# Paleozoic and Mesozoic oceanic anoxic events and biotic crises

Haijun Song, Alexandre Pohl, Alexander Dunhill, Xi Chen

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Haijun Song, Alexandre Pohl, Alexander Dunhill, Xi Chen. Paleozoic and Mesozoic oceanic anoxic events and biotic crises. Global and Planetary Change, In press, pp.105329. 10.1016/j.gloplacha.2026.105329. hal-05486510

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# Paleozoic and Mesozoic oceanic anoxic events and biotic crises

Haijun Songª, Alexandre Pohlᵇ, Alexander Dunhillᶜ, Xi Chenᵈ

ªState Key Laboratory of Geomicrobiology and Environmental Changes, School of Earth Sciences, China University of Geosciences, Wuhan, 430074, China

ᵇUniversité Bourgogne Europe, CNRS, Biogéosciences UMR 6282, 21000 Dijon, France

ᶜSchool of Earth and Environment, University of Leeds, LS2 9JT Leeds, United Kingdom

ᵈInstitute of Earth Sciences, China University of Geosciences, Beijing 100083, China

## Abstract

Anthropogenic global warming is causing rapid deoxygenation of the modern oceans, threatening marine ecosystems. The multiple episodes of oceanic anoxic events (OAEs) that occurred during geological history provide a crucial perspective for understanding their mechanisms and consequences. This paper presents the classification system of ocean redox states (from oxic to euxinic), and summarizes the different indicators used for paleo-redox reconstructions such as lithofacies, biofacies, minerals, redox-sensitive elements, isotopes, and biomarkers. Based on the 30 studies in this special issue, we synthesize the following advances: (1) spatio-temporal evolution of ocean anoxia from the Cambrian to the Late Paleozoic and its coupling with the carbon and nitrogen cycles; (2) triggering and driving roles of volcanic activity, climate fluctuations, and enhanced weathering on OAEs; (3) responses of biota to OAEs, including extinction, recovery, and community reorganization; (4) responses of terrestrial ecosystems to OAEs, such as vegetation turnover and increase of lacustrine organic carbon burial; and (5) the modulating role of astronomical orbital cycles (especially eccentricity) on organic matter burial during OAEs. These multidisciplinary results deepen our understanding of the mechanisms and ecological responses of deep-time OAEs and provide geological analogies and theoretical bases for assessing the future trends of modern ocean deoxygenation.

**Keywords:** Phanerozoic, redox, hypoxia, deoxygenation, mass extinction

## 1. Introduction

The oceans are undergoing a rapid process of deoxygenation due to anthropogenic global warming (Breitburg et al., 2018b; Keeling et al., 2010; Limburg et al., 2020), which poses a serious threat to marine
life and ecosystems (Morée et al., 2023; Rose et al., 2024). How to accurately predict the future process of global ocean deoxygenation and assess its impacts on the ecosystem is a major scientific question. The answer to this question may lie in the fossil record as there have been many periods of oceanic deoxygenation, many of which coincide with mass extinction events, throughout Earth history (Kiessling et al., 2024; Meyer and Kump, 2008).

During the Paleozoic and Mesozoic, ocean deoxygenation occurred many times, leading to regional or global oceanic anoxic events (OAEs) and mass extinctions, providing a window into the mechanisms of ocean-atmosphere-biosphere coupling during ocean deoxygenation. Here, we present a special issue entitled "Paleozoic-Mesozoic Oceanic Anoxic Events and Biotic Crises". This special issue includes 30 original research papers that deepen our understanding of the driving mechanisms, biogeochemical processes and ecological responses of past oceanic anoxic events and their implications for current ocean deoxygenation. These papers cover multidisciplinary topics such as redox change (5 papers), causes of OAEs (5 papers), biotic responses to OAEs (11 papers), the impact of OAEs on terrestrial ecosystems (4 papers), and astronomical cycles during OAEs (5 papers).

This special issue mainly focuses on deep-time anoxic events. To better understand the characteristics and evolution patterns of anoxic events, we have summarized the classification system of marine redox conditions, clarified the definition standards, environmental and biological response features of different redox states. Additionally, we have summarized the key research methods for reconstructing paleoceanic redox conditions. These methods include both the recent technologies applied in the research papers of this special issue, and other classic proxies used in the literature. Through a systematic review of the classification and methods, we hope to provide readers with the necessary background and tools to deeply understand the research results of this special issue and the field.

## 2. Classification of redox conditions

Redox conditions can be classified in various ways across different perspectives. The most fundamental classification is based on dissolved oxygen level in seawater. According to Tyson and Pearson's classification scheme (Tyson and Pearson, 1991), it can be divided into four regimes: oxic (2.0-8.0 mL/L dissolved oxygen), dysoxic (0.2-2.0 mL/L), suboxic (0-0.2 mL/L), and anoxic (0 mL/L). Considering that euxinic (hydrogen sulfide-rich) is also a common type of seawater condition in anoxic oceans (Fakhraee et al., 2025; Meyer and Kump, 2008), we list euxinic ([O₂] = 0 mL/L, [H₂S] > 0 mL/L) as a unique redox state in the ocean (Fig. 1).
Corresponding to the environment facies classified based on oxygen content levels, biofacies can be divided into five types (Fig. 1). The first four follow the classification schemes of previous researchers (Tyson and Pearson, 1991; Wignall and Hallam, 1991): aerobic (high diversity and abundance of benthic organisms, widespread bioturbation, and well-mixed sediments, with laminations rarely preserved); dysaerobic (decreased benthic diversity, dominance of low-oxygen-tolerant groups, weakened but still present bioturbation); quasi-anaerobic (extremely poor benthic communities, only a few tolerant/opportunistic groups briefly appear, very weak bioturbation, and laminations are relatively well preserved); and anaerobic (difficult for macrobenthic organisms to survive, bioturbation is usually absent or extremely weak, organic matter is easily preserved, and laminations are well preserved). The fifth type sulfidic anaerobic refers to the biofacies under anoxic conditions accompanied by hydrogen sulfide ($H_2S$). This biofacies shows that oxygen and oxidants such as nitrate/iron and manganese oxides have been exhausted, and the system is dominated by sulfate reduction (Fakhraee et al., 2025; Meyer and Kump, 2008). Under sulfidic anaerobic conditions, the abundance of sulfide minerals (such as pyrite) increases (Li et al., 2025a; Lyons and Severmann, 2006).

Normoxic, hypoxic, and anoxic mainly reflect the physiological responses of living organisms at different oxygen content levels (Kamykowski and Zentara, 1990; Tyson and Pearson, 1991). Normoxic conditions support the physiological activities of all regular aerobic organisms, and there is no hypoxic stress response in organisms. Hypoxic conditions pose physiological stress to most aerobic organisms, and organisms will initiate a series of stress and adaptation mechanisms such as improving respiratory efficiency and changing metabolic pathways (Kamykowski and Zentara, 1990). Under anoxic conditions, only a few animals (e.g., some foraminifera and polychaetes) can survive for a short time through anaerobic metabolism, and anaerobic microorganisms can survive (Diaz and Rosenberg, 2008; Levin et al., 2009). To align with environmental facies and biofacies, this paper also discusses the physiological responses of organisms under two other redox states (near-anoxic, euxinic) (Fig. 1). Under near-anoxic conditions, the vast majority of aerobic organisms cannot survive for a long time, and only a few can survive briefly (Seibel, 2011). In euxinic environments, sulfur-oxidizing bacteria and photosynthetic sulfur bacteria are common, and a few animals like polychaetes *Ampharete* and *Magelona* can also live (Levin et al., 2009; Li et al., 2025a).

Deoxygenation refers to the process in which the dissolved oxygen content in water column or sediments decreases (Keeling et al., 2010). Ocean deoxygenation focuses more on the "process/trend", while dysoxia, suboxia, and anoxia are more inclined to represent "states" or "thresholds" (Fig. 1). The main features of ocean deoxygenation include the expansion of hypoxic areas, the vertical extension of low-oxygen water masses, and the prolongation of hypoxic durations (Breitburg et al., 2018a;
Stramma et al., 2008). Dysoxia, suboxia, and anoxia are the common states of the ocean deoxygenation process at different stages (Tyson and Pearson, 1991; Wignall and Hallam, 1991). Modern ocean deoxygenation is mainly driven by multiple factors such as climate warming, intensified water stratification, eutrophication, and human activities (Dai et al., 2023; Diaz and Rosenberg, 2008). Similarly, ancient OAEs were usually associated with global warming events (Kiessling et al., 2024; Song et al., 2021; Song et al., 2019; Wu et al., 2025). The degree of modern ocean deoxygenation is relatively mild compared to typical ocean anoxic events of the geological past, representing an early stage of deoxygenation (Fig. 2). Especially from the perspective of the area of ocean anoxia, the difference is substantial: the current ocean anoxic area is ~0.2-0.3% (Reershemius and Planavsky, 2021), while the maximum estimated area of ocean anoxia during the Permian-Triassic boundary interval was 65% (Zhang et al., 2018). Ocean oxygenation is the reverse process of ocean deoxygenation (Fig. 1), referring to the process by which the dissolved oxygen content in the marine environment recovers and dysoxic-anoxic areas shrink. The early ocean was anoxic, and it experienced a long period of ferruginous and euxinic conditions, along with multiple major oxidation events, to become an oxygenated ocean (Lyons et al., 2014; Pohl et al., 2022).

**3. Methods for paleo-redox reconstruction**

There are many methods for studying the redox conditions of the ocean, including lithofacies, biofacies, mineralogy, elements, isotopes, and biomarkers (Table 1). Since a single indicator is prone to being affected by diagenesis, provenance, and regional factors, reconstructing the redox state of ancient oceans and lakes often requires a comprehensive synthesis of multiple indicators (Algeo and Li, 2020). Most of the related studies in this special issue have used multiple methods for the characterization of redox conditions.

Lithofacies is mainly used to reconstruct the redox conditions of the ocean based on the type of rocks and the degree of bedding development. Black shale typically indicates a low-oxygen to anoxic environment, mainly because organic matter is more easily preserved under low-oxygen to anoxic conditions and causes the sediment color to darken (Arthur and Sageman, 1994; Wignall, 1994). The change from "unclear bedding" to "clear and continuous lamination" usually reflects the transition from strong to weak bioturbation and the decrease of benthic community activity, which is often interpreted as a trend of decreasing dissolved oxygen at the ocean-sediment interface and an increase in the degree of anoxia (Pascal et al., 2024; Savrda and Bottjer, 1986).

Biofacies uses the fossil record to infer redox conditions, with the main indicators including the
community composition and diversity of benthic organisms as well as the intensity of bioturbation (Table 1). From oxygen-rich to oxygen-poor environments, the types and diversity of benthic organisms show regular changes: in oxic environments, the benthic community is prosperous and highly differentiated; as oxygen content decreases, sensitive species disappear first, opportunistic species (such as certain worms) may temporarily thrive; in suboxic environments, benthic macrofauna almost disappears (Diaz and Rosenberg, 2008; Rhoads and Morse, 1971; Wignall and Hallam, 1991). The intensity of bioturbation changes concomitantly, from strong, large-scale, and diverse bioturbation structures in oxic environments, to weak, small-scale traces in suboxic environments, and to the complete absence of bioturbation structures in anoxic environments (Droser and Bottjer, 1986; Savrda and Bottjer, 1986; Taylor and Goldring, 1993).

The formation of certain authigenic minerals is related to specific redox conditions. Authigenic pyrite is the final product of sulfate reduction and its abundance (usually expressed as the degree of pyritization, DOP) increases with increasing severity of anoxic conditions (Raiswell and Berner, 1985; Raiswell and Canfield, 1998). The size distribution of framboidal pyrite has important indicative significance: framboidal pyrite formed in oxic to dysoxic bottom water has large grain sizes and a wide size-distribution range; while in euxinic water columns, due to high hydrogen sulfide concentration and rapid nucleation, the grains of framboidal pyrite are small and their size distribution is narrower, which is one of the key mineralogical indicators for determining the euxinic environment (Li et al., 2022; Wignall and Newton, 1998; Wilkin et al., 1996). The formation of authigenic barite is often related to the interface reaction between barium released during organic-matter degradation and seawater sulfate, and its enrichment is an important indication of ancient euxinic facies (Han et al., 2022; Paytan and Griffith, 2007; Torres et al., 1996).

Redox-sensitive elements (RSEs) such as V, U, Mo, and Ce exhibit a certain degree of enrichment in sediments or rocks, which is a crucial method for quantitatively assessing ancient redox states (Algeo and Tribovillard, 2009; Bennett and Canfield, 2020; Tribovillard et al., 2006). Under reducing conditions, some elements transform from soluble to insoluble states and accumulate in the sediments. Generally, the contents of V, U, Mo, Ni and Cu increase markedly from oxic to anoxic/sulfidic environments (Tribovillard et al., 2006). The iron speciation is another key indicator, including the ratio of pyrite iron (Fe<sub>py</sub>), highly active iron (Fe<sub>HR</sub>) to total iron (Fe<sub>T</sub>) (such as Fe<sub>HR</sub>/Fe<sub>T</sub>), which can effectively distinguish between oxic, anoxic non-sulfidic (ferruginous) and euxinic environments (Poulton, 2021; Poulton and Canfield, 2011; Raiswell and Canfield, 1998). The Ce anomaly can serve as a sensitive indicator of the water-column redox state, showing a remarkable increase in the ratio of insoluble Ce (IV) to dissolved Ce (III) and resulting in a positive anomaly in seawater under anoxic
conditions (German and Elderfield, 1990; Song et al., 2012; Tostevin, 2021).

A variety of isotope systems provide a unique perspective for redox studies. The U and Mo isotopes usually change systematically with the degree of ocean reduction, making them excellent redox indicators. The fractionation degree of uranium isotopes (δ²³⁸U) in carbonates can be used to estimate the area of global anoxia because uranium undergoes pronounced isotopic fractionation when buried in anoxic sediments (Andersen et al., 2014; Weyer et al., 2008). The molybdenum isotopes (δ⁹⁸Mo and δ⁹⁷Mo) in shales are used to distinguish between the oxic and anoxic conditions of seawater (Arnold et al., 2004; Dickson, 2017; Scott and Lyons, 2012). Iron isotopes (δ⁵⁶Fe) can also reflect the redox cycle (Johnson et al., 2008). The fractionation amplitude of sulfur isotopes (δ³⁴S) in pyrite is controlled by the rate of sulfate reduction and the openness of the sulfate reservoir, and can be used to infer the ancient marine sulfate concentration and microbial activity intensity (Canfield, 2001; He et al., 2020; Song et al., 2014). The changes in nitrogen isotopes (δ¹⁵N) reflect the transformation of the nitrogen cycling processes (such as denitrification, anaerobic ammonia oxidation), which are closely related to anoxic water bodies (Du et al., 2025b; Pellerin et al., 2024; Sigman et al., 2003).

Biomarkers are molecular fossils left by specific organisms in sediments, providing direct evidence of biogeochemical processes. The key biomarkers used for redox determination mainly come from green sulfur bacteria. These photosynthetic bacteria require light, hydrogen sulfide, and anoxic conditions, and their markers such as isorenieratene are important evidence for proving the euxinic environment (Brocks et al., 2005; Cui et al., 2020; Summons and Powell, 1987). The abundance changes of isorenieratene can semi-quantitatively indicate the degree of euxinia and persistence of euxinic conditions within the photic zone (Kuypers et al., 2004). The Pr/Ph (pristane/phytane) ratio is often used in petroleum geochemistry to distinguish relative oxidizing/reducing sedimentary environments, but this indicator can be affected by the source of organic matter and thermal maturity (Didyk et al., 1978; Evenick, 2016).

# 4. Major processes in this special issue

## 4.1 Redox change

Revealing the changes in ocean redox in Earth history is crucial for understanding the causes of ocean anoxic events and their impact on marine life. The following studies explore the evolution of redox states during multiple geological periods through diverse geochemical proxies (Table 2; Fig. 3), including the Cambrian (Li et al., 2026a; Meng et al., 2026; Wang et al., 2025b), mid-Silurian (Maaleki-
Moghadam et al., 2025), and late Pennsylvanian (Yang et al., 2026).

Meng et al. (2026) identified three negative carbon isotope events (with remarkable negative shifts in δ¹³Corg and δ¹³Ccarb) in the Yangtze Block, South China, and used δ¹⁵N and Ce anomalies to investigate redox changes. The results showed that the Fortunian-early Stage 2 was an oxic to suboxic environment, which was replaced by anoxic-euxinic conditions in Stage 3. The Co×Mn index suggests strong upwelling activity in Fortunian-early Stage 3. The upwelling transported ¹³C-depleted dissolved organic carbon (DOC) to the shallow continental shelf, where denitrification in the oxygen minimum zone (OMZ) produced ¹³C-depleted dissolved inorganic carbon (DIC), thereby triggering negative carbon isotope events. This study coupled redox changes with the carbon and nitrogen cycles, emphasizing that "nitrate as an oxidant" may, to some extent, delay the further depletion of O₂, while increases in nutrients and DIC promoted the early biological proliferation and bio-mineralization.

Li et al. (2026a) used carbonate I/(Ca+Mg) and Ce/Ce* to reconstruct the redox changes in the eastern Yangtze shallow shelf of South China. They found persistent anoxia from the top of the Cambrian Stage 3 to the end of Stage 4. The oxygen content in the shallower K2 section was higher than that in the deeper WN2 section, indicating a gradient of dissolved oxygen between shallow and deep waters. By integrating the published data, they inferred oxygen-limited conditions in the Stage 4, suggesting that shallow-water anoxia might be the cause of metazoan extinction in this interval.

Wang et al. (2025b) reported positive shifts in δ¹³Ccarb and δ¹³Corg during the SPICE (Steptoean positive carbon isotope excursion) event in South China, accompanied by a negative shift in δDorg. The onset timing and intensity of the SPICE event varied in different sedimentary environments. This finding challenged the "global synchronous anoxic event" perspective. The authors proposed that the enhancement of organic carbon burial and isotope fractionation changes might be related to the expansion of phytoplankton, with different types of phytoplankton generating different isotope responses across facies.

Maaleki-Moghadam et al. (2025) examined the mid-Silurian Mulde/Lundgreni event in two basins: Laurentia (Canadian Arctic) and Baltica (Latvia). All sections showed synchronous positive shifts in carbon and sulfur isotopes, indicating an increase in organic carbon and pyrite burial. The bottom water in the Laurentia basin remained ferruginous and anoxic throughout the study interval, while the bottom water in Baltica experienced larger redox fluctuations, shifting from possibly oxic to ferruginous conditions during the Mulde carbon-isotope excursion. The study argued that climate-driven sea-level changes dictated temporal trends in ocean deoxygenation, thereby influencing the timing and selectivity
of biotic extinction.

Yang et al. (2026) used the δ¹⁵N records from three carbonate slope sections in South China across the late Pennsylvanian Kasimovian-Gzhelian boundary to investigate the changes in marine redox conditions and primary productivity. The results showed that the δ¹⁵N records from all three sections exhibited a four-stage change: a slow decline (~12‰ to ~8‰), a rapid increase to 15‰, a negative excursion (15‰ to 8‰), and subsequent stabilization. The dynamics of δ¹⁵N reflect the adjustment of denitrification and nitrogen fixation processes in the water column in response to marine redox changes. The negative shift of δ¹⁵N below the Kasimovian-Gzhelian boundary indicates enhanced nitrogen fixation. This phenomenon is closely related to the expansion of the oxygen-deficient area in the ocean during the Kasimovian-Gzhelian warming event.

## 4.2 Causes of OAEs

The following studies explore the causes of OAEs from different geological periods through geochemical analysis, numerical simulation, stratigraphy and biostratigraphy methods (Table 2; Fig. 3). The driving factors of OAEs include volcanic activity (Du et al., 2025a; Zhang et al., 2026c), sea-level changes and glaciation (Algeo et al., 2026), changes in nutrient input caused by enhanced continental weathering (Cui et al., 2025), and the synergistic effects of multiple factors (Li et al., 2026b). These studies show that multiple factors and their synergistic effects have significant impacts on the initiation and evolution of OAEs.

Du et al. (2025a) identified four Late Triassic carbon-isotope excursions (CIEs) in the Tela section in southern Tibet. The Initial-CIE related to end-Triassic mass extinction is considered to be associated with the activity of the Central Atlantic Magmatic Province (CAMP), while the causes of the other three CIEs during the Sevatian (upper Norian) remain unclear. The long-term volcanic activity related to seafloor spreading in the Neotethys and the rifting of the Lhasa Block from eastern Gondwana, which occurred from the Sevatian to Norian-Rhaetian boundary, may be important driving factors for global carbon cycle disturbances during this interval.

Zhang et al. (2026c) measured late Cenomanian to early Turonian mercury (Hg) concentrations and isotopes in the southern Tethyan Himalaya. Constrained with these data, results of an oceanic Hg box model suggest that volcanic Hg emissions increased by a factor of 8 to 16 during OAE2, hence quantitatively supporting the volcanic trigger of OAE2. Mercury enrichment factors showed remarkable spatial heterogeneity globally. Volcanic eruption type (submarine vs. subaerial), paleogeographic isolation and the distance from the large igneous provinces may have driven the
regional differences in Hg enrichment. These findings suggest that when using Hg as "volcano trigger evidence", spatial patterns arising from transportation and deposition need to be considered.

Algeo et al. (2026), based on the geochemical data from the West Valley core in western New York State, USA, inferred that there was a brief (about 126,000 years) glaciation during the Frasnian-Famennian mass extinction, characterized by orbitally-paced climate cycles. There were three salinity reduction events at the multi-million-year timescale near the Frasnian-Famennian boundary, suggesting that sea-level fall caused the study area to briefly enter a low-salinity surface water layer. At a higher temporal frequency during the glaciation, orbital-scale climate fluctuations controlled sea level and the flux of organic matter and terrigenous sediments from the continent, and thereby modulated the redox state of water and the formation of organic carbon-rich layers. This study demonstrated a modulation of the sedimentary content in organic carbon by climate changes and associated cycles of ice-sheet waxing and waning at the Frasnian-Famennian boundary.

Cui et al. (2025) used an Earth system model forced with data-derived $pCO_2$ changes to quantify the relative contributions of nutrient input, phosphorus recycling and ocean circulation changes to the expansion of oceanic anoxia during OAE2. Model results indicate that bottom-water anoxia already existed in the equatorial Atlantic and along the Tethyan margins prior to OAE2, and only changes in $pCO_2$ had a relatively limited impact on the area of anoxia. Instead, increased nutrient supply from enhanced continental weathering can substantially expand anoxia, and phosphorus recycling within low-oxygen marine areas can further amplify this expansion. Through comparison with geological records, this study estimated that the anoxic bottom water of OAE2 may only account for about 20-40% of the global seafloor area, and the continental phosphorus weathering intensity increased by about 40-60% compared to pre-event.

Li et al. (2026b) found that the mean $\delta^{13}$C value in the late Albian eastern Neo-Tethys was significantly higher than that in other areas, challenging the perspective of OAE1d being a strictly synchronous event globally. The research suggests that the low-latitude location, a restricted epicontinental sea environment, abundant nutrient supply, and volcanic activity jointly created nutrient-rich water masses leading to conditions of high marine primary productivity, which are the key drivers for the increase in $\delta^{13}$C and the development of OAE1d.

## 4.3 Biotic response

Ocean hypoxia or anoxia is one of the most crucial factors influencing the evolution and extinction of marine animals (Pohl et al., 2023; Song et al., 2024; Wang et al., 2025a). The following articles
investigate the responses of organisms to marine anoxic events and related environmental disturbances (Table 2; Fig. 3). The organisms covered include microorganisms, plankton, and macro-invertebrates. The response types include extinction (Cao et al., 2025), recovery (Zhang et al., 2026b), community changes (Fernández et al., 2026; Huang et al., 2026; Sequero et al., 2026), diversity and abundance variations (Chang et al., 2025; Li et al., 2025b), productivity (Sosa-Montes de Oca et al., 2026), malformation (De Backer et al., 2026; Vancoppenolle et al., 2026), and exceptional fossil preservation (Varejão et al., 2025).

Cao et al. (2025) examined the linkage between nitrogen-cycle perturbations and biotic responses during the terminal Ediacaran period. They found that the ocean had a stratified structure with oxic surface water and anoxic to euxinic deep water during the deposition of the Doushantuo Formation. The isotopic anomalies in the lower Shibantan Formation reflect expansion of deep-water anoxia and enhanced organic-carbon burial, together with intensified anaerobic nitrogen-cycling processes. The expansion of anoxic water and the subsequent nitrogen limitation might constitute an important environmental stress for the decline of biota in the end-Ediacaran.

Zhang et al. (2026b) reported that the late Ordovician mass extinction was related to coeval marine anoxia, sea-level fluctuations, and glacial events. In the early Silurian (Rhuddanian), increasing surface-ocean oxygenation drove a gradual biotic recovery, with plankton recovering earlier than benthos, indicating that redox improvement first alleviated ecological stress in surface habitats. The biotic radiation in South China started at the beginning of Aeronian (ca. 440.8 Ma), earlier and faster than in Laurentia, Baltica, and Avalonia, highlighting the driving role of ocean oxygenation in regional heterogeneity in the pace of recovery.

De Backer et al. (2026) explored the drivers and biotic consequences of Silurian OAEs through a detailed study of chitinozoan malformations and geochemistry in the Mid-Ludfordian OAE of Gotland, Sweden. Analysis of over half a million specimens reveals malformations coinciding with Mn, Zn, and Pb enrichments, often preceding carbon isotope excursions, indicating that marine biotic stress anticipated global carbon cycle perturbations. Coupled with rapid shifts in marine Sr isotopes, these findings support a model in which episodic venting of metalliferous hydrothermal brines drove pulsed trace metal enrichments, triggering both environmental stress and anoxia. The study highlights how microscopic fossil indicators can sensitively record the interplay between geochemical disturbances and marine ecosystem responses during deep-time anoxic events.

Vancoppenolle et al. (2026) evaluated the usefulness of chitinozoan malformation frequency as a
proxy for biological stress during the early Wenlock Ireviken oceanic anoxic event. Analyses of chitinozoan specimens from Gotland, Sweden show that the chitinozoan malformation rate during the Ireviken event was 3-33 times the background value. The authors suggest that chitinozoan malformation was related to the redox instability and metal pollution.

Fernández-Martínez et al. (2026) investigated the effects of Late Devonian anoxic pulses, including the Kellwasser Event, on macrobenthic communities in the North American Seaway (NAS) using detailed ichnological analysis. Across five sites, they document shifts in ichnoassemblages reflecting alternating oxic and anoxic conditions, with Givetian deposits showing the highest ichnodiversity and Famennian deposits recording more persistent anoxia. Regional variability is pronounced: inner NAS sites maintained oxic conditions during the Upper Kellwasser Event, whereas open-marine settings experienced stronger biotic impacts and suppressed bioturbation. Their study demonstrates that Late Devonian anoxic episodes were heterogeneously expressed across different marine environments, highlighting the critical role of local and regional conditions in shaping the intensity of global biotic crises and benthic recovery dynamics.

Chang et al. (2025) used the grain size and morphology of framboidal pyrite from the Baoshan block to reconstruct the redox fluctuations spanning the latest Devonian through middle Permian. The study suggests that frequent anoxic/euxinic events during the Asselian and Sakmarian were driven by glacial pulses and sea-level fluctuations; contemporaneous biota was dominated by brachiopod-bryozoan-crinoid assemblages with relatively low diversity. In contrast, the post-glacial Kungurian-Wordian interval was characterized mainly by stable, dysoxic conditions and a marked increase in faunal abundance. The authors conclude that both global cooling and warming can influence biodiversity and abundance through redox changes.

Huang et al. (2026) used biomarker data to explore the response of marine eukaryotic algae after the Permian-Triassic mass extinction. They found that the algal community in the Early Triassic was not characterized by a simple collapse-recovery pattern, but presented a complex reorganization process. Algae that preferentially produced more $C_{28}$ steranes (similar to certain groups in modern oceans) became abundant in the Early Triassic.

Li et al. (2025b) investigated foraminiferal changes during the Late Triassic Carnian Pluvial Episode (CPE) in southwestern Guizhou and southern Tibet. The study found that although the global climate and marine environment underwent drastic changes during the CPE, the foraminifera in the study areas did not show significant extinction. Foraminiferal diversity and abundance dropped sharply at Erguan section in Guizhou, whereas no clear shift was observed at Quxia section in Tibet. This
regional heterogeneity may be related to the differences between the two regions in the magnitude of warming, the input of siliciclastic material, and the severity of marine anoxia.

Sequero et al. (2026) analyzed the response of the biotic and environmental system during the onset of OAE1a. The study found that within thousands of years after the initiation of the carbon cycle disturbance, biotic and environmental responses developed, and these responses occurred earlier than widespread marine anoxia and the enhancement of primary productivity. Major environmental shifts (including extensive anoxia and elevated productivity) occurred ~29 kyr after the onset of the carbon cycle perturbation, while the methane injection event occurred 6 kyr after the initiation of OAE1a.

Varejão et al. (2025) discussed biotic responses from a taphonomic perspective, focusing on the coupling between multiple Lower Cretaceous exceptional fossil preservation deposits (Konservat-Lagerstätten) in Brazil’s Araripe Basin and transgressive-anoxic conditions. The multiple marine transgressions during the opening process of the South Atlantic gradually submerged the marginal/continental environments and expanded the oxygen minimum zone, subsequently creating a window for exceptional preservation. The Batateira fossil layer of the Barbalha Formation and the Santana Lagerstätte of the Romualdo Formation corresponded to OAE1a and OAE1b, respectively, indicating that global geological events could lead to extreme preservation conditions through regional marine transgression and anoxia.

Sosa-Montes de Oca et al. (2026) analyzed multiple proxies, including biomarkers, inorganic geochemical compositions, and carbon isotope values across the K-Pg boundary at Caravaca in southeast Spain. The section spans ~22 kyr before to ~220 kyr after the K-Pg impact event. This study reveals that some aquatic biomarkers increase preceding and spanning across the K-Pg boundary from 5 kyr before to 30 kyr after, which are interpreted as a possible reflection of a pulse of productivity. This study highlights a complex ecological response to K-Pg climatic and environmental change and argues against an extended global shutdown of primary productivity.

## 4.4 Terrestrial ecosystems

The terrestrial ecosystem is an important component of the global ecosystem, but its linkage with OAEs is less understood. The following articles study the relationship between climate and the terrestrial ecosystem during the Late Triassic Carnian Pluvial Episode and T-OAE intervals (Table 2; Fig. 3), showing that the plant communities shifted (Rial et al., 2025; Zhang et al., 2026a) and that organic carbon in lakes was extensively buried, becoming an important terrestrial carbon sink, which affected the carbon cycle (Luan et al., 2025; Zhang et al., 2025).
Zhang et al. (2026a) examined the profound terrestrial ecosystem and climate changes during the Late Triassic Carnian Pluvial Episode (CPE; ~234-232 Ma) using high-resolution palynological records from the Jiyuan Basin, North China. Their data reveal four phases of vegetation turnover, shifting from xerophytic to hygrophytic communities, each closely aligned with volcanic activity from the Wrangellia Large Igneous Province, as evidenced by carbon isotope excursions and mercury anomalies. While most eruptions prompted resilient vegetation recovery resembling pre-eruption floras, the second eruption triggered a more dramatic shift, increasing plant diversity and establishing a modernized floral composition that persisted beyond the CPE. These findings illuminate the complex interplay between LIP volcanism, CO₂-driven climate perturbations, and differential terrestrial ecosystem responses during deep-time anoxic and greenhouse events.

Rial et al. (2025) used systematic palynological analysis of the Rodiles Formation in the northwest of Spain to find that the plant communities underwent remarkable changes during the T-OAE interval. Before the T-OAE, vegetation was dominated by Cheirolepidiaceae (drought-tolerant conifers), indicating a dry and semiarid background. During the T-OAE interval, ferns became dominant, indicating an increase in humidity or stronger seasonality. After the event, the Cheirolepidiaceae forests recovered, reflecting a return to drier conditions and a tree-dominated ecological structure.

Luan et al. (2025) found that the T-OAE in the Sichuan Basin spans approximately 360 kyr (182.88 -182.52 Ma) and is divided into three stages: Onset (~198 kyr), Body (~80 kyr), and Recovery (~75 kyr). The multi-taper method analysis of organic carbon isotopes shows that precession forcing affected the carbon cycle during the T-OAE, indicating that the terrestrial system is sensitive to orbital forcing. The accumulation rate of lacustrine organic carbon mass is 2 to 5 times that of the background, supporting the key role of lakes in regulating the carbon cycle.

Zhang et al. (2025) identified a $\delta^{13}$C$_{org}$ negative excursion associated with sedimentary organic carbon enrichment corresponding to the Jenkyns Event (T-OAE) in the Sangonghe Formation of the high-latitude lacustrine succession in the Junggar Basin. Overall, the interval shows inland aridification punctuated by intermittent humid phases, reflecting the hydrological instability under strong greenhouse climate. This study proposes that enhanced productivity and anoxia jointly promote the large-scale burial of organic carbon in high-latitude lakes, making them important terrestrial carbon sinks during extreme warming intervals.
## 4.5 Astronomical cycles

The following five articles focus on cyclostratigraphy and explore the relationship between astronomical orbital forcing and anoxic events (Table 2; Fig. 3). They show that the eccentricity plays a significant role in multiple events, exerting important modulation over a range of environmental effects, including carbon cycle perturbations (Chu et al., 2025; Wang et al., 2026), organic matter burial (Huygh et al., 2026), marine anoxia (Wei et al., 2026), amplification of the greenhouse effect (Chen et al., 2026), and sea level change (Wang et al., 2026). Obliquity cycles dominated sea-level change in the late Devonian period (Huygh et al., 2026).

Wang et al. (2026), based on high-resolution cyclostratigraphy and carbon-isotope data from the middle-late Cambrian Xixiangchi Formation in the Sichuan Basin (South China), identified multiple Milankovitch cycles including long eccentricity of 2.4 Myr, short eccentricity, obliquity, and precession cycles. The 2.4 Myr eccentricity peak corresponded to sea-level rise and positive δ¹³C shifts (up to +2‰), which was interpreted as widespread shallow-sea expansion promoting organic carbon burial and algal productivity enhancement. Conversely, the low sea level period corresponded to negative δ¹³C shifts, attributed to enhanced weathering and organic matter oxidation.

Huygh et al. (2026) conducted cyclostratigraphic analysis on Ti, Ti/Al, and TOC from the West Valley core in New York State, USA, revealing a strong astronomical imprint of marine anoxia/black shale deposition during the Kellwasser crisis and Rhinestreet event. By aligning the TOC record to a stable 405 kyr eccentricity cycle, a 5.67-Myr floating astronomical chronology was established. Titanium content was also modulated by eccentricity, revealing the link between organic matter preservation, detrital sediment supply, and astronomical forcing. Sea-level change was dominated by 1.2 Myr and 2.4 Myr obliquity cycles.

Wei et al. (2026) identified 405 kyr eccentricity, ~36 kyr obliquity, and ~20 kyr precession signals in the Late Permian Dalong Formation in South China. The duration of the Dalong Formation was limited to approximately 1.6 Myr (about 4 long eccentricity cycles, 253.75-252.15 Ma). Organic matter accumulation was mainly regulated by orbital eccentricity, and the maximum eccentricity value enhanced the nutrient transport and water stratification driven by the monsoon, leading to bottom water anoxia and promoting organic matter preservation.

Chen et al. (2026) established a high-resolution chronostratigraphic framework for OAE2 in the northeastern Neo-Tethys by integrating biostratigraphy, carbon isotope stratigraphy, and cyclostratigraphy. The authors defined the onset of OAE2 as 94.52 ± 0.06 Ma, with a duration of
approximately 755 kyr. The results show that the onset of OAE2 is more likely to correspond to the pulses of the Caribbean LIP (CLIP) and Ontong Java Plateau (OJP) rather than the earlier Kerguelen Plateau (KP) and High Arctic LIP (HALIP). The onset of OAE2 aligns with the first maximum eccentricity after the pulses of CLIP/OJP, indicating that the eccentricity might amplify the environmental perturbations driven by LIPs.

Chu et al. (2025) constructed an astronomical timescale framework for the Paleocene-Eocene Thermal Maximum (PETM) based on three high-resolution sections from the Fushun, Nanyang, and Tarim basins in China. The results showed that the carbon cycle perturbation was synchronous during the PETM, and the negative shift of carbon isotopes persisted for ~100 kyr, with the onset lasting from 5 kyr to 20 kyr. The spectral power of orbital eccentricity increased near the beginning of the PETM, and the two small carbon isotope excursions before and after the event in some records coincided with the maximum of the ~100 kyr short eccentricity, confirming that eccentricity regulation was a key driving factor for the carbon cycle during the PETM interval.

Declaration of Competing Interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Acknowledgements

We sincerely express our gratitude to editors Howard Falcon-Lang and Maoyan Zhu for their valuable advice and editorial work during the selection of papers for this special issue, and to all of the authors for their contributions to this special issue. We thank all the reviewers for their insightful comments. This special issue is a contribution to the State Key R&D Project of China (2023YFF0804000), the National Natural Science Foundation of China (42325202), the Natural Science Foundation of Hubei (2023AFA006) and the French *Agence Nationale de la Recherche* (ANR) project CYCLO-SED (ANR-23-CE01-0003).
Figures

<figure><img src="image_2.png" /><figcaption>**Fig. 1** Classification of redox state in multiple regimes. Terminology and classification are modified from Tyson and Pearson (Tyson and Pearson, 1991), 1 mL/L ≈ 1.43 mg/L ≈ 44.7 µmol/L (µM).</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>**Fig. 2 Relationship between oceanic oxygen level and anoxic seafloor area.** Oceanic anoxic events (OAEs) are ranked according to the global ocean seafloor area overlain by anoxic waters (Reershemius and Planavsky, 2021). PETM, Paleocene-Eocene Thermal Maximum; OAE2, Cenomanian-Turonian oceanic anoxic event; FF-OAE, Frasnian-Famennian oceanic anoxic event; T-OAE, Toarcian oceanic anoxic event; SPICE, Steptoean positive carbon isotope excursion; PT-OAE, Permian-Triassic oceanic anoxic event; H-OAE, Hirnantian oceanic anoxic event.</figcaption></figure>
<figure><img src="image_4.png" /><figcaption>**Fig. 3** Oceanic anoxic events (OAEs) in the Paleozoic and Mesozoic. Only OAEs that discussed in the special issue are included in this figure. The geologic timescale is from the International Chronostratigraphic Chart (2024 version), www.stratigraphy.org. The ages of Permian-Triassic boundary and T-OAE are from references (Dai et al., 2025; Kemp et al., 2024).</figcaption></figure>
Tables

**Table 1 Classification of redox conditions.** See the discussion in the section of classification of redox conditions for the source of data. The enrichment patterns of the redox-sensitive elements (RSE) are specific to the clastic (detrital) sedimentary system; however, in carbonate rocks or fossil-hosted systems, their enrichment characteristics often differ or even are opposite. Moreover, uranium isotopes are applied to the carbonate system, while molybdenum isotopes are typically measured in clastic sediments. BI, Bioturbation Index; DOP, degree of pyritization of iron; FeHR, highly reactive iron; FeT, total iron; Fepy, pyrite iron; Pr/Ph, pristane/phytane.

<table><thead><tr><th>Redox</th><th>Lithofacies</th><th>Biofacies</th><th>Minerology</th><th>Elements</th><th>Isotopes</th><th>Biomakers</th></tr></thead><tbody><tr><td>Oxic</td><td>No lamination</td><td>Abundant benthos BI 4-6</td><td>Little or no pyrite DOP 0-0.5</td><td>Low V, U, Mo, Ni, Cu FeHR/FeT ≤ 0.22</td><td>Very negative δ<sup>238</sup>U Negative δ<sup>98</sup>Mo</td><td>Pr/Ph >3</td></tr><tr><td>Dysoxic-Suboxic</td><td>Grey shale Discontinuous lamination</td><td>Rare benthos BI 1-3</td><td>Big framboid pyrite DOP 0.5-0.75</td><td>Low V, U, Mo, Ni, Cu FeHR/FeT 0.22-0.38</td><td>Slightly negative δ<sup>238</sup>U Positive δ<sup>98</sup>Mo</td><td>Pr/Ph 1-3</td></tr><tr><td>Anoxic</td><td>Black shale Continuous lamination</td><td>Little or no benthos BI 0</td><td>Abundant small framboid pyrite DOP 0.75-1</td><td>High V, U, Mo, Ni, Cu FeHR/FeT ≥ 0.38</td><td>0 to positive δ<sup>238</sup>U Positive δ<sup>98</sup>Mo</td><td>Pr/Ph<1</td></tr><tr><td>Euxinic</td><td>Black shale Continuous lamination</td><td>Little or no benthos BI 0</td><td>Abundant small framboid pyrite DOP 0.75-1</td><td>Very high V, U, Mo, Ni, Cu Fepy/FeHR ≥ 0.8</td><td>Very positive δ<sup>238</sup>U Very positive δ<sup>98</sup>Mo</td><td>Pr/Ph<1> Abundant Isorenieratane</td></tr></tbody></table>
<table><caption>Table 2 Summary of papers in this special issue.</caption><thead><tr><th>No</th><th>Study area</th><th>Events/intervals</th><th>Age (Ma)</th><th>Topic</th><th>References</th></tr></thead><tbody><tr><td>1</td><td>Fushun, Nanyang, Tarim basins, China</td><td>PETM</td><td>56</td><td>Astronomical cycles</td><td>Chu et al., 2025</td></tr><tr><td>2</td><td>Caravaca, Spani</td><td>K-Pg</td><td>66</td><td>Biotic responses</td><td>Sosa-Montes de Oca et al., 2026</td></tr><tr><td>3</td><td>Global scale</td><td>OAE2</td><td>94</td><td>Causes</td><td>Zhang et al., 2026c</td></tr><tr><td>4</td><td>Global scale</td><td>OAE2</td><td>94</td><td>Causes</td><td>Cui et al., 2025</td></tr><tr><td>5</td><td>Tarim Basin, China</td><td>OAE2</td><td>94</td><td>Astronomical cycles</td><td>Chen et al., 2026</td></tr><tr><td>6</td><td>Coqen Basin, Tibet, China</td><td>OAE1d</td><td>113-94</td><td>Causes</td><td>Li et al., 2026b</td></tr><tr><td>7</td><td>Carbonero, Spain</td><td>OAE1a</td><td>120</td><td>Biotic responses</td><td>Sequero et al., 2026</td></tr><tr><td>8</td><td>Araripe Basin, Brazil</td><td>OAE1a, OAE1b</td><td>121-113</td><td>Biotic responses</td><td>Varejão et al., 2025</td></tr><tr><td>9</td><td>Sichuan Basin, China</td><td>T-OAE</td><td>183</td><td>Terrestrial ecosystem</td><td>Luan et al., 2025</td></tr><tr><td>10</td><td>Junggar Basin, China</td><td>T-OAE</td><td>183</td><td>Terrestrial ecosystem</td><td>Zhang et al., 2025</td></tr><tr><td>11</td><td>Northern Iberia, Spain</td><td>T-OAE</td><td>183</td><td>Terrestrial ecosystem</td><td>Rial et al., 2025</td></tr><tr><td>12</td><td>Southern Tibet, China</td><td>late Triassic</td><td>206-201</td><td>Causes</td><td>Du et al., 2025a</td></tr><tr><td>13</td><td>Yangtze Block and South Tibet, China</td><td>CPE</td><td>234-232</td><td>Biotic responses</td><td>Li et al., 2025b</td></tr><tr><td>14</td><td>Jiyuan Basin, North China</td><td>CPE</td><td>234-232</td><td>Terrestrial ecosystem</td><td>Zhang et al., 2026a</td></tr><tr><td>15</td><td>Sverdrup Basin, Canada; Yangtze Block, China</td><td>PT- OAE</td><td>252</td><td>Biotic responses</td><td>Huang et al., 2026</td></tr><tr><td>16</td><td>Yangtze Block, China</td><td>late Permian</td><td>254-252</td><td>Astronomical cycles</td><td>Wei et al., 2025</td></tr><tr><td>17</td><td>Yangtze Block, China</td><td>Kasimovian-Gzhelian</td><td>305-302</td><td>Redox change</td><td>Yang et al., 2026</td></tr><tr><td>18</td><td>Baoshan Block, China</td><td>Late Paleozoic Ice Age</td><td>360-265</td><td>Biotic responses</td><td>Chang et al., 2025</td></tr><tr><td>19</td><td>Appalachian Basin, USA</td><td>Frasnian-Famennian OAE</td><td>372</td><td>Causes</td><td>Algeo et al., 2025</td></tr><tr><td>20</td><td>Appalachian Basin, USA</td><td>Frasnian-Famennian OAE</td><td>372</td><td>Astronomical cycles</td><td>Huygh et al., 2025</td></tr><tr><td>21</td><td>North American Seaway</td><td>Late Devonian</td><td>382-358</td><td>Biotic responses</td><td>Fernandez-Martinez et al., 2026</td></tr><tr><td>22</td><td>Gotland, Sweden</td><td>Mid-Ludfordian OAE</td><td>424</td><td>Biotic responses</td><td>De Backer et al., 2026</td></tr><tr><td>23</td><td>Arctic Canada; western Latvia</td><td>mid-Silurian Mulde Event</td><td>431-426</td><td>Redox change</td><td>Maaleki-Moghadam et al., 2025</td></tr><tr><td>24</td><td>Gotland, Sweden</td><td>Ireviken OAE</td><td>433</td><td>Biotic responses</td><td>Vancoppenolle et al., 2026</td></tr><tr><td>25</td><td>Yangtze Block, China</td><td>Late Ordovician OAE</td><td>448-438</td><td>Biotic responses</td><td>Zhang et al., 2026b</td></tr><tr><td>26</td><td>Yangtze Block, China</td><td>late Cambrian SPICE</td><td>497-494</td><td>Redox change</td><td>Wang et al., 2025b</td></tr><tr><td>27</td><td>Sichuan Basin, China</td><td>middle-late Cambrian</td><td>504-488</td><td>Astronomical cycles</td><td>Wang et al., 2026</td></tr><tr><td>28</td><td>Yangtze Block, China</td><td>Cambrian Stage4</td><td>516-508</td><td>Redox change</td><td>Li et al., 2026a</td></tr><tr><td>29</td><td>Yangtze Block, China</td><td>Ediacaran-Cambrian</td><td>539-514</td><td>Redox change</td><td>Meng et al., 2026</td></tr><tr><td>30</td><td>Yangtze Block, China</td><td>Ediacaran-Cambrian</td><td>640-530</td><td>Biotic responses</td><td>Cao et al., 2025</td></tr></tbody></table>
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