1.9 Tmol Si year^–1, where the positive numbers indicate cation release into seawater mainly due to feldspar (e.g. plagioclase) dissolution and the negative numbers denote cation uptake by the precipitation of unspecified authigenic clay minerals and carbonates. However, these estimates did not consider Si burial through biogenic opal precipitation.

More specifically, it has been reported that element sequestration associated with glauconite formation in modern shallow marine settings (0–200 m water depth) sequesters ∼0.06 Tmol year^–1 of K, Mg and Al, ∼0.1 Tmol year^–1 of Fe and ∼0.3 Tmol year^–1 of Si, respectively (see Fig. 4 in Baldermann et al., 2022). This is equivalent to ∼3–6% of the total oceanic K inventory that is delivered from riverine and hydrothermal sources and thus of the same order of magnitude as the K burial flux attributed to authigenic Fe-illite formation in mangrove forests (up to 0.1 Tmol year^–1; Cuadros et al., 2017). Although these estimates have a high degree of uncertainty, Li et al. (2022a) proposed that glauconite is an important sink of seawater K (∼0.2 Tmol year^–1) based on K isotope measurements of modern marine ‘bulk’ sediments. Regarding the marine Mg cycle, glauconite formation at the marine shelf has been estimated to constitute ∼10–38% of the Mg consumption by authigenic clays forming in the Amazon deltaic sediments (∼0.21 Tmol year^–1) and ∼1–3% of the Mg sink that is associated with fresh oceanic crust alteration (Sun et al., 2016; Huang et al., 2018). Shallow-water glauconite formation may be even more important for the marine Al and Si cycles. It involves ∼27–82% removal of the dissolved riverine Al influx entering the modern oceans (∼0.1 Tmol year^–1; Maring & Duce, 1987) and ∼10% removal of the total Si uptake by authigenic clay minerals on a global scale (∼4.7 ± 2.3 Tmol year^–1; Rahman et al., 2016; Tréguer et al., 2021). In addition, glauconite formation acts as a significant sink for Fe, sequestering ∼8–33% of the dissolved and particulate riverine fluxes of Fe that enter the ocean (∼0.48–0.86 and ∼0.63 Tmol year^–1, respectively; Raiswell, 2006).

In contrast, the elemental uptake rates by glauconite formation in modern deep-marine sediments are currently poorly constrained, but Fe sequestration at Ocean Drilling Project (ODP) Site 959 (Ivory Coast–Ghana Marginal Ridge) has been estimated to be ∼80 µmol cm^–2 kyr^–1. Although this oceanic sink of Fe is two orders of magnitude lower than that of modern shelf areas, it is (on average) six times higher than that related to the formation of pyrite in suboxic subsurface sediments 5 m below the seafloor or shallower (Baldermann et al., 2015). This indicates that glauconite formation in the deep sea may also act as an important Fe sink, in addition to marine Fe-sulfide minerals, which develop under euxinic to anoxic conditions deeper within the sedimentary pile (e.g. Raiswell & Canfield, 2012). As for the modern deep-sea Al, Si, K and Mg cycles, however, glauconite precipitation seems to be less important, which has been attributed to the slower reaction rates at the lower temperatures (<5°C) of deep-sea settings (e.g. Baldermann et al., 2013, 2015, 2022). However, Logvinenko (1982) has reported that modern deep-sea sediments contain ∼1–2 wt.% glauconite or glauconite-smectite on average. In modern calcareous and fossiliferous deep-sea sediments, the glauconite content is even higher, reaching 2–3 wt.% on average, according to Giresse et al. (2004) and Wigley & Compton (2007). These relationships may suggest that the elemental uptake rates of Al, Si, K and Mg by glauconite formation in the deep ocean are currently underestimated.

In hydrothermally influenced sediments, such as at the Galapagos spreading centre, the Mid-Atlantic Ridge, the Crozet Island Archipelago, the Costa Rica Margin, the Pacific–Antarctic Ridge and the Chile Ridge, precipitation of Fe-bearing smectites occurs in the form of nontronite, Fe(III)-montmorillonite and saponite. The formation of these minerals can locally take up ∼1000 µmol Fe cm^–2 kyr^–1, suggesting that smectite authigenesis at thermally active sites of the deep sea can also constitute an important (but locally restricted) Fe sink in the modern oceans (e.g. Chester, 2000). Additional reverse weathering reactions occur via the low- and high-temperature interaction of fresh oceanic crust and seawater in the vicinity of mid-ocean ridge sites (e.g. Elderfield & Schultz, 1996; Higgins & Schrag, 2015; Voigt et al., 2020). These have been proposed to contribute strongly to the elemental and isotopic composition of seawater, as indicated by in situ Rb-Sr dating of celadonite forming within veins and amygdales ∼20 Myr after ocean crust accretion (Laureijs et al., 2021a). That study also concluded that widespread celadonite formation occurring in ocean floor sediments and ophiolite lavas could act as an important sink for seawater K (but also Al, Si, Li, B, Mg and Fe), but without providing robust mass balance constraints to support their hypothesis. This could not be achieved due to the incomplete sample recovery associated with ocean-floor drilling and related challenges in quantifying the heterogeneous distribution of K-rich phases in the upper oceanic crust (Laureijs et al., 2021b).

While the influence of reverse weathering on major element cycles is becoming increasingly recognized, especially with regard to often easily extractable (millimetre-sized) and quantifiable glauconite and celadonite grains, there are few constraints on the impacts of reverse weathering on the cycling of minor and trace elements in the ocean (e.g. Tostevin & Ahmed, 2023). This is largely due to the difficulty of separating and purifying typically micrometre-scale authigenic clay minerals, which are usually closely intergrown with detrital clay minerals of similar size. It has been proposed, based on porewater analyses, that the modern marine rare earth element (REE) cycle is largely controlled by the balance of REE release from detrital clay mineral dissolution versus uptake of REEs by authigenic clay minerals, Fe/Mn-(hydr)oxides and phosphates (e.g. Abbott et al., 2019). Recent analyses of authigenic glauconite separates from the West African continental margin and comparison to the detrital clay mineral-rich muds in which they form support this model (Bayon et al., 2023), whereby detrital clay mineral dissolution coupled with REE uptake by marine authigenic clay minerals is responsible for the large sedimentary flux of REEs in the modern oceans (e.g. Abbott et al., 2015, 2019; Haley et al., 2017). In addition, the work of Bayon et al. (2023) suggests that glauconite formation from a detrital clay mineral precursor assemblage is associated with net sinks of Mo, Cr, V and Rb but represents a source of U, Li, Co, Ni, Hf, Zr, Th and Ba. Similarly, authigenic smectites in a Cenozoic sequence from the Ross Sea (Antarctica) have been reported to be enriched in V, Cr, Co, Ni and Sc compared to co-occurring detrital clay minerals (Setti et al., 2004). Other recent work suggests that authigenic clay mineral formation may act as a Be sink (Bernhardt et al., 2020). In addition, clay mineral synthesis experiments suggest that greenalite may have been an important sink of Cu, V and Zn in the Precambrian oceans (Tostevin & Ahmed, 2023).

Isotope proxies used to trace clay mineral reactions

The development of multi-collector inductively coupled plasma mass spectrometry (MC-ICP-MS) and thermal ionization mass spectrometry (TIMS) marked a significant step towards the improved measurement of the isotopic composition of minerals. These instruments are used to accurately measure elements and their isotopes down to ppm and ppb concentrations with light to large atomic masses (Li, B versus Ca, Fe, Mg, Sr, Zn, etc.) in a suite of components, such as carbonates, phosphates and silicates. These analytical techniques allowed the establishment of a range of new tools and proxy signals to identify and quantify (bio)geochemical processes occurring in Earth’s (surface) settings, such as chemical weathering reactions, element cycles and fluid–mineral interactions (e.g. Pogge von Strandmann et al., 2008; Oelkers et al., 2015; Schott et al., 2016; Andrews et al., 2020; Chanda et al., 2023; Farkaš et al., 2025). While the elemental and isotopic compositions of carbonates have long been used to reconstruct conditions prevailing at the time of their formation (i.e. palaeo-pH, temperature, redox conditions and growth rate), and thus seawater chemistry over time (e.g. Eiler, 2007), applying such methods to determine palaeo-environmental information from authigenic and detrital clay minerals is less well established. This is, among other factors, due to: (1) the lack of well-constrained fractionation factors for clay mineral–fluid systems and substantial isotopic differences between adsorbed/exchangeable and structurally incorporated constituents (e.g. Köster et al., 2019; Li et al., 2022a); (2) the difficulties in synthesizing well-crystallized and close to pure clay minerals at ambient temperature (e.g. Kloprogge, 1998; Baldermann et al., 2024); and (3) the challenges associated with isolating authigenic clay minerals from phyllosilicates of detrital origin (e.g. Zhang et al., 2010; Tosca & Masterson, 2014; Han et al., 2022a, 2024; Bayon et al., 2023).

Nevertheless, field and experimental studies using the isotopic composition of clay minerals to reconstruct marine element cycles and weathering/alteration pathways occurring in various sedimentary environments are currently on the increase. This is because marine authigenic clay minerals (1) are probably more resistant to post-depositional alteration compared to the more reactive carbonates, (2) incorporate more elements of interest (e.g. K, Si and Li) and in greater abundance than carbonates (which may also be affected by as-yet unknown vital effects), allowing for the development of more new proxies (Li et al., 2022a, 2022b), (3) represent important sinks for many elements, meaning that the isotopic characterization of authigenic clay minerals can potentially better constrain modern and past marine elemental budgets via isotope mass balance approaches (e.g. Zheng et al., 2022), and (4) are datable through conventional K-Ar and ⁴⁰Ar-³⁹Ar analysis or advanced in situ Rb-Sr analysis by laser ablation collision cell inductively coupled plasma mass spectrometry (LA-ICP-MS/MS; e.g. Kelley, 2002; Clauer et al., 2012; Hogmalm et al., 2017; Scheiblhofer et al., 2022). This is particularly the case for glauconite, for which new isotope proxies are actively being developed (Fig. 9), taking advantage of the relative ease with which typically millimetre-sized glauconite pellets can be separated from their host sediments and purified using standard sieving, acid leaching and magnetic separation techniques or can be directly targeted for in situ analyses.

Figure 9. Illustration showing potential environmental geochemical and isotopic proxies recorded in glauconite and surrounding sediments. Red text indicates primary proxies of high significance; blue text indicates secondary proxies of lower significance (i.e. these are often influenced by other environmental controls and therefore difficult to interpret). Ap = apatite; Dol = dolomite; Ilt = illite.

Clay minerals, as hydrous phyllosilicates, contain water (i.e. δ¹⁸O and δD isotopic systems) that can be (1) adsorbed on the external surface sites, (2) weakly bound in the interlayer sheet in the case of swellable clay minerals and (3) more strongly bound as hydroxyls in the octahedra. Whereas the first two types of water can isotopically exchange with ambient fluids within minutes to days, structurally bound water frequently shows slow isotope exchange rates depending on temperature and the type of clay mineral (e.g. Clauer et al., 2014). Importantly, the re-equilibration of H isotopes takes place faster at lower temperatures than that of O isotopes, so that the δ¹⁸O and δD signatures of clay minerals can provide insights into the nature of mineral reactions (e.g. Schön et al., 2016). For example, the temperature of smectite formation, its post-depositional (re)equilibration/transformation into illite and the kinetics of isotopic exchange between glauconite, host sediment and pore fluid can be resolved (e.g. Savin & Epstein, 1970).

Stable silicon isotopes (δ³⁰Si) offer a sensitive proxy for studying dissolution–precipitation reactions of silicate minerals in marine sediments (e.g. Rahman et al., 2017). Light Si isotopes (²⁸Si) are preferentially taken up over heavy Si isotopes (³⁰Si) from the pore fluid during precipitation of authigenic clay minerals, shifting pore fluid δ³⁰Si to higher values, often exceeding seawater values (e.g. Ehlert et al., 2016; Geilert et al., 2023). Elevated pore fluid δ³⁰Si is thus commonly inferred to reflect authigenic clay mineral formation. However, present constraints on marine authigenic clay mineral δ³⁰Si derive almost entirely from sequential leaching approaches (e.g. Pickering et al., 2020; Huang et al., 2023), with inherent ambiguity related to the potential dissolution of non-target phases, such as detrital clay minerals. Alternative estimates of authigenic clay mineral δ³⁰Si based on isotope mass balance are helpful (e.g. Rahman & Trower, 2023) but are impacted by uncertainties in the present-day marine Si cycle. More direct constraints from in situ laser ablation analysis of authigenic clay minerals (e.g. Geilert et al., 2023) or conventional solution-based analyses of authigenic clay mineral separates (e.g. glauconite pellets; Geilert et al., 2024) have the potential to substantially reduce these uncertainties and therefore greatly improve quantification of benthic Si fluxes, alkalinity and cation budgets related to marine silicate alteration (e.g. Wallmann et al., 2008; 2023; Geilert et al., 2024; Farkaš et al., 2025). For example, the δ³⁰Si signatures of a bulk sediment rich in diatoms, its porewater, the terrigenous fraction and the authigenic Fe-smectite or glauconite-smectite, obtained by laser ablation, were recently used to demonstrate rapid marine silicate alteration and related clay mineral formation during a coastal El Niño extreme weather event at the Peruvian margin (Geilert et al., 2023). This case study illustrated the fact that clay mineral authigenesis can impact the marine element cycle and the CO₂ cycle on timescales of weeks to months.

The fractionation behaviour of boron isotopes (δ¹¹B) of clay minerals depends on temperature, B speciation, fluid δ¹¹B composition, pH, fluid/rock ratio and (re)equilibration time between the solid and fluid phase (e.g. Williams et al., 2001; Deyhle & Kopf, 2005). Boron isotopes have been used for reconstructing the fluid source(s) involved in the formation of ash-derived bentonite deposits in marine and terrestrial low-temperature environments (e.g. Köster et al., 2019; Clauer et al., 2022). Moreover, B concentrations and δ¹¹B signatures of diagenetically modified and rejuvenated glauconites can record either ancient seawater δ¹¹B composition or incorporation of organic-derived light ¹⁰B into glauconite during diagenesis, which enables reconstruction of the evolutionary history of glauconite (Środoń et al., 2023).

Another isotope system that has seen great interest in recent years is the stable potassium system (δ⁴¹K; e.g. Farkaš et al., 2025). The balance between silicate weathering and marine clay mineral authigenesis acts as an important control on Earth’s long-term carbon cycle. Since these processes also represent important influences on the marine K cycle (e.g. Li et al., 2022a) and are thought to be associated with substantial K isotopic fractionations (e.g. Li et al., 2019b, 2021c; Teng et al., 2020), the δ⁴¹K signature of seawater has the potential to become an important new tool for understanding long-term changes to the carbon cycle (e.g. Li et al., 2019c; Hu et al., 2020). Preferential uptake of the lighter K isotope (³⁹K) during authigenic clay mineral formation or during adsorption processes by clay minerals in the marine sediment is likely to be a key factor explaining the surprisingly heavy isotopic composition of present-day seawater (∼0.12‰ ± 0.07‰), considering the K isotope difference between modern seawater and Bulk Silicate Earth of ∼0.6‰ (e.g. Li et al., 2022a). The large K isotopic difference between seawater and Earth’s silicate crust (i.e. upper continental crust; –0.44‰ ± 0.05‰; Huang et al., 2020) remains enigmatic because modern seawater shows a significant excess of the heavy ⁴¹K isotope compared to the main input fluxes of K into the ocean. We note, however, that no direct measurements of modern or recent authigenic clay mineral δ⁴¹K have been published to date (only from bulk sediment), which is due to the difficulties in isolating these authigenic clay minerals from the host sediment. However, pellet-forming authigenic clay minerals, such as glauconite and celadonite (e.g. Santiago Ramos et al., 2020), are promising targets and have the potential to become an important archive for the reconstruction of seawater δ⁴¹K changes over time. Alternative sedimentary archives, which may also record seawater δ⁴¹K, include biogenic carbonates and K-rich evaporite salts, but these are either convoluted by vital effects and low K contents or show a very patchy temporal distribution (e.g. Li et al., 2022a). Direct measurement of authigenic clay mineral δ⁴¹K is thus a key research priority that can be used to assess the relative impact on seawater δ⁴¹K caused by changes in continental weathering intensity (e.g. Hu et al., 2020; Li et al., 2022a), the rates of low-temperature seafloor alteration and the magnitude of authigenic clay mineral formation in marine sediments (e.g. Wang et al., 2021).

We have already mentioned the important role of reverse weathering in the marine REE budget, with the sedimentary REE fluxes and the related composition of the marine bottom waters being controlled by the dissolution of detrital clay minerals versus the uptake of REEs by authigenic clay minerals (e.g. Haley et al., 2017; Abbott et al., 2019; Bayon et al., 2023). This has important implications for the REE concentration and the Nd isotopic composition of water masses, which have been shown to evolve over distance, location and time (e.g. Abbott et al., 2019). It has been proposed that authigenic clay minerals, such as glauconite, may be promising recorders of bottom water or porewater Nd isotopic compositions (e.g. Giresse et al., 2021), with potential applications for reconstructing changes to ocean circulation over time. Similarly, radiogenic Sr isotope (⁸⁷Sr/⁸⁶Sr) signatures have been used to track the evolution of authigenic clay minerals, showing that ⁸⁷Sr/⁸⁶Sr of detrital clay minerals is progressively lost as they are replaced by authigenic clay minerals, causing ⁸⁷Sr/⁸⁶Sr to approach a composition that is more similar to that of seawater (e.g. Clauer et al., 1992). This observation is also commonly used to justify anchoring of Rb-Sr isochrons using expected seawater Sr isotope values (e.g. Redaa et al., 2023).

Other promising isotope systems that have only rarely been applied to marine authigenic clay minerals include (1) magnesium (δ²⁶Mg) and calcium (δ⁴⁴Ca) for obtaining insights into biogeochemically coupled processes, weathering rates and marine element budgets (e.g. Dunlea et al., 2017; Huang et al., 2024), (2) lithium (δ⁷Li) to quantify continental palaeo-weathering regimes versus rates of oceanic crust alteration (e.g. Zhang et al., 2022a; Yin et al., 2023) and (3) zinc (δ⁶⁶Zn), iron (δ⁵⁶Fe) and chromium (δ⁵³Cr) to constrain palaeo-redox conditions during clay mineral precipitation (cf. Fig. 9).

Another (re)emerging application for authigenic clay minerals that is likely to gain importance aims to determine the absolute depositional ages of marine sedimentary rocks. The most widely used radiometric techniques for obtaining absolute depositional ages are presently U-Pb zircon dating of volcanic ash (e.g. Linnemann et al., 2018), ⁴⁰Ar/³⁹Ar dating of lava flows (e.g. Jourdan et al., 2005) and Re-Os dating of organic-rich black shales (e.g. Rooney et al., 2015; Millikin et al., 2022). Because these target lithologies are often restricted to specific tectonic or environmental settings, dating sedimentary sequences remains among the most challenging tasks in geochronology. Glauconite geochronology is a potential alternative approach where ash beds or black shales are absent. In fact, ∼40% of all ages used to constrain the geological timescale of the last 250 Myr up until the mid-1990s were obtained from glauconite (Smith et al., 1998). Although glauconite bulk separates have long been dated using the Rb-Sr, K-Ca, K-Ar and ⁴⁰Ar-³⁹Ar techniques, frequent mismatches with independent chronostratigraphic constraints (e.g. Selby, 2009; Cecil & Ducea, 2011) mean that glauconite geochronology is no longer widely employed. Erroneous ages have been attributed to factors including the presence of detrital Rb- and K-bearing phases in glauconite separates, the presence of younger and isotopically distinct clay mineral overgrowths, as well as post-depositional disruption or resetting of the Rb-Sr isotope system by exposure to increased temperatures and/or fluid–rock interactions (e.g. Rafiei et al., 2023). Various sample preparation and screening approaches, based on improved mineral separation and more aggressive chemical cleaning, have been developed to address these problems, but with varying degrees of success (e.g. Derkowski et al., 2009).

The advent of in situ Rb-Sr and K-Ca dating offers a promising new approach that has the potential to resolve many of these limitations, leading to a recent revival in authigenic clay mineral-based geochronology (e.g. Zack & Hogmalm, 2016; Hogmalm et al., 2017; Laureijs et al., 2021a; Redaa et al., 2023). Unlike traditional solution-based Rb-Sr dating, which requires the dissolution of bulk sample powders followed by wet chemical separation of Rb and Sr, reaction cell mass spectrometry resolves the spectral overlap of, for example, ⁸⁷Rb and ⁸⁷Sr via the addition of a reaction gas (e.g. N₂O) to a reaction cell (e.g. Hogmalm et al., 2017; Gorojovsky & Alard, 2020; Redaa et al., 2023). In the case of the Rb-Sr system, Rb is measured based on mass while Sr isotopes are measured as mass-shifted oxides (i.e. produced by reaction with the gas), allowing quantitative online separation of ⁸⁷Rb and ⁸⁷Sr. When combined with recent advances in sedimentary petrography (Rafiei et al., 2020; Han et al., 2022a) this approach permits rapid in situ Rb-Sr dating of carefully screened glauconite or celadonite grains, which opens up a range of new applications (e.g. Laureijs et al., 2021a; Scheiblhofer et al., 2022; Rafiei et al., 2023; Redaa et al., 2023). For Precambrian mudrocks rich in authigenic clay minerals, in situ dating of marine authigenic illite has also shown great promise (e.g. Subarkah et al., 2022, 2024).

Clay minerals in geoengineering: porewater alkalinity enhancement

A major challenge facing today’s society is to limit anthropogenically induced global warming to 1.5°C above the pre-industrial value (e.g. United Nations Framework Convention on Climate Change, 2015). As a result, various strategies for CO₂ removal at an industrial scale have been proposed over the past three to four decades, such as carbon capture and storage or carbon capture and utilization methods.

A novel and sustainable solution to decrease atmospheric CO₂ concentrations is to bind carbon in near-surface marine sediments and directly in the oceans (e.g. Morrow et al., 2020; Meier et al., 2022) by enhancing the chemical weathering rates of silicates and the resulting alkalinity of marine porewaters (e.g. Bellamy et al., 2012; Bach et al., 2019; Fakhraee et al., 2023; Hartmann et al., 2023; Iglesias-Rodríguez et al., 2023; Dale et al., 2024; Marín-Samper et al., 2024). As described earlier, silicate mineral weathering achieves this by releasing cations and binding CO₂ (i.e. HCO₃^– and CO₂^2– ions), thus facilitating the uptake of atmospheric CO₂ by the ocean (e.g. Feng et al., 2017). The spreading of crushed mafic rocks, such as basalt, and mafic minerals (here: forsterite, Mg₂SiO₄) in marine settings is increasingly considered in geoengineering to stabilize Earth’s climate through long-term binding and storage of atmospheric CO₂ following Equation 10 (e.g. Wood & Kleppa, 1981; Taylor et al., 2016):

(10)\begin{equation}{\text{M}}{{\text{g}}_{\text{2}}}{\text{Si}}{{\text{O}}_{\text{4}}}{\text{ + 4C}}{{\text{O}}_{\text{2}}}{\text{ + 4}}{{\text{H}}_{\text{2}}}{\text{O}} \to {\text{2M}}{{\text{g}}^{{\text{2 + }}}}{\text{ + 4HC}}{{\text{O}}_{\text{3}}}^{\text{--}}{\text{ + Si}}{\left( {{\text{OH}}} \right)_{\text{4}}}\end{equation}

During this irreversible dissolution reaction, 1 mol of olivine sequesters 4 mol of CO₂, releasing stoichiometric amounts of HCO₃^– and dissolved metal (cat)ions to the marine pore fluid. However, this process can subsequently trigger marine clay mineral authigenesis and hence reverse weathering (Fig. 10). Depending on the species formed and the reaction pathway, these reactions may lead to the recycling or release of CO₂, meaning that the net carbon mass balance remains unknown. Although such a CO₂ sequestration technique still requires verification and optimization of the total CO₂ binding capacity at individual test sites, as well as estimation of the costs and potential environmental risks, this approach has great potential for application on a large scale (e.g. Montserrat et al., 2017).

Figure 10. (a) Field-scale (carbonate) alkalinity enhancement test site situated in the pioneer vegetation zone of salt marshes in Portugal. (b) Illustration showing mineral reactions and element transfer at the interface between (1) seawater and substrate (deployment: reactive mafic rock rich in olivine), (2) substrate and biologically active top layer and (3) top layer and marsh sediments. Olivine dissolution consumes CO₂ and liberates carbonate alkalinity and elements, causing smectite formation. (c) Evolution of the alkalinity profiles in sediment porewater with time at control sites without deployments (marked by the geogenic background curve with a low alkalinity) and at deployment sites (marked by the green curve with a ∼2.5-fold higher alkalinity due to enhanced olivine weathering). Note that the deployment sites approach the geogenic background alkalinity level in a relatively short timeframe, which suggests that reverse weathering takes place (i.e. CO₂ release).

Based on CO₂ mass balance calculations (Equation 10), olivine deployments in high-energy marine–coastal environments have been proposed (e.g. Meysman & Montserrat, 2017). On main shipping routes, olivine release from vessel tanks is currently under consideration in the open ocean (e.g. Köhler et al., 2013). Laboratory and modelling studies have demonstrated the great potential of alkalinity enhancement by chemical weathering of silicates, with recognition of mineral-specific CO₂ uptake rates, nutrient release rates and ecological effects (e.g. Monserrat et al., 2017; Bach et al., 2019; Fuhr et al., 2022).

Despite the significant interest, field-scale experiments in marginal marine environments (under natural conditions) are still scarce, and the counteracting role of any reverse weathering reaction (e.g. Wallmann et al., 2023) remains largely unknown. In a recent study, He & Tyka (2023) proposed that at favourable near-coastal sites, such as in north Madagascar, California, Brazil, Peru and locations close to the Southern Ocean, a carbon uptake efficiency of ∼0.6–0.8 mol CO₂ per mol of alkalinity could be reached after 3–4 years. Moreover, Mendes et al. (unpublished data) are currently conducting alkalinity enhancement experiments by deploying picro-basalt and basanite in the pioneer vegetation zone of salt marshes in Portugal and monitoring their impacts on the ecosystem and porewater chemistry (Fig. 10). Preliminary results indicate that: (1) porewater alkalinity increases by a factor of three within a few days of the deployments, maintaining an alkalinity that is ∼10–20% higher compared to that at control sites (without deployments) after 6 months; and (2) the release of dissolved Si(OH₄), Mg²⁺ and Fe²⁺ ions from olivine, pyroxene and feldspar dissolution triggers the formation of smectite at as-yet unknown rates, confirming the importance of reverse weathering reactions in marine settings (e.g. Han et al., 2022b; Mendes et al., unpublished data).

However, before this type of geoengineering can be applied to marine settings on a larger scale (e.g. gigaton scale; He & Tyka, 2023), potential side effects and consequences have to be identified and carefully monitored, such as: (1) the formation of distinct chemical and microbial environments on deployed mineral particles and grain surfaces; (2) the response of the biological community to the nutrients released by the deployments; (3) the evolution of porewater geochemistry; (4) the long-term links between increased carbonate alkalinity, porewater pH and carbonate mineral saturation indices; (5) the rate of CO₂ re-equilibration with the atmosphere; and (6) the impact of smectite authigenesis on the physicochemical properties of the marine sediment (e.g. Simon et al., 2002; van der Jagt et al., 2018; Hartmann et al., 2023).

Clay minerals as slow-release fertilizers

Clay minerals that formed during continental chemical weathering, as well as during reverse weathering, are increasingly being used in agriculture for soil conditioning and soil environmental improvement. The addition of these minerals helps mitigate the accumulation of excess nutrients in soils and water bodies worldwide caused by the application of conventional, easily soluble (K-N-P-based) fertilizers. The overuse of common fertilizers leads to burning or nitrate poisoning of plants and to the increased eutrophication of freshwater bodies, as well as to increased emissions of the greenhouse gas nitrous oxide. Hence, modern agricultural practices focus on the efficient use of sustainable ‘green’ fertilizer technologies (e.g. Tilman et al., 2001). One solution to this challenge is the use of functional fertilizers, which enable the controlled or slow release of nutrients (e.g. Sharma, 1979; Liu et al., 2022). Slowly dissolving silicate-based fertilizers, such as glauconite (or greensand), aim to deliver nutrients to plants in a targeted (and directed) manner, thus increasing crop yields while reducing environmental pollution. These controlled-release products typically consist of nutrients (e.g. N, P or other macro- and micronutrients) that are loaded onto a clay mineral-based carrier (e.g. glauconite), whose properties determine the release rate of the nutrients (e.g. Fu et al., 2018; Rahman et al., 2021; Duan et al., 2023; Moradi et al., 2023). Moreover, it has been found that, in the long term, additional secondary clay mineral assemblages may form, such as ferruginous smectite, illite-smectite and metal (hydr)oxides (e.g. Pestitschek et al., 2012), which help to maintain the favourable physicochemical properties of the soil, such as a greater pH buffer capacity, an elevated organic matter content and increased water retention and ion-exchange capacities (e.g. Li et al., 2019a).

Clay minerals such as smectites (e.g. Pereira et al., 2015; Ren et al., 2024), kaolinite (e.g. Lei et al., 2018; AlShamaileh et al., 2022), halloysite (e.g. Maximov et al., 2023) and glauconite (e.g. Vakal et al., 2020; Rudmin et al., 2023b) have been investigated in the search for cost-effective and environmentally friendly nutrient-carrier materials. Current research has focused on glauconite because this mineral can act as a carrier for nitrogen (N) and as a natural source of K. For example, glauconite loaded with ammonium (NH₄⁺) ions, carbamide (urea) or ammonium dihydrogen phosphate has been proposed as a promising slow-release fertilizer (e.g. Rudmin et al., 2023b). It was found that the adsorption of NH₄⁺ ions into the micropores of glauconite or in the interlayer sites of smectite within interstratified glauconite-smectite occurs immediately after the preparation of the mixtures. It was suggested that the degree of NH₄⁺ adsorption by globular glauconite depends on the distribution and abundance of surface-accessible pores and on the smectite content. Importantly, such glauconite fertilizers exhibited prolonged release times for N and K, making them suitable for diverse slow-release fertilizing purposes (e.g. Rudmin et al., 2023b). These findings advance our understanding of the influence of environmental conditions on reverse weathering. Future studies should make use of isotopes to quantify the rates of silicate weathering versus near-surface clay mineral neoformation and also explore how the application of large volumes of slow release clay-based fertilizers to soils may eventually, after chemical degradation and transport, impact reverse weathering processes in marine environments.

Conclusion and perspectives

Reverse weathering processes produce a suite of authigenic clay minerals in marine sediments over comparatively short geological timescales. Despite the generally sluggish reaction kinetics of near-surface clay mineral authigenesis, there is increasing evidence from modelling, isotopic and experimental studies for the occurrence of faster reactions and substantial element sequestration (K, Si, Mg, REEs, Li, B, etc.) associated with reverse weathering. This takes place in shallow-marine and deep-marine depocentres, as well as in marine hydrothermal surroundings. A complex interplay occurs between global climate change and elemental fluxes to the oceans, which are intimately related to the rates of continental and seafloor chemical weathering, the evolution of seawater biogeochemistry and elemental uptake fluxes associated with marine clay mineral sinks. Novel analytical approaches used to quantify clay mineral reactions are advancing, which will allow better estimations of the efficiency of reverse weathering. These advances are driving applications and proposals for clay mineral-related geoengineering for CO₂ sequestration and for slow-release fertilizers as soil amendments. However, further progress in understanding the links between silicate weathering and reverse weathering is needed, which requires continued field studies, advanced laboratory approaches and numerical modelling. Such advancement underlines the great relevance of reverse weathering reactions on past and present marine element cycles and Earth’s climate.