Cretaceous large igneous provinces: from volcanic formation to environmental catastrophes and biological crises

: TheCretaceousPeriodwasmarkedbytheformationofnumerouslargeigneousprovinces(LIPs),sev-eralofwhichwereassociatedwithgeologicallyrapidclimate,environmentalandbiosphereperturbations,includ-ingtheearlyAptianandlatestCenomanianoceanicanoxicevents(OAEs1aand2,respectively).Inmostcases,magmaticCO 2 emissions are thought to have been the major driver of climate and biosphere degradation. This work summarizes the relationships between Cretaceous LIPs and environmental perturbations, focusing on how volcanism caused climate warming during OAE 1a using osmium-isotope and mercury concentration data.ThenewresultssupportmagmaticCO 2 outputfromsubmarineLIPactivityastheprimarytriggerofclimate warming and biosphere stress before / during OAE 1a. This submarine volcanic trigger of OAE 1a (and OAE 2), twoofthemostclimatically / bioticallysevereCretaceousevents,highlightsthecapacityofoceanicLIPstoimpact Earth ’ s environment as profoundly as many continental provinces. Cretaceous magmatism (and likely output of CO 2 andtrace-metalmicronutrients)wasapparentlymostintenseduringthoseOAEs;furtherstudiesareneededto better constrain the eruption histories of those oceanic plateaus. Another open question is why the Cretaceous PeriodoverallfeaturedahigherrateofmagmaticactivityandLIPformationcomparedwithbeforeandafterwards. Supplementary material: Geochemical datasets from Deep Sea Drilling Project (DSDP) Site398 andAlstätte-1 and a summary of Cretaceous Os isotope literature are available at https: // doi.org / 10.6084 / m9. ﬁ gshare.c. 7026011

Large igneous provinces (LIPs) represent the formation of huge volumes of igneous material emplaced into and/or onto the Earth's continental or oceanic crust over a geologically short time interval (Coffin and Eldholm 1994).This emplacement occurs through a combination of both intrusive magmatism and extrusive volcanic activity: the latter is most famously in the form of continental flood basalts, some of which featured individual lava flows with an aerial extent on the order of 10 4 km 2 (e.g.Coffin and Eldholm 1994;Bryan and Ernst 2008;Bryan and Ferrari 2013;Ernst 2014).Several different formal definitions for large igneous provinces have been proposed, characterizing them as magmatic provinces with high volumes of igneous material that can form in a range of intraplate tectonic settings, with the majority of magmas emplaced over a geologically short space of time (e.g..0.1 Mkm 3 in pulses of c. 1-5 Myr; Bryan and Ernst 2008; or .0.1 Mkm 3 , frequently .1 Mkm 3 , in ,5 Myr, often ,2 Myr; Ernst et al. 2021).
The history of LIP formation over the last 300 Myr, since the formation of the Pangaean supercontinent, is relatively well established.Several provinces are preserved as flood basalt units and/or large-scale intrusive sill and dyke swarms on the continents, or as oceanic plateaus in the Pacific and Indian ocean basins (see e.g.Coffin and Eldholm 1994;Ernst 2014;Kerr 2014;Ernst et al. 2021).These LIPs are generally better preserved than those that formed prior to 300 Ma, and are typically estimated to have had original volumes of emplaced magma greater than 0.1 Mkm 3 ; frequently above 1 Mkm 3 (Bryan and Ernst 2008;Ernst et al. 2021).However, the Greater Ontong-Java Plateau in the western Pacific Ocean may have been an order of magnitude more voluminous still (Gladczenko et al. 1997;Taylor 2006;Hoernle et al. 2010).The rapid emplacement of most provinces (within c. 1 Myr; see Kasbohm et al. 2021, and references therein) highlights that the rate of magma production during LIP formation was typically both higher and sustained for longer time periods than for any observed volcanic activity in human history.In this context, it is notable that the large majority of geologically rapid changes to Earth's global climate and environment during the Phanerozoic Aeon, particularly in the last 300 Myr, broadly coincided with an interval of LIP formation.These events include at least four of the so called 'Big Five' mass extinctions (e.g.Wignall 2001;Courtillot and Renne 2003;Bond and Wignall 2014;Ernst et al. 2021;Kasbohm et al. 2021).
The Cretaceous Period (143-66 Ma; Gale et al. 2020) was marked by the highest rate of LIP formation in at least the last 300 Myr.Numerous provinces were emplaced into both the continental and oceanic crust, and estimates of the total number of Cretaceous LIPs vary (but see recent lists by e.g.Torsvik 2019;Ernst et al. 2021).However, there are seven major Cretaceous LIPs that comprise both high volumes of igneous material and have been widely linked with climate/environmental change and/or elevated biotic stress/extinction during that period (Fig. 1).Continental flood-basalt provinces include the Paraná-Etendeka LIP, the High Arctic LIP (HALIP), the Madagascan LIP, and the Deccan Traps (although the latter three also comprise offshore components).Additionally, the Greater Ontong-Java, Kerguelen, and Caribbean plateaus were emplaced into continental margins and ocean crust to form oceanic plateaus, which are considered to represent large areas of elevated and thickened basaltic ocean floor formed through mantle-plume activity rather than seafloor spreading-related magmatic processes (Kerr 2014).Indeed, owing to the subduction of most pre-Jurassic crust, many currently known oceanic LIPs are Cretaceous in age, including those that are best preserved and most extensively investigated (see Kerr 2014).This work reviews the seven major Cretaceous LIPs that are thought to have impacted Earth's environment and/or biosphere, and the relationship between these phenomena, utilizing both new datasets for one such environmental episode, the Early Aptian oceanic anoxic event (OAE 1a), and published datasets relating to this and other Cretaceous intervals and LIPs.

Cretaceous LIPs and episodes of environmental change
The Paraná-Etendeka LIP and Valanginian 'Weissert' Event The Paraná-Etendeka LIP covers an area of c. 4 Mkm 2 in western South America (principally Brazil, but also Paraquay, Uruguay and Argentina) and western Africa (Namibia and Angola), and has long been associated with the opening of the South Atlantic Ocean (Peate 1997).Extrusive magmas largely comprise tholeiitic basalt flows, with some silicic and alkaline units and intrusive sills and dyke swarms, particularly in the South American Paraná part of the province (e.g.Erlank et al. 1984;Milner et al. 1992Milner et al. , 1995;;Peate 1997;Marsh et al. 2001;De Min et al. 2018; see also Gomes and Vasconcelos 2021, and references therein).The volatile budget of Paraná-Etendeka magmas may have been relatively low compared with other LIPs, as calculated from phenocrysts and measured from melt inclusions (maximum of 900-1100 ppm S, 125 ppm Cl and 450 ppm F; Callegaro et al. 2014;Marks et al. 2014).Moreover, the magmas were emplaced within or erupted on to country rocks that largely consisted of aeolian sandstones, organic-lean shales and crystalline basement (Barreto et al. 2016;Jones et al. 2016).These lithologies were all probably volatile depleted and would not have been a major source of thermogenic carbon or sulfur following heating by the intruding magmas (cf., Svensen et al. 2004Svensen et al. , 2009;;Heimdal et al. 2018).
Paraná-Etendeka volcanism is widely attributed to have caused an episode of prolonged environmental perturbation during the Valanginian , known as the Weissert Event (Erba et al. 2004).This episode of environmental change is characterized in the upper Valanginian stratigraphic record by a positive carbon-isotope (δ 13 C) excursion of ∼1.5-2 ‰ in carbonate and up to 4 ‰ in bulk organic matter (e.g.Lini et al. 1992;Weissert et al. 1998;Erba et al. 2004;Price and Mutterlose 2004;Gröcke et al. 2005;McArthur et al. 2007;Bornemann and Mutterlose 2008;Littler et al. 2011;Price et al. 2018;Jelby et al. 2020).The end of the Weissert Event is marked by the maximum value of the positive carbon-isotope excursion in upper Valanginian strata (Erba et al. 2004), although the δ 13 C values decrease gradually through lower Hauterivian strata.This δ 13 C shift was initially hypothesized to result from enhanced burial of organic carbon (which is isotopically light) in sediments as they were deposited, leaving the residual seawater carbon inventory, and any carbonates or organic material subsequently formed from it, isotopically heavier (Weissert et al. 1998;Erba et al. 2004).However, while some Valanginian strata do preserve organic-rich facies consistent with this hypothesis, the majority do not.Thus,  (2015); and the KPg from Claeys et al. (2002).
enhanced primary productivity and burial of organic carbon in terrestrial sediments have been proposed as alternative contributors towards the carbon-cycle change (Westermann et al. 2010).
The nature of climate change during the Weissert Event is debated.The Early Valanginian interval prior to and during the onset of the Weissert Event was marked by a warm global climate (Littler et al. 2011), and evidence for surface warming at that time has been documented from bulk-rock oxygenisotope (δ 18 O) records in some NW Tethyan and proto-Atlantic archives (Duchamp-Alphonse et al. 2007;Charbonnier et al. 2020c).Clay mineralogy and calcareous nannofossil and spore-pollen assemblages also indicate enhanced humidity in and around the former region (Kujau et al. 2013;Möller et al. 2020;Charbonnier et al. 2020c).Thus far, this warming has not been as clearly recorded by other palaeotemperature proxies or robustly documented in other areas around the world, with only modest temperature increases of at most 1°C that largely do not significantly exceed pre-event variations (e.g.Littler et al. 2011;Price et al. 2018;Charbonnier et al. 2020c;Cavalheiro et al. 2021).In contrast, in several Valanginian records, particularly those from the Boreal Realm but also in the NW Tethys and southern high latitudes, the turning point in δ 13 C values at the peak of the positive excursion stratigraphically correlates with evidence for the onset of a transient fall in surface temperatures, (e.g.Erba et al. 2004;McArthur et al. 2007;Meissner et al. 2015;Price et al. 2018;Cavalheiro et al. 2021).This correlation between the δ 13 C peak and onset of temperature decrease is consistent with enhanced atmospheric CO 2 sequestration through the hypothesized increase in organic-carbon burial.However, proto-Atlantic sea-surface temperature reconstructions based on the biomarker-based TEX 86 palaeothermometer suggest that stable warm conditions persisted throughout the Valanginian in that low-latitude environment (Littler et al. 2011), potentially highlighting a steepening of global temperature gradients during the Weissert Event (Charbonnier et al. 2020c;Cavalheiro et al. 2021).
Age constraints for Paraná-Etendeka volcanism are based on magnetostratigraphy and both argonargon ( 40 Ar/ 39 Ar) and uranium-lead (U-Pb) geochronology.Collectively, these data suggest a duration of at least 2-3 Myr of effusive volcanic activity for the LIP as a whole, with Etendeka magnetostratigraphy suggesting even more protracted eruptions in that part of the province (see compilations by Mena et al. 2011;Dodd et al. 2015;Gomes and Vasconcelos 2021;Bacha et al. 2022;Jiang et al. 2023).Uranium-lead and 40 Ar/ 39 Ar dating of Paraná and Etendeka units by several studies have outlined a pronounced pulse of volcanic activity that probably reached its maximum between 134 and 135 Ma (Renne et al. 1992(Renne et al. , 1996; following recalculation by Thiede and Vasconcelos 2010;also Ernesto et al. 1999;Janasi et al. 2011;Pinto et al. 2011;Florisbal et al. 2014;Almeida et al. 2018;Gomes and Vasconcelos 2021;Bacha et al. 2022).These magmatic ages broadly overlap with the onset date of the Weissert Event based on both magnetostratigraphic (133.9Ma;Cavalheiro et al. 2021) and cyclostratigraphic (134.56 Ma;Martinez et al. 2023) age modelling.However, recent highprecision U-Pb dating of Paraná rocks has shown that at least part of the intense volcanic activity postdated the start of the Weissert Event (Rocha et al. 2020).To what extent the older v. younger dates of Paraná-Etendeka magmas are representative of rates of extrusive volcanic activity across the province as a whole is unclear.Because the majority of dated rocks are coeval with the onset of the Weissert Event, it is widely assumed that there was a causal relationship between the two phenomena (e.g.Erba et al. 2004;Martinez et al. 2015Martinez et al. , 2023;;Gomes and Vasconcelos 2021;Bacha et al. 2022), although it cannot be excluded that significant Paraná-Etendeka volcanism occurred later, potentially questioning the nature of any link (Rocha et al. 2020).

The Greater Ontong-Java Plateau and Early Aptian OAE (OAE 1a)
The total magmatic volume comprised by the Greater Ontong-Java Plateau (G-OJP) is debated, but is often defined as the Ontong-Java Plateau together with subsidiary nearby flood-basalt provinces such as the Nauru, East Mariana, Manihiki and Hikurangi (after Ingle and Coffin 2004;Charbonnier and Föllmi 2017; see also Svensen et al. 2019, and references therein).These provinces were collectively emplaced on to the western Pacific oceanic crust during the Early Cretaceous, covering an area over 2 Mkm 2 , and comprising a combined total volume of several tens of millions of cubic kilometres, significantly greater than any other preserved LIP (Taylor 2006;Hoernle et al. 2010).Moreover, part of the original plateau has likely been subducted since its emplacement (see Schlanger et al. 1981;Larson 1991).Accreted and exposed Ontong-Java magmas in Malaita (Solomon Islands) consist of massive sheeted and pillowed basalt flows, with basaltic rocks also recovered from ODP drill sites on the plateau (e.g.Saunders et al. 1996;Tejada et al. 1996;Neal et al. 1997;Fitton and Godard 2004).It is likely that much of the G-OJP was emplaced via submarine volcanic activity (Saunders et al. 1996).However, preservation of phreatomagmatic deposits at Ocean Drill Program (ODP) Site 1184 in the Eastern Salient of the Ontong-Java Plateau shows that at least some eruptions occurred near or above the sea surface (Chambers et al. 2004;Thordarson 2004).
Geochronological 40 Ar/ 39 Ar studies on Ontong-Java basalts and volcaniclastic sediments from southern Malaitia, Ramos Island and ODP sites 289, 807 and 1184 indicate a pulse of volcanic activity that covered a huge area around 122 Ma (Mahoney et al. 1993;Tejada et al. 1996Tejada et al. , 2002;;Chambers et al. 2004).Submarine eruptions of a broadly similar age have also been reported from studies of whole rock basalts recovered from the Manihiki and Hikurangi plateaus (Ingle et al. 2007;Hoernle et al. 2010;Timm et al. 2011).Recently, this geochronology has been challenged by Davidson et al. (2023), on the basis that older ages might be affected by recoil, while plagioclase 40 Ar/ 39 Ar dates from Ontong-Java and Manihiki plateau rocks have ages closer to 110 Ma than 120 Ma.However, those authors do not discount the possibility that the G-OJP had a protracted emplacement over several million years across the late Barremian to early Albian interval, in which their dates are from volcanic rocks that were erupted towards the end of the plateau's formation.Mahoney et al. (1993) and Tejada et al. (1996) also report eruption ages of c. 90 Ma at Sigana Island and ODP Site 803, potentially highlighting a second major pulse of volcanism on the G-OJP.Plateau formation through two distinct spells of volcanism 30 Myr apart is at odds with the mantle plume model of LIP emplacement, and it has been suggested that the younger dates are also affected by argon recoil, and may represent minimum ages (Chambers et al. 2002).
The 122 Ma pulse of volcanism has long been associated with OAE 1a (121 Ma), based on the broad temporal correlation between the two phenomena (e.g.Larson and Erba 1999;Courtillot and Renne 2003;Erba et al. 2015).The Early Aptian oceanic anoxic event was marked by the development of marine anoxia/euxinia across the global ocean for approximately 1-1.4 Myr (Li et al. 2008;Malinverno et al. 2010;Leandro et al. 2022).Widespread oxygendepleted conditions were initially interpreted from the preservation of laminated, organic-rich, mudstones around the world, dubbed the Selli Level in the Umbria-Marche Basin of central Italy (e.g.Schlanger and Jenkyns 1976;Weissert 1989;Jenkyns 1995;Pancost et al. 2004;Föllmi et al. 2006;van Breugel et al. 2007; see also reviews by Jenkyns 2010 andRobinson et al. 2017).Global stratigraphic records of OAE 1a are further characterized by a series of δ 13 C excursions, reflecting sequential carbon-cycle perturbations during the Early Aptian (Weissert 1989;Jenkyns 1995;Menegatti et al. 1998;Gröcke et al. 1999;Ando et al. 2008;Robinson et al. 2008;Vickers et al. 2016).These δ 13 C excursions have been subdivided into segments (C3-C7) based on Tethyan archives (Menegatti et al. 1998).Following a relatively stable δ 13 C signature through uppermost Barremian-lowermost Aptian strata (C1-C2), a sharp negative shift is documented at the base of the OAE 1a level (C3), highlighting a large influx of isotopically light carbon to the Earth's surface from one or more of volcanic activity, thermogenic emissions, and methane clathrate destabilization (e.g.Jahren et al. 2001;Méhay et al. 2009;Kuhnt et al. 2011;Naafs et al. 2016;Bauer et al. 2017;Adloff et al. 2020).The negative isotopic signature is followed by a positive rebound (C4), another stable spell (C5), a second positive excursion at the top of the OAE level (C6) and a continuation of elevated δ 13 C values that remain high above it (C7), with these later shifts probably reflecting the widespread deposition of organic carbon (Jenkyns 2010).The various carbon-cycle perturbations impacted Earth's surface climate during OAE 1a, with overarching climate warming and transient interludes of cooling during the event interpreted from palaeontological and geochemical evidence (e.g.Menegatti et al. 1998;Jenkyns 2003Jenkyns , 2018;;Dumitrescu et al. 2006;Ando et al. 2008;Kuhnt et al. 2011;Bottini et al. 2015;Naafs and Pancost 2016).As well as oceanic anoxia and global temperature changes, OAE 1a is thought to have been marked by seawater acidification, biospheric stress, accelerated hydrological cycling and enhanced continental weathering (e.g.Erba 2004;Erba et al. 2010Erba et al. , 2015;;Bottini et al. 2012;Mutterlose et al. 2014;Lechler et al. 2015;Naafs and Pancost 2016).

The Kerguelen Plateau and Aptian-Albian OAE (OAE 1b)
The Kerguelen Plateau began to form during the break-up of India from Australia and Antarctica during the Early Cretaceous, and today encompasses an area of elevated oceanic crust comprising .2Mkm 3 of mainly basaltic rock, with some dacites and rhyolites (Kerr 2014).Additional fragments of the plateau such as Broken Ridge and Ninetyeast Ridge are preserved in the Indian Ocean (see Wallace et al. 2002, and references therein).Some older magmas that formed during the initial break-up of southern Gondwana are also preserved in NE India and western Australia (Frey et al. 1996;Coffin et al. 2002;Kent et al. 2002).Unlike many LIPs, the Kerguelen Plateau has a long volcanic history, commencing in the Early Cretaceous and continuing (less voluminously) to the present day (see Coffin et al. 2002;Jiang et al. 2021).The onshore volcanics all date to 114 Ma or older; Baksi 1995;Frey et al. 1996;Coffin et al. 2002;Kent et al. 2002).The oldest part of the main oceanic province is the Southern Kerguelen Plateau, for which 40 Ar/ 39 Ar dates suggest that volcanic activity commenced by 125 Ma at least, and possibly earlier (Jiang et al. 2022).Volcanic activity on the Southern Kerguelen Plateau continued until 110 Ma (Whitechurch et al. 1992;Coffin et al. 2002;Duncan 2002;Jiang et al. 2022), with later eruptions occurring on the Central Kerguelen Plateau and Elan Bank (100.4+ 0.7 and 107.7 Ma, respectively;Duncan 2002).Further north, Broken Ridge samples have been dated to c. 95 Ma (Duncan 2002), while ages for rocks from Ninetyeast Ridge and Skiff Bank are younger still (c. 85-35 Ma;Duncan 1978Duncan , 1991Duncan , 2002)).Jiang et al. (2022Jiang et al. ( , 2023) ) have recently proposed a revision of the temporal magmatic history of Kerguelen, and argue that the volcanism commenced at 125 Ma or earlier.As such, these authors suggest that early volcanic activity on the plateau may have contributed towards the environmental change associated with OAE 1a, while acknowledging that Kerguelen carbon emissions were unlikely to have caused the OAE alone.More frequently, a link has been postulated between the Kerguelen Plateau and more moderate environmental change during the Aptian-Albian transition (c.113 Ma), often referred to as OAE 1b, based on a better age correlation between the two phenomena (e.g.Trabucho Alexandre et al. 2011;Erba et al. 2015;Sabatino et al. 2018;Matsumoto et al. 2020).This interval of marine anoxia is documented in the Vocontian Basin (SE France) and NW Tethys by the preservation of four main discrete organic-rich shale intervals (the Jacob, Kilian, Paquier or Urbino level, and Leenhardt levels; Bréhéret 1988;Coccioni et al. 2014).Some of these shale horizons have also been identified in Atlantic sites (e.g.Erbacher et al. 2001;Herrle et al. 2004;Trabucho Alexandre et al. 2011).Of those four shale levels, the Kilian and Paquier horizons are the two that have been the most widely identified around the world (see overviews in Coccioni et al. (2014); Bodin et al. 2023), with OAE 1b sometimes referred to as the Paquier Event (Jenkyns 2010).Upper Aptian-lower Albian strata also document a series of δ 13 C excursions, with the Kilian and Paquier levels both marked by sharp negative carbon-isotope shifts of 2-3‰ that have been used to identify OAE 1b strata at sites around the world not marked by deposition of organic-rich sediments (e.g., Gröcke et al. 1999;Millán et al. 2014;Tsikos et al. 2004b;Herrle et al. 2015;Navarro-Ramirez et al. 2015;Phelps et al. 2015;Li et al. 2016;Zhao et al. 2022).Upper Aptian palaeotemperature records from the North Atlantic indicate a prolonged cooling pulse prior to/during the onset of OAE 1b (McAnena et al. 2013), and a cold pulse has also been postulated from study of Vocontian Basin and NW Tethys Ocean sites (Herrle and Mutterlose 2003;Bottini and Erba 2018).However, there is no clear evidence for a similar fall in temperatures in the South Atlantic (Jenkyns et al. 2012).In contrast, the Paquier level (or its stratigraphic equivalent) is marked by evidence for climate warming in the Vocontian Basin, North Atlantic and NW Tethys (Erbacher et al. 2001;Herrle et al. 2003;Huber et al. 2011;Bottini and Erba 2018).
The Caribbean LIP and the latest Cenomanian OAE (OAE 2) The Caribbean (or Caribbean-Colombian) LIP was an oceanic plateau that is preserved today as obducted fragments outcropping in Ecuador, Colombia (including Isla Gorgona), Costa Rica, and several Caribbean islands, most prominently Haiti, Jamaica, Curaçao, and Aruba (Kerr et al. 1997).The plateau also comprises a substantial proportion of the thickened oceanic crust of the Caribbean (Mauffret and Leroy 1997).The fragmentary nature of the remnants of this LIP hinders determination of its original volume, although 4-4.5 Mkm 3 has been estimated (see Courtillot and Renne 2003;Kerr et al. 2003;Kuroda et al. 2007).It is widely accepted that the plateau formed on the Farallon (proto-Pacific) plate and possibly represents the 'head' phase of the Galápagos plume (Kerr et al. 2003;Boschman et al. 2014;Nerlich et al. 2014).After its formation, the southern part of the plateau collided with and accreted onto the NW margin of South America, while the northern portion moved into the inter-American gap that had been opening between north and south America since the Jurassic (see reviews in Kerr et al. 2003;Boschman et al. 2014).Studies of the exposed outcrop and recovered rocks from ODP sites indicate that the emplaced magmas largely consist of tholeiitic pillow basalts, with some basaltic and dolerite sills in addition to picritic and komatiitic lavas also preserved (Kerr et al. 2003).While it is evident that the plateau predominantly formed through submarine volcanism, Buchs et al. (2018) reported layered tuffs with accretionary lapilli and interbedded lahar deposits with rounded clasts of basalt in accreted oceanic plateau sequences from western Colombia.These tuffs indicate that, like other Cretaceous Pacific plateaus, the Caribbean LIP became subaerial.The preservation of corals and carbonized tree-trunk fragments in sediments interbedded with LIP basalts in the Western Cordillera of Colombia supports the occurrence of eruptions in either shallow water depths or a subaerial context (Hall et al. 1972;Moreno-Sanchez and Pardo-Trujillo 2003).Dating of Caribbean LIP samples through 40 Ar/ 39 Ar and Re-Os geochronology has highlighted a wide range of ages spanning the Cretaceous to the Early Paleogene (see Kerr et al. 2003).However, 40 Ar/ 39 Ar ages from Deep Sea Drilling Project (DSDP) Leg 15, Haiti, Gorgona Island, Costa Rica, Hispaniola, Western Colombia and Curaçao all suggest a major pulse of volcanism between 93 and 88 Ma (Walker et al. 1991(Walker et al. , 1999;;Sinton et al. 1998;Kerr et al. 2004;Snow et al. 2005; see also Kasbohm et al. 2021).
The Cenomanian-Turonian boundary has been dated to 93.9 Ma (Meyers et al. 2012), and records another interval of widespread oceanic anoxia (OAE 2), which has been widely linked to the coeval Caribbean volcanism (e.g.Sinton and Duncan 1997;Kerr 1998;Snow et al. 2005;Turgeon and Creaser 2008;Du Vivier et al. 2014, 2015;Scaife et al. 2017;Percival et al. 2018).Stratigraphic archives of this episode of global environmental perturbation are characterized by a positive δ 13 C excursion of up to 6‰ in all of carbonates, bulk-, and compoundspecific organic matter (e.g.Scholle and Arthur 1980;Hasegawa 1997;Tsikos et al. 2004a;Erbacher et al. 2005;Sageman et al. 2006;Jarvis et al. 2011).Well-preserved laminated organic-rich shales have been reported from uppermost Cenomanian-lowermost Turonian strata in numerous sites, particularly in the Atlantic, NW Tethys and Boreal Realm (Schlanger and Jenkyns 1976;Arthur et al. 1987;Linnert et al. 2010; see also reviews by Jenkyns 2010; Robinson et al. 2017).Thus, this isotopic shift is typically interpreted as reflecting enhanced organic-carbon burial in the global ocean.Palaeotemperature reconstructions based on carbonate oxygen-isotope (and in some locations, TEX 86 ) trends indicate a rise in temperatures at the onset of OAE 2, with both proxies highlighting that the warm conditions continued after the end of the event (Paul et al. 1999;Forster et al. 2007;Jarvis et al. 2011;van Helmond et al. 2014van Helmond et al. , 2015;;O'Connor et al. 2020).
However, the elevated temperatures were punctuated by a transient interval of climate cooling (and seawater re-oxygenation) across large parts of the marine realm, dubbed the Plenus Cold Event owing to its initial recognition by the southward migration of the boreal belemnite species Praeactinocamax plenus (Gale and Christensen 1996).This cooling/reoxygenation pulse is further documented by multiple geochemical palaeothermometers, a shift in the abundance of specific nannofossils (notably cold-water species), and often a return to pre-OAE lithologies and the disappearance of organic-rich shales (e.g.Tsikos et al. 2004a;Forster et al. 2007;Sinninghe-Damsté et al. 2008;Linnert et al. 2010;van Helmond et al. 2014van Helmond et al. , 2015van Helmond et al. , 2016;;Desmares et al. 2016;O'Connor et al. 2020).The transient temperature decrease probably resulted from sequestration of organic carbon caused by one or both of the documented organic-carbon burial and enhanced silicate weathering (Pogge von Strandmann et al. 2013;Robinson et al. 2019;Percival et al. 2020;Papadomanolaki et al. 2022).However, this fall in temperatures did not occur synchronously across the world, suggesting that local oceanographic and climatic conditions significantly influenced the regional manifestation of any cooling (O'Connor et al. 2020;Percival et al. 2020).

The High Arctic LIP (HALIP) and its relation to the Cretaceous OAEs
Numerous studies have documented the preservation of Cretaceous magmatic units across much of the High Arctic, spanning Svalbard, Franz Josef Land, Novaya Zemlya, the Barents Sea, Northern Greenland, the Canadian Arctic islands, and New Siberian Islands, together with the offshore Alpha-Mendeleev Ridge and Chukchi Plateau (see e.g.Tarduno 1998;Maher 2001;Buchan and Ernst 2006;Tegner et al. 2011;Corfu et al. 2013;Senger et al. 2014;Naber et al. 2021;Bédard et al. 2021a, b;Senger and Galland 2022).Comprising both extrusive volcanic units and intrusive dyke swarms in particular, these magmatic formations have been postulated to collectively represent the HALIP (Tarduno 1998).Thus far, geochronology studies of the HALIP have primarily focused on areas in Arctic Canada (particularly the islands of Axel Heiberg and Ellesmere), with U-Pb and 40 Ar/ 39 Ar dates indicating that the LIP formed through numerous spells of emplacement over a 40 Myr interval of the Cretaceous (Tegner et al. 2011;Dockman et al. 2018).In particular, two major magmatic pulses at 135-120 Ma and 105-90 Ma have been identified (e.g.Corfu et al. 2013;Evenchick et al. 2015;Estrada et al. 2016;Polteau et al. 2016;Dockman et al. 2018;Kingsbury et al. 2018;Naber et al. 2021;Deegan et al. 2022).Dating of alkaline HALIP rocks in Northern Greenland potentially indicates a later episode of HALIP-related alkaline volcanism between 85 and 70 Ma (Tegner et al. 2011;Thórarinsson et al. 2015).Recognized magmatic pulses of HALIP overlap in age with both OAEs 1a and 2; consequently, the province has been linked with both episodes of environmental change (e.g.Tegner et al. 2011;Zheng et al. 2013;Polteau et al. 2016;Adloff et al. 2020).In the case of OAE 1a, this link may be related to the emission of thermogenic carbon following intrusion of organic-rich sedimentary rocks by basaltic sills, providing a source of isotopically light carbon to the Earth's surface that could have caused the C3 negative excursion in δ 13 C (Polteau et al. 2016;Adloff et al. 2020;Deegan et al. 2022).

The Madagascan LIP
The Madagascan LIP largely comprises tholeiitic flood basalt sequences, with minor alkaline and silicic units, and associated sill and dyke swarms and intrusive complexes, and has been attributed to the Marion plume prior to the separation of Madagascar from the Indian subcontinent (e.g.Storey et al. 1995;Torsvik et al. 1998Torsvik et al. , 2000;;Kumar et al. 2001).Remnants of the province are well exposed and preserved along Mahajanga and the eastern coastal area, and in the Morondava sedimentary basins in western Madagascar (Mahoney et al. 1991;Storey et al. 1995;Melluso et al. 2001;Cucciniello et al. 2022).Further volcanism is also recorded through intrusions into and extrusions onto Precambrian basement.The original volume of the province is poorly constrained, but probably exceeded 1 Mkm 3 , including the Madagascar Plateau flanking the south coast of the island, as well as the Conrad Rise (Storey et al. 1995).Age estimates for Madagascar LIP volcanism largely range between c. 93 and 86 Ma, based largely on 40 Ar/ 39 Ar and some U-Pb dating, together with biostratigraphic constraints based on interbedded sedimentary rocks (Storey et al. 1995;Torsvik et al. 2000;Melluso et al. 2001Melluso et al. , 2005;;Pande et al. 2001;Mahoney et al. 2008;Cucciniello et al. 2010Cucciniello et al. , 2011Cucciniello et al. , 2013Cucciniello et al. , 2022; see also Jiang et al. 2023).Because these ages are close to that of the Cenomanian-Turonian boundary, it has been postulated that the Madagascan LIP might have contributed to environmental change associated with OAE 2. However, statistical appraisals of the available ages shows that volcanic activity initially peaked in northern Madagascar and spanned 3 Myr between 93 and 90 Ma (Cucciniello et al. 2022), slightly postdating OAE 2. Thus, it is unlikely that the Madagascan LIP played a major role in triggering environmental change associated with that OAE.

The Deccan-Traps and the Cretaceous-Paleogene interval
Major eruptions associated with the Deccan Traps (western India) commenced during the latest Maastrichtian, 300-400 kyr prior to the end of the Cretaceous Period (66.04 Ma;Gale et al. 2020) and continued into the earliest Paleogene (Schoene et al. 2019;Sprain et al. 2019 and references therein).It has been suggested that a smaller-scale phase of Deccan eruptions occurred earlier in the Cretaceous (68-67 Ma; see Chenet et al. 2007;Parisio et al. 2016).However, subsequent data have cast doubt on the occurrence of major eruptions significantly prior to the main-phase of Deccan volcanism (Schoene et al. 2015).By far the most intensely studied part of the Deccan Traps is the Western Ghats region, which comprises a cumulative thickness of c. 3000 m of tholeiitic basalts in multiple flood basalt units stratigraphically separated by oxidized 'redbole' palaeosols (Widdowson et al. 1997, and references therein).Magnetostratigraphic studies of the Western Ghats basalts show that the main phase of Deccan volcanism began just prior to the start of the C30n-C29r chron reversal (66.38 Ma), and concluded just after the C29r-C29n reversal at 65.7 Ma (Courtillot et al. 1986(Courtillot et al. , 2000;;Chenet et al. 2009).This ,1 Myr duration is supported by numerous recent geochronological investigations utilizing both 40 Ar/ 39 Ar ages of Deccan basalts (66.31-65.72Ma; e.g.Renne et al. 2015;Sprain et al. 2019) and U-Pb dating of zircons preserved in the red boles interbedded between with them (66.26-65.63Ma; e.g.Schoene et al. 2015Schoene et al. , 2019;;Eddy et al. 2020).Recent modelling of eruption rates in the Western Ghats, based on 40 Ar/ 39 Ar geochronology, highlighted a pronounced rise in eruption volume between the Bushe and Poladpur formations, around the time of the Cretaceous-Paleogene (KPg) transition and Chicxulub impact (Renne et al. 2015;Sprain et al. 2019).An acceleration in volcanism beginning with Poladpur eruptions is also shown by U-Pb-based modelling, but manifesting as a series of discrete pulses that began 50-100 kyr prior to the KPg event and bolide impact (Schoene et al. 2019).However, U-Pb geochronology of red-bole zircons from the Malwa plateau indicate that there were highly voluminous Deccan eruptions occurring in the northern part of the province up to 350 kyr prior to the KPg event (Eddy et al. 2020).Additionally, some of the Deccan magmas are preserved offshore from the Indian coast and may comprise a volume significantly greater than the well-studied onshore volcanics, with little known regarding their eruptive history (Mittal et al. 2022).
The last 300-400 kyr of the Cretaceous were marked by up to 2-3°C of climate warming (the Late Maastrichtian Warming Event), which has been documented in shallow and deep marine settings around the world by both oxygen-isotope and TEX 86 data (e.g.Li and Keller 1998a;Westerhold et al. 2011;Esmeray-Senlet et al. 2015;Vellekoop et al. 2016;Barnet et al. 2018;Woelders et al. 2017Woelders et al. , 2018;;Hull et al. 2020).Cyclostratigraphic and magnetostratigraphic age models across a number of these sites demonstrate that the warming commenced coevally with the onset of Deccan volcanism, supporting a causal link between the two phenomena (Barnet et al. 2018;Woelders et al. 2017;Hull et al. 2020;Gilabert et al. 2022).Intriguingly, temperatures then declined around 100-200 kyr before the end of the Cretaceous, despite continuing Deccan activity (Vellekoop et al. 2016;Barnet et al. 2018;Woelders et al. 2017Woelders et al. , 2018;;Hull et al. 2020).A rise in SO 2 contents of Deccan basalts from 750 to 1800 ppm, and thus of assumed volatile emissions, just prior to the end of the Cretaceous may have aided this cooling pulse (Callegaro et al. 2023).A second (geologically brief ) warming pulse in the final 10 s kyr of the period has been postulated, although evidence for this second temperature increase has been reported from only a few locations (Stüben et al. 2003;Keller et al. 2020;O'Connor et al. 2023;cf., Vellekoop et al. 2016;Woelders et al. 2017;Barnet et al. 2018;Hull et al. 2020).It is assumed that these temperature changes occurred in response to carbon cycle disturbances following Deccan CO 2 emissions (e.g.Li and Keller 1998a;Robinson et al. 2009;Esmeray-Senlet et al. 2015;Vellekoop et al. 2016;Hull et al. 2020).However, there is no clear δ 13 C excursion in upper Maastrichtian strata comparable in magnitude with those of the Cretaceous OAEs (Woelders et al. 2017;Barnet et al. 2018).The lack of a negative isotopic shift may result from Deccan carbon output being primarily magmatic in origin, rather than thermogenic, although whether magmatic volatiles alone could have caused the warming is unclear (Self et al. 2006; see also Ganino and Arndt 2009).
The KPg boundary, which marks the most recent of the 'Big Five' Phanerozoic mass extinctions (Raup and Sepkoski 1982), also overlapped in time with Deccan volcanism.This extinction has been compellingly linked to the coeval Chicxulub impact by global enrichment of iridium and other platinum group elements in the extinction horizon, together with documented microspherules, shocked-quartz grains and tsunami sediments (e.g.Alvarez et al. 1980;Schulte et al. 2010;Goderis et al. 2021).Nonetheless, it has been hypothesized that Deccan eruptions contributed towards causing this biospheric crisis, either in combination with the Chicxulub impact or independently (e.g.Courtillot and Renne 2003;Keller 2012;Font et al. 2016;Petersen et al. 2016;Callegaro et al. 2023;Cox and Keller 2023).The different hypotheses proposed regarding the links between the Chicxulub impact, Deccan Traps and the extinction event are outlined in more detail elsewhere (see e.g.Schulte et al. 2010;Hart et al. 2012Hart et al. , 2019;;Richards et al. 2015;Punekar et al. 2016;Leighton et al. 2017;Schoene et al. 2019;Hull et al. 2020;Callegaro et al. 2023;Cox and Keller 2023;Senel et al. 2023), and are not further discussed here.

The link between LIP volcanism and environmental change
Many major environmental perturbations during the Cretaceous were marked by pronounced climate warming on a global scale (with the Weissert Event and parts of OAE 1b the likely exceptions), and LIP-related emissions of carbon dioxide are widely accepted to have been a key trigger of these temperature rises (e.g.Weissert and Erba 2004;Jenkyns 2010;Bond and Wignall 2014;Bodin et al. 2015).Carbon outgassing of LIP basalts remains poorly constrained, however, and it is uncertain whether magmatic emissions would have been sufficient to greatly impact atmospheric CO 2 levels (see Self et al. 2006Self et al. , 2014;;Black et al. 2021).As noted above, thermogenic carbon emissions associated with the heating of organic-rich sediments by intruding magmas have been proposed as an additional/ alternative trigger of climate warming during OAE 1a (Polteau et al. 2016;Deegan et al. 2022).This mechanism has also been suggested for the Late Maastrichtian Warming Event (Eddy et al. 2020;Hernandez Nava et al. 2021), although the limited evidence for organic-rich sediments that were intruded by Deccan magmas and the lack of a δ 13 C negative excursion in uppermost Cretaceous strata do not support this hypothesis.Regardless of the specific carbon source, any episode of Cretaceous climate warming probably resulted in acceleration of the global hydrological cycle, enhancing continental weathering and nutrient runoff to the oceans, and consequently stimulating eutrophication and seawater oxygen depletion, aided by ocean stratification and stagnation directly caused by oceanic temperature increase (see Jenkyns 2010; Bond and Sun 2021, and references therein).The carbon emissions themselves could also have directly resulted in ocean acidification and associated environmental stress, at least during some events (Erba et al. 2010).
Establishing better constraints on the timing of volcanic activity from sedimentary records that document surface temperature fluctuations and/or changes in environmental and biospheric conditions is key to resolving the complex relationships between LIP emplacement and their impact on Earth's climate and environment.In recent years, the development of several geochemical proxies of volcanism has enabled reconstruction of the timing of LIP activity in the stratigraphic archives that record episodes of environmental and/or biota perturbation.Osmiumisotope compositions (specifically 187 Os/ 188 Os) and mercury (Hg) concentrations represent two such proxies (see reviews by Grasby et al. 2019;Dickson et al. 2021;Percival et al. 2021a).Differences in the compatibility of osmium and rhenium (Re) within silicates result in mantle rocks and LIP basalts derived from mantle-plume volcanism typically being enriched in primitive osmium compared with rhenium, giving them an unradiogenic Os-isotope composition ( 187 Os/ 188 Os ≈ 0.13; Allègre et al. 1999).Thus, intense LIP volcanism and/ or the weathering (or submarine alteration) of juvenile basalts should cause a geologically rapid increase in the flux of unradiogenic Os to the global ocean relative to the comparatively constant background input from mid-ocean ridges and extraterrestrial material, lowering seawater 187 Os/ 188 Os ratios (e.g.Peucker-Ehrenbrink and Ravizza 2000).Conversely, the felsic crust is marked by higher Re/Os ratios, with decay of 187 Re to 187 Os over long geological timescales, eventually resulting in a more radiogenic osmium-isotope composition of continental material and the riverine runoff derived from its erosion (average 187 Os/ 188 Os ≈ 1.4 today; Peucker-Ehrenbrink and Jahn 2001).Thus, seawater 187 Os/ 188 Os ratios can also increase if continental weathering rates increase.Because osmium has a seawater residence time of tens of thousands of years, any significant change in the oceanic 187 Os/ 188 Os composition should be recorded in the hydrogenous phase of sedimentary rocks deposited throughout the global ocean, apart from hydrographically restricted basins where local sources may dominate (e.g.Paquay and Ravizza 2012;Dickson et al. 2015).
Mercury is emitted to the atmosphere as a volcanic volatile, with this source acting as a major natural source of the element to the Earth's surface today (Pyle and Mather 2003).In the stratosphere, mercury has a residence time of 0.5-2 years, enabling it to be distributed over a hemispherical-global scale before being deposited in sediments (Schroeder and Munthe 1998;Selin 2009).Hence, numerous studies have utilized Hg contents as a proxy for volcanic activity in Earth's past, normalizing against total organic carbon (TOC) to account for the element typically being bound to organic compounds during deposition (see reviews by Grasby et al. 2019;Shen et al. 2020;Percival et al. 2021a).Under highly euxinic conditions where free sulfides precipitate in the water column, Hg may be associated with that phase (Shen et al. 2019(Shen et al. , 2020)).Mercury may be adsorbed onto clay minerals if neither organic carbon or sulfides are present (Kongchum et al. 2011;Shen et al. 2020).Adsorption onto iron-manganese oxides or hydroxides has also been reported in oxidized settings, particularly for some riverine, lacustrine and estuarine environments (e.g.Quémerais et al. 1998;Driscoll et al. 2013;Maher et al. 2020), although studies of other sites have reported preferential binding onto organic compounds (e.g.Feyte et al. 2010;Cossa et al. 2021).
Any or all of remobilization of sedimentary mercury related to water-column and/or sediment redox changes, input from wildfires or enhanced terrigenous runoff, variation in organic matter type (from marine to terrestrial) and diagenetic removal of sedimentary TOC can alter sedimentary Hg contents and Hg/TOC ratios independently of volcanism (e.g.Hammer et al. 2019;Them et al. 2019;Charbonnier et al. 2020a;Frieling et al. 2023).Thus, reconstructing a volcanically triggered perturbation to the global mercury cycle depends on documenting stratigraphically correlative Hg enrichments in multiple sites distributed across the world and considering any potential local inputs of the element.Moreover, submarine volcanic activity may have a very limited dispersal range of mercury owing to the rapid scavenging and transformation of mercury species in hydrothermal vents, resulting in an apparent nearfield drawdown of Hg emitted directly into seawater (within hundreds of kilometres; Bowman et al. 2015; see also Scaife et al. 2017;Percival et al. 2018Percival et al. , 2021b)).
Despite the advances in tracking LIP volcanism in stratigraphic records of major climate/environmental change, relatively few studies have combined geochemical proxies of volcanism and surface temperature in the same site (but see Robinson et al. 2009;Bottini et al. 2015;Schoene et al. 2019;Percival et al. 2020).Thus, questions remain regarding the extent to which volcanic activity triggered climate warming and consequent environmental degradation during the Cretaceous.These questions particularly pertain to OAE 1a, with magmatic CO 2 , thermogenic carbon, and methane clathrate release all suggested as a cause of climate warming during the onset of the event (e.g.Jahren et al. 2001;Méhay et al. 2009;Naafs et al. 2016;Adloff et al. 2020;Deegan et al. 2022).Moreover, hydrothermal nutrient (e.g.Fe, V, Cu) output from submarine G-OJP volcanism may have directly enhanced primary productivity and triggered widespread marine anoxia during OAE 1a independently of global temperature change (see Erba et al. 2015).
Here, the record of LIP volcanism during OAE 1a is investigated at two sites that have been previously studied for palaeotemperature trends (Fig. 2): DSDP Site 398 (Vigo Seamount, Atlantic Ocean) and Alstätte-1 (Lower Saxony Basin, NW Germany).The documented relationship between volcanism and environmental change during OAE 1a is placed into the wider context of how LIP activity impacted the global climate during that event.This scenario is then discussed in a broader overview of the geochemical records and environmental and biotic impacts of large igneous provinces during the Cretaceous Period.

Study sites
The expanded Early Aptian stratigraphic record recovered from DSDP Site 398 consists of repeating dark grey-green turbiditic sedimentary sequences spanning several tens of metres in thickness.The strata are composed primarily of dark mudstones and claystones with some layers of calcareous sandstones and graded quartz-based sandstones-mudstones, deposited on the southern flank of the Vigo Seamount, just off the Iberian Margin (Ryan et al. 1979;Sigal 1979).Preservation of terrigenous material (including terrestrial organic matter) indicates that the sediments were probably deposited as part of a submarine fan, potentially related to deltaic progradation (Arthur 1979).However, a generally low BIT (branched and isoprenoid tetraether) index of organic matter studied from the sediments suggests that terrestrial organics do not dominate the setting (Naafs and Pancost 2016).There is no black-shale layer at this locality comparable with those typically associated with Lower Aptian stratigraphic records (Arthur 1979).Instead, OAE 1a strata have been identified by trends in the δ 13 C composition of organic matter, in combination with calcareous nannofossil biostratigraphy (Bralower et al. 1994;Li et al. 2008).Palaeoenvironmental reconstructions from this site, based on TEX 86 , have highlighted warm sea-surface temperatures prior to OAE 1a, which rose slightly during the event but significantly declined after it (Naafs and Pancost 2016).
The Alstätte-1 section represents a boreal record of OAE 1a, and primarily consists of organic-rich laminated mudstones and marlstones that were deposited into the epicontinental Lower Saxony Basin during the Early Aptian (Hoffmann and Mutterlose 2011).Palaeogeographic reconstructions, geochemical data, palaeontological findings and the preservation of organic-rich sediments deposited under anoxic conditions all indicate that the basin was probably hydrographically restricted to some degree during Barremian-Aptian times, potentially including the OAE 1a interval (Pauly et al. 2013).However, Aptian fauna recovered from strata below and in the OAE 1a interval are mainly of Boreal endemic composition and co-occur with cosmopolitan taxa (ammonites, belemnites), with the earliest cosmopolitan species documented a few metres below the OAE 1a level.Highly distinctive Tethyan belemnites are preserved directly above OAE 1a strata, suggesting strong sea-way connectivity and faunal exchange with the NW Tethys immediately following the event (Mutterlose 1998;Mutterlose and Böckel 1998).Lower Aptian strata and the OAE 1a interval have been identified from a combination of belemnite, ammonite and nannofossil biostratigraphy, integrated with δ 13 C studies (Bottini and Mutterlose 2012).TEX 86 -derived seasurface temperature reconstructions highlight a sharp rise from ,30°C prior to OAE 1a to 33°C during the event, but mild cooling during the middle of the OAE and after it had concluded (Mutterlose et al. 2014).The initial warming is potentially supported by a negative excursion in belemnite δ 18 O correlative with the shift in TEX 86 (Mutterlose et al. 2014).

Methodology Osmium-isotope analysis
Osmium abundances and isotopic values were determined by isotope dilution and negative thermal ionization mass spectrometry on a Thermo Fisher Scientific TRITON at the Japan Agency for Marine-Earth Science and Technology (JAMSTEC), Japan.Sample preparation followed the protocol described in Ishikawa et al. (2014) and Matsumoto et al. (2020), and broadly follows the procedures developed by Cohen and Waters (1996) and Birck et al. (1997).Briefly, powdered samples were spiked with a 185 Re-and 190 Os-rich solution and sealed in quartz glass tubes with 4 ml of inverse aqua regia.After being heated at 230°C for 48 h, osmium was separated and purified by CCl 4 extraction, HBr backextraction and standard micro-distillation techniques.Rhenium was purified from the aqua regia phase through anion chromatography, with Re concentrations determined via quadrupole inductively coupled plasma mass spectrometry (Thermo Fisher Scientific iCapQ) at JAMSTEC.Procedural blanks averaged 0.10 + 0.08 pg Os and 22 + 10 pg Re, with a 187 Os/ 188 Os ratio of 0.166 + 0.004 (1 SD, n = 3).Analytical uncertainty and precision were monitored through repeated measurements of the JMC Os standard, giving 187 Os/ 188 Os values of 0.10691 + 0.00008 (1 SD, n = 3), consistent with previously reported values ( 187 Os/ 188 Os = 0.106838 + 0.000015; Nozaki et al. 2012).The past seawater Os-isotope composition at the time of deposition ( 187 Os/ 188 Os (i) ) is calculated from the modern-day measured 187 Os/ 188 Os and 187 Re/ 188 Os compositions of a sedimentary rock sample and its age in order to account for the postdepositional decay of 187 Re to 187 Os (Cohen et al. 1999).The initial sedimentary Os concentration at the time of deposition [Os (i) ] is determined by utilizing the determined 187 Os/ 188 Os (i) ratio together with the measured 187 Os/ 188 Os composition and Os and 192 Os concentrations of a rock sample, assuming a 192 Os/ 188 Os of 3.08271 (Percival et al. 2021b).

Mercury concentrations
Mercury concentrations were determined on an Advanced Mercury Analyser 254.7 at the Vrije Universiteit Brussel (VUB), Belgium, following the procedure outlined in Liu et al. (2021).The limit of detection for the Advanced Mercury Analyser based on repeated blank analyses (n = 33 across the two sample sets) was 0.09 ng, calculated as 3 times the standard deviation on the blank measurements, one to two orders of magnitude below the content in a measured sample.A 100 + 2 mg aliquot of powdered sample was used per analysis, with repeatability of the results monitored through at least two measurements for each sample.Analytical accuracy was further tested through repeated measurements of the certified reference materials MESS-3 (marine sediment) and JP-1 (peridotite), yielding average concentrations of 89.51 + 1.89 ng/g (1 SD, n = 25) and 4.86 + 0.38 ng/g (1SD, n = 13), respectively, across analysis of the two sample sets, consistent with established compiled values (MESS-3 = 91 + 9 ng/g; JP-1 = 5.3 ng/g).

Total organic carbon and total sulfur contents
Total organic carbon and total sulfur (TS) contents were determined on a Nu Instruments Horizon 2 isotope ratio mass spectrometer coupled to a Eurovector elemental analyser, EuroEA3000, at the VUB.Prior to TOC analyses, the samples were decarbonated with 10% HCl following the procedure outlined in Liu et al. (2021), in order to remove inorganic carbon (assumed to be entirely in carbonate form).Carbon content calibration was performed using the certified reference material IAEA-CH-6 (sucrose).Data accuracy and reproducibility were monitored through repeated analyses of international standards IVA33802151 (organic-rich sediment) and IVA33802153 (organic-poor soil), with analytical uncertainty typically better than +0.1 wt% (1 SD).The measured carbon content in decarbonated aliquots of each sample was converted to a bulk rock TOC value by accounting for the percentage sample mass lost following decarbonation.The TS contents were measured on bulk-rock powder samples without any chemical pre-treatment.Sulfur content calibration on the isotope ratio mass spectrometer was performed using the certified reference material IAEA-S-1 (silver sulfide), with data accuracy and reproducibility again better than +0.1 wt% (1 SD) based on repeated analyses of international standards: IVA33802151 and MESS-2 (estuarine sediment).

Results
Early Aptian strata from DSDP Site 398 record a shift in 187 Os/ 188 Os (i) ratios from a mean background of c. 0.60 to a highly unradiogenic composition of c. 0.20 across the OAE 1a level, before a return towards more radiogenic pre-event values above it (Fig. 3a).The sample with the most unradiogenic Os-isotope composition is also marked by a large enrichment in [Os] (i) (1323 pg/g, compared with a mean of 173 pg/g for the rest of the record).A transient rise in 187 Os/ 188 Os (i) up to 0.69 in basal OAE1a strata is superimposed onto the initial unradiogenic signal.Mercury concentrations show significant variation, ranging from 18 to 100 ng/g (mean 49 ng/g) below the OAE 1a level and in basal OAE strata, before rising slightly to an average of 63 ng/g with a peak of 129 ng/g in the upper part of the OAE level, and further still above it (averaging 86 ng/g with spikes up to 220 ng/g).While TOC and TS contents also vary considerably at DSDP Site 398 (between 0.19-2.09and 0.10-1.25 wt%, respectively), there is no clear stratigraphic trend.Thus, the Hg enrichments are largely mirrored by peaks in Hg/TOC and Hg/TS ratios.There is a clear Hg/TOC peak in the middle part of the OAE 1a level between 1563 and 1557 mbsf (average Hg/TOC = 197 ng/g/wt%; average Hg/TS = 342 ng/g/wt%), and additional spikes of up to 394 ng/g/wt% and 1054 ng/g/wt%, respectively, above it (Fig. 3a).There is no clear linear relationship between Hg and either TOC or TS (R 2 values of 0.09 and 0.07, respectively).
The 187 Os/ 188 Os (i) trends documented at Alstätte-1 differ markedly from those of DSDP Site 398, rising from a background average of c. 0.77 to over 0.9 at the base of the OAE 1a level, but with values remaining high ( 187 Os/ 188 Os (i) .0.7) throughout the OAE strata (Fig. 3b).A shift to more unradiogenic Os-isotope compositions of c. 0.34 (and slight increase in [Os] (i) concentrations from 95 to 173 pg/g) only takes place above OAE 1a strata.Mercury concentrations at Alstätte-1 are also elevated around the OAE 1a level (mean 49 ng/g) with three spikes to c. 100 ng/g or more, compared with background contents of 27 ng/g (Fig. 3b).However, the TOC and TS contents also rise in OAE strata; thus, there is no systematic increase in Hg/TOC or Hg/TS in basal OAE 1a strata.There is an increase in Hg/TS from c. 32 to 100 ng/g/wt % in the upper part of the OAE level, but this rise is not correlative with any increase in Hg concentrations, rather with a fall to very low TS contents.Some small spikes in Hg/TOC (up to 109 ng/g/wt % compared with a background of 29 ng/g/wt%) do correlate with the three peaks in Hg concentrations; however, those samples also have very high sulfur contents.Indeed, there is a strong linear correlation between Hg and TS concentrations at Alstätte-1 (R 2 = 0.73; compared with R 2 = 0.05 between Hg and TOC), potentially suggesting that local drawdown with sulfide was the main control on Hg deposition under the highly oxygen-depleted conditions in the Lower Saxony Basin.OAE 1a black shales from Alstätte-1 are known to yield abundant pyrite, suggesting that sulfide sorption was the main control on Hg contents in this oxygen-depleted setting.

OAE 1a as a case study of how LIPs impacted the Cretaceous global environment
The 187 Os/ 188 Os (i) and [Os] (i) trends recorded at DSDP Site 398 closely resemble those from previous studies of OAE 1a.Very low 187 Os/ 188 Os (i) values recorded in middle-upper OAE 1a strata support a major influx of unradiogenic osmium to the global ocean during that event, probably linked to G-OJP activity (Tejada et al. 2009;Bottini et al. 2012;Adloff et al. 2020;Martínez-Rodríguez et al. 2021;Percival et al. 2021b).Recorded Os-isotope trends from several other sites also document a volcanically triggered shift towards a more unradiogenic seawater composition immediately prior to OAE 1a, with this change also potentially documented at DSDP Site 398 by a single data point (1171.33mbsf ).Basal OAE 1a strata only show a clear mercury enrichment at a single site very close to that LIP (Percival et al. 2021b).This non-dispersal of mercury is consistent with the short residence time of volcanic Hg emitted directly into seawater (see Bowman et al. 2015), and supports submarine volcanism on the G-OJP as the main source of unradiogenic Os to the global ocean during OAE 1a.
The shift to more radiogenic 187 Os/ 188 Os (i) compositions across OAE 1a strata at Alstätte-1 represents a markedly different trend from that recorded at all other sites.This disparity probably indicates that the Lower Saxony Basin was too hydrographically restricted for significant water-mass exchange with the global ocean, despite having enough connection to permit faunal migration.In this context, the shift to higher 187 Os/ 188 Os (i) values across the OAE 1a level at Alstätte-1 may highlight a rise in local riverine runoff to the Lower Saxony Basin from enhanced continental weathering in response to elevated atmospheric CO 2 and global temperatures, given that this flux typically represents the main source of radiogenic osmium to seawater.Notably, the radiogenic shift at Alstätte-1 begins at the base of C3 (3.8 m), and age-equivalent C3 strata at DSDP Site 398 (1169.8-1164.54mbsf ) and other locations are also marked by a transient rise in 187 Os/ 188 Os (i) values between the unradiogenic pulses (Fig. 4).This stratigraphic correlation supports previous interpretations that the transient radiogenic Os-isotope shift recorded in C3 strata at DSDP Site 398 and elsewhere probably reflects a weathering pulse superimposed on an overarching signal of submarine LIP activity, rather than a spell of volcanic quiescence between two distinct pulses of volcanism (e.g.Tejada et al. 2009;Bottini et al. 2012;Martínez-Rodríguez et al. 2021).Strontium-, calcium-and lithium-isotope studies of OAE 1a also support a rise in global weathering rates during the onset of the event (Jones and Jenkyns 2001;Blättler et al. 2011;Lechler et al. 2015).
When stratigraphic trends in 187 Os/ 188 Os (i) are compared with temperature proxies from records of OAE 1a, it is clear that the highest temperatures indicated by δ 18 O or TEX 86 data are documented above the radiogenic osmium shift and the C3 excursion, stratigraphically correlative with the most unradiogenic 187 Os/ 188 Os (i) compositions recorded in OAE 1a strata (upper C3-lower C4; see also Bottini et al. 2015).These trends support an intensification in LIP volcanism and associated carbon emissions as the main driver of this further rise in global temperatures during OAE 1a.A pronounced increase in Hg contents and Hg/TOC enrichments across upper C3-lower C4 strata at DSDP Site 398 and some (although not all) other sites across the globe might reflect a potential intensification in volcanic activity or switch to subaerial eruptions after the start of OAE 1a (see Charbonnier and Föllmi 2017;Percival et al. 2021b;Vickers et al. 2023).The latter possibility is consistent with a scenario in which the G-OJP was initially emplaced through submarine volcanism (enabling worldwide dispersal of unradiogenic Os but only nearfield Hg enrichment), before becoming emergent above the sea surface and leading to more global distribution of mercury from subaerial eruptions.However, it should be noted that the C3-C4 Hg enrichments could also reflect localized environmental perturbations such as redox changes or enhanced wildfires/runoff of terrestrial organic matter (see Percival et al. 2021b).A not exclusively volcanic cause of the C3-C4 mercury enrichments is supported by the enhanced Hg and Hg/TOC values recorded above the OAE 1a level at DSDP Site 398, in strata that post-date the time of LIP volcanism reconstructed from 187 Os/ 188 Os (i) trends.Unless subaerial eruptions occurred later than and entirely separately from submarine volcanism (which is implausible if both were related to the same LIP), it is highly unlikely that these post-event Hg peaks were volcanically derived.Instead, the post-OAE Hg enrichments may have been caused by variability in the input of terrestrial organic matter from the nearby palaeoshoreline (Hammer et al. 2019), or redox fluctuations/the return to a more oxygenated water column (Frieling et al. 2023), although there is no direct evidence of major redox change at DSDP Site 398.
There is also a potential correlation between an initial shift towards more unradiogenic Os-isotope values and possible higher sea-surface temperatures below the C3 negative δ 13 C excursion at DSDP Site 398.This earlier temperature rise is less pronounced than the later one during the OAE itself, and the fall in 187 Os/ 188 Os (i) values is based on one data point.Nonetheless, an unradiogenic shift in the marine Os-isotope composition and rise in surface temperatures just prior to the C3 excursion and onset of OAE 1a is supported by data from other stratigraphic archives in the NW Tethys and the Subbetic margin (e.g.Tejada et al. 2009;Keller et al. 2011;Bottini et al. 2012Bottini et al. , 2015;;Lorenzen et al. 2013;Castro et al. 2021;Martínez-Rodríguez et al. 2021;Fig. 4).A crisis in nannoconid fauna is also recorded below the base of OAE 1a strata sensu stricto in several sites (see the section on the biotic response to Cretaceous LIP formation).Lead-isotope and trace-element datasets further support an onset of volcanic activity that was coeval with the nannoconid crisis but prior to the C3 excursion (Kuroda et al. 2011;Erba et al. 2015; Figure 4).Assuming that there was a single prolonged pulse in G-OJP activity recorded by the unradiogenic shifts in 187 Os/ 188 Os (i) , which began before and continued after a transient increase in continental weathering rates, the new and previously published geochemical and palaeontological datasets both suggest that magmatic carbon emissions associated with submarine LIP volcanism probably triggered the initial warming and nannoconid crisis immediately prior to the onset of OAE 1a.These datasets also support a non-volcanic carbon source, probably related to thermogenic emissions or methane clathrate release, as the driver of the C3 δ 13 C negative excursion.Osmium-isotope trends in those strata primarily record enhanced continental weathering rather than LIP activity.Furthermore, there is no clear global mercury signal of volcanism at that time.Both methane clathrates and thermogenic carbon are isotopically lighter than magmatic CO 2 , with the latter potentially sourced from intrusive magmatism related to the HALIP (see above; also Polteau et al. 2016).Thus, either carbon source could potentially have caused the δ 13 C negative shift without greatly changing atmospheric CO 2 levels or affecting global temperatures (Méhay et al. 2009;Naafs and Pancost 2016;Adloff et al. 2020).This scenario is consistent with the geochemical and palaeontological evidence for volcanism as the main driver of climate warming and biota stress prior to and during OAE 1a, but with a non-volcanic carbon source as the driver of the C3 negative excursion (Méhay et al. 2009;Adloff et al. 2020).However, it is less clear which carbon source (if either) played the dominant role in initiating the widespread marine anoxia that characterized the OAE.

Overview of geochemical evidence for the (variable) links between LIP volcanism and Cretaceous episodes of major environmental change
The entirety of Cretaceous stratigraphic history has been investigated for trends in oceanic strontium-isotope compositions (see reviews by Jones and Jenkyns 2001;McArthur et al. 2020), and large parts also for seawater 187 Os/ 188 Os (i) reconstructions (see e.g.Peucker-Ehrenbrink and Ravizza 2000;Matsumoto et al. 2022;and Fig. 5).Reconstructed 187 Os/ 188 Os (i) trends for most Cretaceous episodes of environmental change are similar to OAE 1a, showing pronounced unradiogenic shifts that are assumed to result from one or more of submarine volcanism, the weathering of juvenile LIP basalts or some other form of basalt-seawater interaction (Fig. 5a, and references therein).Some more localscale episodes of environmental perturbation are also marked by unradiogenic shifts of a smaller magnitude (Matsumoto et al. 2021b).Thus, osmium-isotope stratigraphy strongly supports the link between LIP emplacement and Cretaceous climate/environmental change.Additionally, a more gradual 187 Os/ 188 Os (i) trend towards more unradiogenic compositions in the Late Cretaceous following OAE 2, an interval not coeval with environmental change, has been linked with weathering of Caribbean and Madagascan LIP basalts (Matsumoto et al. 2023a).Interestingly, clear shifts towards primitive-mantle compositions of strontium isotopes are only documented for OAE 1a and OAE 2 (Fig. 5b), albeit with less sharp stratigraphic trends owing to the longer residence time of strontium in seawater (.2 Myr;Palmer and Edmond 1989).Those intervals are also associated with the highest magnitude unradiogenic Os-isotope shifts in the Cretaceous.Collectively, the strontium and osmium trends may suggest that LIP activity, or at least the flux of mafic elements to seawater, was greater during those intervals than for other Cretaceous events.In addition to strontium and osmium datasets, evidence for some form of magmatic activity during OAEs 1a and 2 has also been inferred on the basis of lead-, neodymium-, chromium-and zinc-isotope studies, and trace-element enrichments (Larson and Erba 1999;Snow et al. 2005;Kuroda et al. 2007;Zheng et al. 2013;Erba et al. 2015;Holmden et al. 2016;Sweere et al. 2018).
Stratigraphic mercury records of Cretaceous events also yield information on the style and/or provenance of magmatic activity.Like OAE 1a, where mercury enrichment in basal C3 strata was only documented at a site proximal to the G-OJP, there is limited global evidence for peaks in Hg and Hg/TOC in OAE 2 strata, with small enrichments at some records deposited relatively proximally to LIPs (Scaife et al. 2017, Percival et al. 2018).This pattern of a more localized mercury enrichment proximal to a LIP source is consistent with largely submarine volcanism during both events (Scaife et al. 2017;Percival et al. 2018Percival et al. , 2021b)).Mercury records associated with OAE 1b are more limited, but spikes have been reported from the NW Tethys (Sabatino et al. 2018).If these Hg enrichments were indeed volcanically derived, it may indicate that some Kerguelen Plateau volcanism at that time was subaerial.Albian-age mercury enrichments have also been reported from a lacustrine record in North China (Zhao et al. 2022), although stratigraphic uncertainties hinder a clear correlation with the Tethys record or a connection to Kerguelen volcanism.
Interestingly, mercury records of Paraná-Etendeka and Deccan volcanism show considerable variability across different sites, despite both LIPs forming through subaerial eruptions (e.g.Percival et al. 2018;Fendley et al. 2019;Charbonnier et al. 2020b).Hg enrichments have been reported from strata that record the onset of the Weissert Event (Charbonnier et al. 2017(Charbonnier et al. , 2020b)), but because the duration of Paraná-Etendeka volcanism was at least 2-3 Myr (see the section introducing the Paraná-Etendeka LIP; also Dodd et al. 2015;Gomes and Vasconcelos 2021;Bacha et al. 2022), those peaks cannot reflect the entire eruptive history of that LIP.It is possible that the less clear global mercury signals of the Paraná-Etendeka LIP reflect a generally lower volatile output or slower eruption rate than for other subaerial provinces (see Callegaro et al. 2014;Dodd et al. 2015).In this case, the Hg enrichments reported by Charbonnier et al. (2017Charbonnier et al. ( , 2020b) ) in basal Weissert Event strata could record a peak in the intensity of Paraná-Etendeka volcanism coeval with the onset of the crisis, consistent with the geochronological models of Gomes and Vasconcelos (2021) and Bacha et al. (2022).However, local environmental or sedimentological causes for those Hg peaks cannot be ruled out, particularly if significant Paraná-Etendeka volcanism post-dated the onset of the Weissert event, as suggested by Rocha et al. (2020).Similarly, uppermost Maastrichtian Hg peaks have been associated with Deccan eruptions (Keller et al. 2020), most notably just below the KPg boundary (e.g.Font et al. 2016;Keller et al. 2018;Percival et al. 2018;Fendley et al. 2019).This peak stands in contrast to a comparative lack of evidence for a global-scale systematic enrichment in mercury at the onset of Deccan volcanism 300-400 kyr prior to the end of the Maastrichtian (Percival et al. 2018).Thus, it may reflect an intensification of volcanism immediately prior to the extinction, consistent with the U-Pb geochronological model of Deccan eruptions (Schoene et al. 2019).However, given that this Hg enrichment is not recorded globally, a non-volcanic cause cannot be excluded (which would be potentially consistent with the Deccan 40 Ar/ 39 Ar geochronology of Renne et al. 2015, andSprain et al. 2019).Mercury enrichments have also been reported for more minor episodes of environmental change, such as OAE 1d and the Faraoni Event (Charbonnier et al. 2018;Yao et al. 2021), but the possible existence and nature of any causal relationship between those crises and volcanism remains unclear.
Cyclostratigraphic and magnetostratigraphic correlation of KPg records clearly shows that the main phase of Deccan volcanism commenced coevally with the onset of late Maastrichtian climate warming (Schoene et al. 2019;Hull et al. 2020), as appears to have been the case for LIP activity and global temperature increases during OAE 1a.Geochemical evidence for a causal link between volcanism and climate warming during OAE 2 is also well documented at the stratigraphically condensed record of ODP Site 1260 (Forster et al. 2007;Turgeon and Creaser 2008).More expanded OAE 2 records, with better temporal resolution, show that LIP activity began prior to the onset of the OAE sensu stricto, leaving the exact chronology of volcanism v. climate warming at that time unclear (e.g.Du Vivier et al. 2014).However, evidence of transient cooling pulses during both OAE 1a and OAE 2, such as the Plenus Cold Event (see the section introducing the Caribbean LIP and OAE 2), is recorded in strata which also record Os-isotope evidence of intense LIP activity (Bottini et al. 2015;Jenkyns 2018;Percival et al. 2020).The late Maastrichtian climate warming was also alleviated while Deccan volcanism was ongoing (e.g.Sprain et al. 2019;Hull et al. 2020).There are two possible explanations for these transient falls in temperature during times of LIP volcanism.The first model is that the LIP eruptions associated with those events featured spells during which volcanic carbon emissions were reduced (e.g.Sprain et al. 2019;Hull et al. 2020;Hernandez Nava et al. 2021).Alternatively, and more plausibly for the OAEs, magmatic CO 2 output may have been offset by one or both of enhanced organic-carbon burial and an increase in global silicate weathering rates (Pogge von Strandmann et al. 2013;Jenkyns et al. 2017;Jenkyns 2018;Robinson et al. 2019;Percival et al. 2020;Papadomanolaki et al. 2022).While the latter scenario might be expected to increase the input of radiogenic osmium to the ocean, apparently at odds with the low 187 Os/ 188 Os values documented during the Plenus Cold Event and other transient intervals of cooling during OAEs 1a and 2, it is possible that any continental weathering signal was overprinted by the unradiogenic Os flux from volcanism, or that primitive LIP basalts were the main lithology being eroded.
Climate cooling associated with the Weissert Event may have been caused by carbon sequestration during weathering of newly erupted Paraná-Etendeka basalts, based on stratigraphic correlations between unradiogenic Os-isotope shifts and evidence of surface temperature decrease in Valanginian strata (Percival et al. 2023).Elevated organic-matter burial in the terrestrial realm may have further drawn down carbon at that time (Westermann et al. 2010).Alternatively, volcanic SO 2 emissions might have caused the cooling by leading to the formation of stratospheric sulfate aerosols, which may also have helped cause transient temperature falls during the latest Maastrichtian (see Self et al. 2006;Schmidt et al. 2016;Callegaro et al. 2023).However, the short atmospheric residence time of those compounds and the slow eruption rate and potentially low sulfur content of Paraná-Etendeka basalts do not support this link (Callegaro et al. 2014;Dodd et al. 2015).Alternatively, the surface-temperature warming reconstructed from some NW Tethyan and proto-Atlantic records of the Weissert Event could suggest significant volcanic emission of CO 2 and other volatiles to the atmosphere, which may be supported by Hg enrichments in sedimentary records of that time (Charbonnier et al. 2017(Charbonnier et al. , 2020b)).Direct stratigraphic comparisons between global temperature and volcanic activity are currently lacking for other Cretaceous events.Nonetheless, it is clear that at least some Cretaceous LIPs had a pronounced impact on global climate, through volcanic CO 2 emissions and/or weathering of newly formed basalts.
Aside from changes in global climate, the environmental perturbations associated with the emplacement of Cretaceous LIPs were highly variable.Evidence of ocean acidification has been reported from stratigraphic records of OAEs 1a and 2, and potentially the late Maastrichtian, with this fall in seawater pH probably resulting chiefly from the same carbon emissions that triggered climate warming during those time intervals (e.g.Erba et al. 2010;Dameron et al. 2017;Jones et al. 2023).To date, these postulated episodes of ocean acidification have largely been inferred from sedimentological and palaeontological evidence (see the section on the biotic response to Cretaceous LIP formation), with little geochemical information regarding seawater pH.
While the rises in global temperature during OAEs 1a and 2 were marked by the development of widespread marine anoxia/euxinia in both epicontinental shelf basins and across the global ocean, evidence for similar oceanic redox changes during the Late Maastrichtian warming or Weissert Event is limited.For example, numerous sedimentary records of OAEs 1a and 2 from both epicontinental shelf basins and open-ocean sites are marked by the appearance of organic-rich shales and changes in organic biomarker data, iron speciation and variations in sedimentary elemental contents (particularly P, V, Mn, Mo, U and TOC/P and I/Ca ratios) and N-, S-, Fe-, Mo-, Tl-and Uisotope compositions, all of which are redoxsensitive proxies (e.g.Sinninghe-Damsté and Köster 1998; Kuypers et al. 2004;Pancost et al. 2004;Jenkyns et al. 2007Jenkyns et al. , 2017;;Junium and Arthur 2007;Mort et al. 2007;van Bentum et al. 2009;Lu et al. 2010;Montoya-Pino et al. 2010;Föllmi 2012;Owens et al. 2012Owens et al. , 2013;;Ruvalcaba Baroni et al. 2015;Dickson et al. 2016;Ostrander et al. 2017;Clarkson et al. 2018;Siebert et al. 2021; see also a review by Jenkyns 2010).In the case of the Weissert Event, the limited expansion of anoxic water masses (and lithological/geochemical evidence thereof ) may be attributed to the apparent lack of a clear temperature increase and associated seawater stratification/nutrient runoff and eutrophication in many parts of the world.However, the Late Maastrichtian Warming Event featured a temperature rise of 2°C (Hull et al. 2020); thus, an increase in seawater deoxygenation and evidence of this phenomenon might be expected.It is possible that the continued break-up of Pangaea through the Cretaceous caused evolution in the global palaeogeography, shelf-basin extent and ocean circulation, potentially decreasing the propensity of marine environments to become oxygen depleted by the Maastrichtian.This mechanism has been proposed to explain the far greater extent of marine anoxia during OAEs 1a and 2 than the Paleocene-Eocene Thermal Maximum (56 Ma), despite the latter event being marked by climate warming of a similar or greater magnitude and at a faster rate than for the Cretaceous events (e.g.Jenkyns 2010; Dickson et al. 2014;Clarkson et al. 2021).
Another hypothesized trigger of marine anoxia during OAEs 1a and 2 is the enhanced output of trace-metal micronutrients from submarine oceanplateau volcanism directly into seawater, elevating primary productivity and the consumption of oxygen from the water-column (Sinton and Duncan 1997;Kerr 1998;Snow et al. 2005;Erba et al. 2015; see Fig. 6).As a mechanism that depends on submarine LIP activity instead of continental volcanism, rather than global temperature increase, this model is consistent with the lack of widespread marine anoxia during the Weissert and Late Maastrichtian events, and is supported by enriched abundances of micronutrient metals in sedimentary records of OAEs 1a and 2, particularly those proximal to the G-OJP and Caribbean LIP (Larson and Erba 1999;Snow et al. 2005;Erba et al. 2015).However, OAE 1b also coincided with submarine volcanism during emplacement of the Kerguelen Plateau, and has also been associated with climate warming, yet anoxic sediments are largely found in Atlantic and NW Tethyan sites (see the section introducing the Kerguelen Plateau and OAE 1b).It is possible that this distribution may reflect the relative paucity of OAE 1b sites studied thus far.However, the Os-isotope record of OAE 1b documents a much lower magnitude unradiogenic shift than those recorded for OAEs 1a and 2 (see Fig. 4).Thus, submarine LIP volcanism and associated trace metal output was probably more intense at the time of OAEs 1a and 2 than during OAE 1b (or any other time in the Cretaceous, see above).OAE 2 was also coeval with enhanced seasonality related to a peak in orbital eccentricity, which might have increased the susceptibility of the Earth system to the development of marine anoxia (Batenburg et al. 2016).Furthermore, the relatively protracted nature of environmental change during OAE 1b (∼4 Myr; Gale et al. 2020;Ait-Itto et al. 2023), compared with the geologically rapid onset of OAEs 1a and 2, may have mitigated the severity of environmental change during the Aptian-Albian transition.

The biotic response to Cretaceous LIP formation
Fossil evidence of biotic changes during times of LIP formation during the Cretaceous highlights considerable variation in the response of the marine biosphere to these major volcanic episodes.Furthermore, diversity patterns of the main planktonic and benthic carbonate producers in the Cretaceous, evaluated at the generic level, document disjunct trends for the two groups (Steuber et al. 2023).These differences suggest the importance of distinct environmental parameters in controlling the radiation and decline of marine groups, following different ecological strategies.The terrestrial biosphere was also impacted to some degree by different Cretaceous LIPs, with indications of at least regional vegetation changes during the Weissert Event (Kujau et al. 2013) and OAE 1a (e.g.Cors et al. 2015;Galloway et al. 2022).A rise in terrestrial organic-matter burial associated with the Weissert Event δ 13 C positive excursion may also reflect changes in the continental biosphere (Westermann et al. 2010).Furthermore, charcoal preservation in Aptian-Albian and OAE 2 strata suggest increased wildfire activity during parts of the Cretaceous (Brown et al. 2012;dos Santos et al. 2016;Wang et al. 2019;Baker et al. 2020;Xu et al. 2022), potentially resulting from climate warming and elevated atmospheric oxygen levels following enhanced primary productivity in the case of the latter time interval (Baker et al. 2020).However, Heimhofer et al. (2004) reported relatively stable terrestrial vegetation during and after OAE 1a.Overall, the comparative paucity of the continental stratigraphic record compared with that of the marine realm means that the ocean biosphere is much more studied and better understood, and is the focus of this section hereafter.
A pronounced impact of environmental change associated with Cretaceous LIPs has been documented for benthic carbonate platform dwellers (dasycladales, larger benthic foraminifera, corals, rudists; see overview by Steuber et al. 2023;Figs 5c-e).Shortly before or during the onset of OAE 1a, carbonate platforms and their benthic biota experienced a substantial decline, particularly in the mid latitudes.Dasycladacean algae and large benthic foraminifera are marked by significant diversity declines across OAE 1a and (especially) OAE 2, with aragonitic taxa chiefly affected.Rudist genera disappeared during OAE 1a from higher latitudes but survived in more equatorial settings.This abrupt decline of aragonitic taxa is thought to be related to the major increases in atmospheric CO 2 and very high sea-surface temperatures caused by LIP eruptions, resulting in a shallowing of the carbonatecompensation depth (e.g.Ridgwell 2005;Foster et al. 2017), reduced seawater carbonate saturation and a subsequent calcification crisis (Bauer et al. 2017).
The same mechanism could have caused the retreat of shallow-water carbonate platforms to low latitudes in the aftermath of OAEs 1a and 2. Coral diversity was significantly reduced in the aftermath of OAE 2, and to a lesser extent after OAE 1a (Fig. 5e), with high sea-surface temperatures and presumably low seawater pH probably reducing the photosymbiotic activity of zooxanthellid corals and large benthic forams (Kiessling and Kocsis 2015;Reddin et al. 2021;Steuber et al. 2023).This interpretation is supported by an increase of azooxanthellate coral genera after OAE 1a, potentially reflecting a bleaching effect.Corals living in symbioses with dinoflagellates were thus partly replaced by corals missing these symbionts, supporting a decline of photosymbiosis owing to high temperatures (e.g.Kiessling and Kocsis 2015).Carbonate platform drowning has also been documented during several intervals in the Early Cretaceous, particularly in the NW Tethys, and is often linked with increased volcanic activity on one or both of the G-OJP or Kerguelen Plateau (e.g.Föllmi et al. 1994;Weissert et al. 1998;Föllmi 2012;Charbonnier et al. 2018;Matsumoto et al. 2020Matsumoto et al. , 2021bMatsumoto et al. , 2022)).A spread of warm-water scleractinian corals has been documented in uppermost Maastrichtian strata of Belgium and the Netherlands, and is correlative with evidence for rising seawater temperatures (Leloux 1999;O'Hora et al. 2022).This relationship supports some benthic response to rising global temperatures associated with early Deccan volcanism.However, the extent of biospheric changes during the Late Maastrichtian Warming Event remains debated (see e.g.Witts et al. 2016;Tobin et al. 2017 regarding high-latitude molluscs).
The diversity of planktonic carbonate producers (calcareous nannofossils, calcispheres, planktonic foraminifera) was apparently less impacted by environmental change related to Cretaceous LIP volcanism.A major turnover in planktonic carbonate producers did occur during the KPg mass extinction (e.g.Smit 1982;Olsson and Liu 1993;Pospichal 1996), but no specific declines in diversity have been recorded at the generic level for other intervals coeval with Cretaceous LIPs (Steuber et al. 2023).Rather, planktonic taxa diversity increased consistently throughout the Cretaceous, reaching a broad maximum between the Albian and Maastrichtian stages, peaking around the Campanian (Suchéras-Marx et al. 2019;Steuber et al. 2023;Fig. 5f).In general, this trend towards increased diversity applies to both calcareous nannofossils and planktonic foraminifera (Suchéras-Marx et al. 2019;Steuber et al. 2023), despite their disparate trophic levels.However, phytoplanktonic organisms show some change in records of the Late Maastrichtian warming, with warm-water species tending to spread in favour of cold-water taxa, together with blooms in both dinoflagellate and calcareous nannoplankton highlighting stressed ecosystems in response to the rise in global temperatures caused by Deccan volcanism at that time (e.g.Thibault and Gardin 2007;Sheldon et al. 2010;Vellekoop et al. 2019).Planktonic foraminifera show a more locally variable response to Maastrichtian warming, possibly related to local controls such as nutrient levels and/or freshwater runoff as well as temperature changes (e.g.Abramovich and Keller 2002;Woelders et al. 2017).
Extensive study of the nannoconid crisis associated with OAE 1a has shown it to be synchronous at a global scale, and apparently representing the culmination of a decrease in nannofossil calcite production that commenced in the latest Barremian (Erba 1994(Erba , 2004;;Bottini and Mutterlose 2012;Aguado et al. 2014;Erba et al. 2015;Bonin et al. 2016;Giraud et al. 2018;Mahanipour et al. 2019).In addition to major changes in nannofossil abundance, species-specific decreases in size have been detected for both OAE 1a and OAE 2. Specifically, Biscutum constans shows a marked reduction in coccolith size across OAE 1a at global scale (Erba et al. 2010;Lübke et al. 2015;Lübke and Mutterlose 2016;Bottini and Faucher 2020).Coccolith dwarfism of B. constans during OAE 2 has also been documented globally, alongside a fall in nannofossil abundance and species richness worldwide (Nederbragt and Fiorentino 1999;Hardas and Mutterlose 2007;Linnert et al. 2010Linnert et al. , 2011;;Corbett and Watkins 2013;Faucher et al. 2017).It should be noted that these changes did not result in extinctions, with species recovering following the alleviation of environmental changes, and some nannofossil taxa actually increasing in abundance (Erba and Tremolada 2004;Erba et al. 2019).Nonetheless, the transient net decrease in carbonate producers reduced the availability of nannofossil micrite and was reflected as a biocalcification crisis, since nannoconids were the main rock-forming nannofossils during the Early Cretaceous (Erba and Tremolada 2004;Weissert and Erba 2004).This interpretation has been challenged recently by Slater et al. (2022), who argued that the apparent decrease in nannoplankton preservation results from post-depositional removal of carbonate from the rock record and that diversity increased during OAEs.However, while both of these phenomena did indeed occur, they do not explain the geologically rapid loss of heavily calcified nannoconids compared with other calcifiers (Erba et al. 2010), which is more consistent with a biocalcification crisis.Furthermore, fossil evidence of reduced biocalcification during OAE 2 is supported by sedimentological changes consistent with a shallowing of the seawater carbonate compensation depth in response to ocean acidification at that time (Erba 2004;Faucher et al. 2017;Jones et al. 2023).
While ocean acidification caused by volcanic CO 2 emissions from LIP eruptions is considered to be the main driver of reduced biocalcification during OAEs 1a, and OAE 2 (Erba 2004;Erba et al. 2010), alternative processes may also have played a role.Enhanced seawater fertility and subsequent eutrophication have also been proposed as a cause of nannoplankton stress and reduced species richness that favoured smaller taxa (Bersezio et al. 2002;Gréselle et al. 2011;Aguado et al. 2014;Duchamp-Alphonse et al. 2014;Mattioli et al. 2014;Giraud et al. 2018;Shmeit et al. 2022).At open oceanic sites, fertilization was potentially driven by the hydrothermal output of biolimiting trace metals from G-OJP or Caribbean LIP activity (Snow et al. 2005;Bottini et al. 2015;Erba et al. 2015;Faucher et al. 2017;Bottini and Erba 2018;Bottini and Faucher 2020).Increased runoff of terrestrial material and nutrients caused by volcanically triggered climate warming could have produced the same effect in continental shelf settings, while potentially also reducing light availability in the photic zone (Lübke et al. 2015;Lübke and Mutterlose 2016).Alternatively, output of toxic volcanic elements may have suppressed biocalcification.
A more progressive decline in nannoconid biocalcification also occurred during OAE 1b, but has been linked with cold conditions (McAnena et al. 2013).Alternatively, volcanic CO 2 output during the formation of the Kerguelen Plateau may have lowered seawater pH enough to reduce carbonate saturation, even if the carbon emissions were insufficient to cause prolonged global temperature rise.Longer-term cooling in the Late Aptian may have been linked to enhanced absorption of CO 2 in seawater, lowering its pH (Bottini et al. 2015).The reduction in size, decrease in abundance and species turnover of planktonic foraminifers (Huber and Leckie 2011) might reflect the response of calcareous zooplankton to volcanically triggered ocean acidification (Bottini et al. 2015;Bottini and Erba 2018).Analogous changes in the abundance, test size and evolutionary rates of calcareous zooplankton were documented for OAE 1a and OAE 2 (Premoli Silva et al. 1989Silva et al. , 1999;;Coccioni et al. 1992;Leckie et al. 2002) and are also consistent with ocean acidification (Kump et al. 2009;Hönisch et al. 2012).The cause of a nannoconid decline associated with the Weissert Event is less clear, although it has also been linked to increased fertility and seawater acidification related to volcanic CO 2 emissions (Erba and Tremolada 2004;Weissert and Erba 2004).However, there is no independent evidence for decreased seawater pH during that crisis, and as noted above, possible climate warming signals have only been reported for NW Tethyan and proto-Atlantic records of the Weissert Event thus far (see compilations by Charbonnier et al. 2020c;Cavalheiro et al. 2021).At present, the limited evidence for global-scale temperature rise at that time (Meissner et al. 2015;Price et al. 2018;Cavalheiro et al. 2021) hinders argument for a clear causal link between nannoconid stress and climate warming or CO 2 release comparable with that of OAEs 1a and 2.
It is clear that while rapid climate warming itself apparently impacted some marine groups during the Cretaceous, biospheric stress during intervals of LIP volcanism may not have been dependent on increased temperatures (Erba 2006).Nonetheless, aside from the KPg mass extinction (for which the influence of Deccan Trap volcanism was complicated by the coeval climatic effects of the Chicxulub impact), OAEs 1a and 2 were the intervals associated with LIP emplacement that featured the most severe perturbations to Earth's biosphere.Both OAEs featured significant climate warming, unlike the Weissert Event or OAE 1b.Moreover, the environmental change during OAEs 1a and 2 was also less protracted in nature and had a more rapid onset (probably owing to the more intense LIP activity coeval with those intervals), which may have helped drive a more severe biotic response.

Remaining questions and future perspectives for understanding Cretaceous LIPs
Previous overviews of LIPs and mass extinctions/ environmental change have largely focused on continental provinces.It has been hypothesized that the impact of oceanic LIPs was somewhat muted owing to the potentially lower degassing efficiency of submarine eruptions occurring under high seawater pressure (e.g.Courtillot and Renne 2003;Ganino and Arndt 2009;Green et al. 2022).However, submarine oceanic-plateau volcanism was probably the main driver of both OAEs 1a and 2, which were marked by the most severe climatic, environmental and biotic changes during the Cretaceous prior to the KPg extinction.Even if the widespread marine anoxia, ocean acidification and biocalcifcation crises during those events were driven by the output of trace-metal micronutrients and volcanic CO 2 and/or halogens into seawater (e.g.Kerr 1998;Erba et al. 2004Erba et al. , 2015;;Snow et al. 2005), both events also featured significant climate warming that was apparently driven by magmatic carbon emissions (Larson and Erba 1999;Jenkyns 2003;Forster et al. 2007;Méhay et al. 2009;Adloff et al. 2020; this study; Fig. 6).Thus, submarine LIPs were capable of raising global temperatures as well as triggering oceanic anoxia and acidification, three key processes thought to have helped cause major extinctions.
Recent studies have emphasized that the rate of eruptions (and, consequently, the assumed rate of magmatic CO 2 output) plays a key role in determining the environmental and biotic impact of a LIP (Green et al. 2022;Jiang et al. 2022).Constraining this CO 2 flux depends on both precise geochronology and robust estimates of the magmatic carbon content.Currently, reconstructed eruption histories vary greatly in terms of detail and precision for different LIPs, both from the Cretaceous and other time intervals (see reviews by Kasbohm et al. 2021;Jiang et al. 2023).Additionally, the CO 2 content of tholeiitic basalts remains relatively poorly constrained.Current estimates for LIP magma carbon contents range between 0.1 and 2 wt%, based on the study of analogous modern basaltic systems (e.g.Self et al. 2006;Saunders 2016), rare olivine melt inclusions (Black et al. 2014) and volatile/nonvolatile elemental ratios, particularly CO 2 /Nb and CO 2 /Ba (Black and Gibson 2019;Hernandez Nava et al. 2021;Boscaini et al. 2022).Thus far, these efforts have only focused on a few continental LIPs, with the Deccan Traps the only Cretaceous province investigated (Hernandez Nava et al. 2021).Future studies refining both the eruptive history of Cretaceous LIPs, and their magmatic carbon content, are needed in order to better constrain the rate of CO 2 emissions from LIP volcanism and its importance in determining their impact on Earth's environment/biosphere.Such data are especially required for oceanic LIPs, which have been investigated by relatively few geochronological or geochemical studies thus far (but see compilations by Coffin et al. 2002;Kasbohm et al. 2021;Davidson et al. 2023;Jiang et al. 2023).Additional studies of trace-metal and micronutrient distribution from oceanic plateaus are also needed in order to better understand the potential impact of this process in triggering marine anoxia during times of submarine LIP volcanism.In particular, it is unclear whether the distribution of volcanic micronutrients directly into seawater could have enhanced primary productivity and water-column deoxygenation on a global scale.If this effect was largely limited to regions close to the LIP source, anoxia at more distal sites may have depended on terrestrial runoff, upwelling of deepwater nutrients, and/or the development of marine stratification (see Jenkyns 2010;Föllmi 2012).
Finally, the sheer number of Cretaceous LIPs remains something of a mystery.As outlined at the start of this chapter, there were at least seven highly voluminous LIPs during the Cretaceous Period, which have been associated with intervals of profound environmental and/or biosphere disturbance.This represents a significantly higher total than for any other geological period in the last 300 Myr since the formation of Pangaea, even accounting for the longer duration of the Cretaceous v. other periods (Fig. 7).Moreover, seven major Cretaceous LIPs represents a conservative estimate.Some authors have proposed that other extensive igneous formations emplaced geologically rapidly during the Cretaceous represent additional LIPs, or that magmatic areas considered here to be part of the Kerguelen or Greater Ontong-Java plateaus were actually distinct provinces (see compilations by Torsvik 2019;Ernst et al. 2021).It may also be noteworthy that the Cretaceous featured the most oceanic LIPs of any known time in the Phanerozoic.This peak may result partly from preservation bias given the absence of most pre-Cretaceous ocean crust, but there was no comparable level of oceanic LIP formation in the Cenozoic, for which the available geological record is also good.
Why the Cretaceous was apparently 'a time of LIPs' is unclear.The middle part of the period was also marked by a long interval without magnetic reversals (the Cretaceous Normal superchron), and it has been hypothesized that both the LIPs and the superchron resulted from atypical mantle dynamics, such as the formation of a superplume or significant changes in subduction rates (e.g.Larson 1991;Courtillot and Olson 2007;Yoshimura 2022).Alternatively, this superplume may have been an inevitable consequence of the formation and subsequent break-up of the Pangaean supercontinent (Vaughan and Storey 2007).While these models are speculative, the fact that the Cretaceous was marked by an unusually high number of LIPs, extensive and fastspreading mid-ocean ridges, enhanced kimberlite emplacement and even the formation of komatiites (the only clear Phanerozoic example of this lithology), suggest that Earth's mantle conditions were highly abnormal during that time (e.g.Kerr 2005;Seton et al. 2009;Heaman et al. 2019).Given that numerous supercontinent cycles have taken place in the Proterozoic and Phanerozoic eras, it is highly unlikely that there has been only one such spell of high LIP formation in Earth's history.This conclusion raises the tantalising possibility that the Cretaceous Period marked the most recent example of a key mantle phenomenon inherent in Earth's longterm tectonic cycling, which subsequently resulted in profound disturbances to Earth's surface climate, environment, and biosphere.

Summary
Large igneous province volcanism has been hypothesized to be the primary trigger of several intervals of environmental perturbation during the Cretaceous Period, both through eruption of continental flood basalts and the submarine emplacement of oceanic plateaus.This study has generated new mercuryconcentration and osmium-isotope datasets to reconstruct the temporal relationships between volcanism, climate warming and environmental change during the Early Aptian oceanic anoxic event (OAE 1a).By incorporating these results into an overview of sedimentary, geochemical and palaeontological evidence of episodes of environmental and biotic change, and the geological, geochronological and geochemical records of LIP activity, the following conclusions are drawn: (1) Climate warming associated with OAE 1a was probably initiated by magmatic CO 2 emissions prior to the onset of the OAE sensu stricto, although volcanic activity and global temperatures did not reach their maximum until the middle part of OAE 1a.The initial carbon was probably sourced from submarine volcanism during formation of the Greater Ontong-Java Plateau.As well as potentially initiating climate warming, this magmatic CO 2 drove biotic stress amongst benthic fauna (dasycladalean algae, larger benthic foraminifera, corals, rudists) and dwarfism in calcareous nannoflora.An additional nonmagmatic source such as methane clathrate or thermogenic emissions associated with the High Arctic LIP probably played a key role in further perturbing the global carbon cycle at the onset of the OAE.(2) The Early Aptian and latest Cenomanian OAEs (OAEs 1a and 2, respectively) represented the episodes of most severe climatic, environmental, and biota stress during the Cretaceous apart from the end-Cretaceous mass extinction (for which the role of volcanism is contested).OAEs 1a and 2 were also coeval with the times of most intense LIP activity and output of mafic trace metals (potentially .29 times that timespan.This compilation includes the Triassic Wrangellia LIP, despite its preservation as relatively small areas of accreted igneous material obducted on to the western margin of North America, since it was probably a large oceanic plateau originally (Kerr 2014, and references therein).Silicic LIPs (see Bryan and Ferrari 2013) are not included, as they are thought to form over very long time intervals, and are less clearly linked to mantle plume activity.LIP, Large igneous province.
including key micronutrients) to the global ocean from oceanic plateaus.Thus, while severe environmental change and mass extinctions are more typically linked with continental LIPs, oceanic plateaus also had the capacity to profoundly impact Earth's surface environment and ecosystems during the Cretaceous Period.
(3) The more limited biosphere impact of some Cretaceous LIPs, such as the Paraná-Etendeka and Kerguelen Plateau, may have resulted from their more protracted formation and lower eruption rate.Kerguelen volcanism was apparently less intense than that of the Greater Ontong-Java and Caribbean/High Arctic LIPs associated with OAEs 1a and 2, potentially leading to lower carbon and tracemetal micronutrient output.Paraná-Etendeka magmas may also have been volatile depleted and did not intrude carbon-rich country rocks.Further work is needed to precisely constrain the eruptive histories and carbon outputs of Cretaceous LIPs (especially oceanic plateaus), and the potential dispersal of micronutrients from submarine volcanism.( 4) The high number of LIP events in the Cretaceous compared with times before or afterwards resulted in Earth's carbon cycle and surface environment being frequently perturbed to an extent not documented for any other geological period in at least the last 300 Myr.Why there was such a high rate of LIP formation (and magmatism in general) during the Cretaceous remains unclear.Further work is needed to determine the cause, and whether there were earlier intervals in Earth's history featuring enhanced Large Igneous Province emplacement, or if the Cretaceous represented a unique 'Age of LIPs'.

Fig. 1 .
Fig. 1.(a) Occurrence of large igneous province (LIP) volcanism (vertical orange bars), and episodes of major environmental change (horizontal grey bars), during the Cretaceous Period.Ages and durations of volcanic activity are primarily based on Jiang et al. (2023) (the range of the Ontong-Java is extended into the Aptian to account for palaeontological and geochemical evidence of activity at that time).Ages and durations of major events from Wagreich (2012), Jenkyns (2010), Schulte et al. (2010), Eldrett et al. (2015), Cavalheiro et al. (2021), Matsumoto et al. (2022) and Martinez et al. (2023).(b) Palaeogeographic maps show the reconstructed continental positions during the Weissert Event, oceanic anoxic events (OAEs) 1a, 1b and 2, and the Cretaceous-Paleogene (KPg) transition, with the LIPs that have been associated with each event indicated in red (with the Chicxulub impact crater also indicated by the yellow star for the KPg).Sources: Palaeogeographical reconstructions are modified from the following sources: Weissert Event from Möller et al. (2020) and Charbonnier et al. (2020b); OAE 1a from Percival et al. (2021b) and Kocsis and Scotese (2021); OAE 1b from Kocsis and Scotese (2021); OAE 2 from Du Vivier et al. (2015); and the KPg from Claeys et al. (2002).

Fig. 2 .
Fig. 2. Palaeogeographic map of the Barremian-Aptian interval, adapted from Percival et al. (2021b), after van Breugel et al. (2007) and Giraud et al. (2018).The reconstructed locations of the HALIP and G-OJP are shown in red.Sites investigated in this study for mercury and osmium data shown by white circles: (a) DSDP Site 398; (b) Alstätte-1.Sites previously investigated using both proxies are shown by black circles (c-e), just for osmium isotopes with black squares (f-g), and just for mercury by black triangles (h-o).

Fig. 3 .
Fig. 3. Stratigraphic trends in δ 13 C, total organic carbon (TOC) and sea-surface temperature (SST) based on TEX 86 , 187 Os/ 188 Os (i) , [Os] (i) , Hg, Hg/TOC and Hg/TS for (a) DSDP Site 398 and (b) Alstätte-1.Vertical scales are in metres.The stratigraphic extent of the OAE 1a level (here defined as strata spanning the C3-C6 segments) is marked by the horizontal grey bars.TS, Total sulfur.Sources: For DSDP Site 398, the lithological column is from Jenkyns (2018), the biostratigraphic and δ 13 C information are from Li et al. (2008); the SST trends are from Naafs and Pancost (2016); all other data from this study.For Alstätte-1, the lithological, biostratigraphic, δ 13 C, TOC and SST information are fromMutterlose et al. (2014); all other data are from this study.Sea-surface temperature trends based on both the widely applied TEX 86 H calibration ofKim et al. (2010) and the deep-time analogue outlined byTierney and Tingley (2014) are presented, as calculated byNaafs and Pancost (2016).A lower reliability limit of 0.2 wt% for TOC follows the protocol ofGrasby et al. (2016), and is also applied here for TS.

Fig. 4 .
Fig. 4. Stratigraphic correlation of OAE 1a geochemical records from (a) DSDP Site 398, (b) Cismon, and (c) Gorgo a Cerbara.Vertical scales are in metres.The stratigraphic extent of the OAE 1a level (here defined as strata spanning the C3-C6 segments) is marked by the horizontal grey bar.SST, Sea-surface temperature; TOC, total organic carbon.DSDP Site 398 is sourced as for Figure 3. Sources for other sites: Cismon δ 13 C and δ 18 O data, and the stratigraphic position of the nannoconid crisis, are from Erba et al. (2010);

Fig. 5 .
Fig. 5. (a) compiled 187 Os/ 188 Os and (b) 87 Sr/ 86 Sr trends from Cretaceous stratigraphic records.The strontium-isotope trend is redrawn from the compilation in McArthur et al. (2020).Osmium-isotope data sourced as shown.Changes in biospheric genera throughout the Cretaceous as compiled by Steuber et al. (2023) shown for (c) larger benthic forams, (d) rudists, and (e) corals, (f) planktonic taxa.Age models for the Os datasets are taken from the respective studies where possible (accounting for updated Stage boundary absolute ages); where no age model exists, a constant sedimentation rate relative to the start and end of the event is assumed.

Fig. 6 .
Fig. 6.Simplified schematic of the climatic and environmental influence, and output of elements used as proxies of LIP activity, during the Cretaceous for (a) continental flood basalts and (b) oceanic plateaus, OMZ, Oxygen Minimum Zone.

Fig. 7 .
Fig. 7. Summary of the number of mafic LIPs (continental flood basalt provinces or oceanic plateaus) emplaced during each geological period since c. 300 Ma, the onset of the Permian.For each period, the total number of mafic LIPs spanning an area of c. 0.5 Mkm 2 or more is shown (based on the recent compilation ofErnst et al. 2021), together with the ratio between the number of LIPs and proportional duration of the Period relative to the 79 Myr long Cretaceous, i.e. the Jurassic (56.3 Myr) is 0.71 times the duration of the Cretaceous, whereas the Neogene (23.03 Myr) is only 0.29 times that timespan.This compilation includes the Triassic Wrangellia LIP, despite its preservation as relatively small areas of accreted igneous material obducted on to the western margin of North America, since it was probably a large oceanic plateau originally(Kerr 2014, and references therein).Silicic LIPs (seeBryan and Ferrari 2013) are not included, as they are thought to form over very long time intervals, and are less clearly linked to mantle plume activity.LIP, Large igneous province.