# Abrupt episode of mid-Cretaceous ocean acidification triggered by massive volcanism

Matthew M. Jones, Bradley B. Sageman, David Selby, Andrew D. Jacobson, Sietske J. Batenburg, Laurent Riquier, Kenneth G. Macleod, Brian T. Huber, Kara A. Bogus, Maria Luisa G. Tejada, et al.

<figure><img src="image_1.png" /></figure>

## To cite this version:

Matthew M. Jones, Bradley B. Sageman, David Selby, Andrew D. Jacobson, Sietske J. Batenburg, et al.. Abrupt episode of mid-Cretaceous ocean acidification triggered by massive volcanism. Nature Geoscience, 2023, 16, pp.169-174. 10.1038/s41561-022-01115-w. insu-04139380

HAL Id: insu-04139380

https://insu.hal.science/insu-04139380v1

Submitted on 23 Jul 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_2.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
21 **Abstract**

Large igneous province volcanic activity during the mid-Cretaceous approximately 94.5 million years ago triggered a global-scale episode of reduced marine oxygen levels known as Oceanic Anoxic Event 2. It has been hypothesized that this geologically rapid degassing of volcanic carbon dioxide altered seawater carbonate chemistry, affecting marine ecosystems, geochemical cycles, and sedimentation.

Here, we report on two sites drilled by the International Ocean Discovery Program offshore of southwest Australia that exhibit clear evidence for suppressed pelagic carbonate sedimentation in the form of a stratigraphic interval barren of carbonate, recording ocean acidification during the event. We then use the osmium isotopic composition of bulk sediments to directly link this protracted ~600-kiloyear shoaling of the marine calcite compensation depth to the onset of volcanic activity. This decrease in marine pH was prolonged by biogeochemical feedbacks in highly productive regions that elevated heterotrophic respiration of carbon dioxide to the water column. A compilation of mid-Cretaceous marine stratigraphic records reveals a contemporaneous decrease of sedimentary carbonate content at continental slope sites globally. Thus, we contend that changes in marine carbonate chemistry are a primary ecological stress and important consequence of rapid emission of carbon dioxide during many large igneous province eruptions in the geologic past.
37 **Main Text**

38 **Introduction**

39 Episodes of ocean acidification in Earth history provide context for predicting the future consequences
40 of anthropogenic CO2 emissions. Potential geologic acidification events have been identified during the
41 Paleocene-Eocene Thermal Maximum (PETM), end-Permian mass extinction, Triassic-Jurassic
42 boundary, and Mesozoic oceanic anoxic events¹⁻⁵. However, some geologic episodes of prolific
43 volcanic CO2 release may have occurred too slowly to significantly decrease marine pH⁶ given that
44 emission rates must outpace fluxes of alkalinity from silicate weathering and seafloor carbonate
45 dissolution, which buffer ocean carbonate chemistry on timescales of 10's of kiloyears. Therefore,
46 precisely resolved stratigraphic records are key to assessing candidate ocean acidification events. Here,
47 we evaluate the long-standing hypothesis that ocean acidification accompanied the mid-Cretaceous
48 Oceanic Anoxic Event 2 (OAE2, ~94 Ma), ultimately resulting in turnover of marine fauna and altered
49 marine geochemical cycles⁷⁻⁹.

50
51 Ocean acidification's first order effects on marine carbonate chemistry are well understood due to the
52 relationship between pH, alkalinity, carbonate speciation, and carbonate mineral saturation states
53 (Ωmineral)¹⁰. However, erosive carbonate dissolution and sediment burndown commonly occur during
54 ocean acidification events, obscuring direct chemostratigraphic proxy records of changes in Ω and
55 pH⁵,¹¹. The stratigraphic record does, however, robustly preserve an alternative marker of ocean
56 acidification in the form of the calcite compensation depth (CCD) – the depth below which pelagic
57 sediments contain low carbonate contents (<20 wt.%) due to calcite undersaturation and dissolution in
58 colder CO2-rich bottom waters. Rapid addition of CO2 to the marine realm decreases calcite saturation
59 and forces the CCD to shallow, as is occurring in the modern ocean due to uptake of anthropogenic
60 CO₂¹². Thus, paleoacidification events may be identified from punctuated carbonate-barren intervals in 61 pelagic marine strata¹,¹⁰,¹³.

62
63 During OAE2, widespread anoxia and euxinia in marine bottom waters led to the deposition of black
64 shales in many, though not all, marine basins globally¹⁴,¹⁵. The organic carbon burial event is defined in
65 stratigraphic records by a positive carbon isotope excursion lasting at least 600 kiloyears across the
66 Cenomanian-Turonian stage boundary ¹⁶,¹⁷. A volcanic trigger for OAE2 is suggested by increased
67 contents of volcanogenic elements in marine sediments, and isotopic excursions indicate enhanced
68 activity from large igneous provinces (LIPs) emplaced around the time of the event, such as the
69 Caribbean and/or High Arctic LIPs⁷,¹⁸ (Fig. 1a). Particularly strong evidence for this trigger is a shift to
70 unradiogenic (i.e., mantle-like juvenile volcanic) initial osmium isotope ratios (Os) in sediments at
71 many marine sites tens of thousands of years prior to the beginning of OAE2¹⁷,¹⁹,²⁰. Kerr (1998)
72 hypothesized that the Caribbean LIP could have influenced marine carbonate chemistry based on the
73 volume of basalts emplaced (~4x10⁶ km³) and a recent estimate of the total CO₂ degassed during OAE2
74 ranges from 14,000 to 46,000 Pg of carbon ²¹. Additionally, pioneering sedimentological observations
75 of deep-sea cores through the OAE2 interval also noted carbonate barren intervals at many pelagic
76 sites, portending a shallowed CCD²²⁻²⁵. Although OAE2 has been implicated in general considerations
77 of acidification through geologic time⁴,⁸, only one recent Ca isotope study has offered direct
78 geochemical evidence for altered calcite saturation states during OAE2⁹. To comprehensively test the
79 OAE2-ocean acidification hypothesis, here we report osmium geochemical proxy evidence for the
80 timing and magnitude of LIP volcanic activity in relation to changes in the CCD preserved in sediment
81 cores from the Mentelle Basin offshore southwest Australia. These osmium data are combined with a
82 compilation of carbonate content records from marine sites to resolve global trends in the CCD during
83 OAE2.

84
85 **CCD shoaling in the Mentelle Basin.** Expedition 369 of the International Ocean Discovery Program
86 (IODP) cored two relatively conformable pelagic successions of OAE2 at Sites U1513 and U1516 in
87 the Mentelle Basin in 2017<sup>26,27</sup> (**Fig. 1a, Extended Data Fig. 1**). At the time of OAE2, these sites were
88 located at ~60°S in the southeastern Tethys Ocean on the subsiding eastern flank of the Naturaliste
89 Plateau at an estimated water depth of ~1100 m<sup>28</sup>. At Site U1516, changing carbonate content<sup>26</sup> and Ca
90 XRF core scanning data reveal a shoaling of the CCD between 471.0-467.3 m rCCSF (revised core
91 composite depth below sea floor) (**Fig. 2**). This 3.7 m thick darker and carbonate-free stratigraphic
92 interval corresponds to much of the encompassing OAE2 interval as constrained by microfossil
93 biostratigraphy and bulk carbonate carbon isotope (δ<sup>13</sup>C<sub>carb</sub>) chemostratigraphy<sup>27</sup>. The absence of
94 carbonate minerals means some details of the δ<sup>13</sup>C<sub>carb</sub> excursion (CIE) through the event are not
95 preserved. However, the base and termination of the event's CIE are apparent in the bracketing
96 carbonate-bearing intervals (**Extended Data Fig. 2**). Based on a floating orbital timescale derived from
97 bandpass filtering of Fe XRF core scanning data, the duration of the carbonate compensation event is
98 approximately six short eccentricity cycles, or ~600 kiloyears, which approaches the total duration of
99 the OAE2 CIE at Site U1516 of approximately seven short eccentricity cycles, or ~700 kiloyears [see
100 Methods and **Extended Data Figs. 3-4**].

101
102 Extremely unradiogenic Osᵢ values (<0.2) and high <sup>192</sup>Os contents in strata across the base of the OAE2
103 interval at Site U1516 provide evidence for volcanic activity (**Fig. 2**). The main unradiogenic Osᵢ
104 excursion (to juvenile volcanic end-member values) and spike in <sup>192</sup>Os contents to ~1,800 parts per
trillion at Site U1516 both culminate between 471-470 m rCCSF, marking major LIP volcanic activity
associated with the onset of OAE2 (Fig. 2). This timing is notable because the initial shoaling of the
CCD in the Mentelle Basin directly coincides with the geologically rapid intensification of LIP
volcanism (Fig. 3). Based on inverse modeling of the Osᵢ and ¹⁹²Os profiles, the LIP Os flux to the
marine reservoir exceeded pre-event baseline mantle/volcanic fluxes by at least 30x during peak
intensity and lasting at least 60 kiloyears through the onset of OAE2 [see Methods and **Extended Data
Tables 1-2**]. If continental weathering fluxes increased as much as 80% during OAE2 as has been
suggested recently²⁹, the additional radiogenic osmium flux would necessitate an even greater LIP
volcanic Os flux approaching 50x baseline levels to recreate the trends in observed Osᵢ curves
(Extended Data Fig. 5). Combined, these observations link the emission of volcanically derived CO₂
to the initial shoaling of the CCD at Site U1516 and ocean acidification during OAE2.

116

More radiogenic (higher) Osᵢ datapoints and low ¹⁹²Os contents occur above 469.6 m rCCSF at Site U1516, signaling a waning of LIP volcanic activity in the upper interval of OAE2 roughly 180±60 kiloyears after the start of the event (Figs. 2-3). Five unradiogenic Osᵢ datapoints are intermittently present in this interval (469.4-466.0 m rCCSF) and could possibly represent punctuated renewed LIP eruptions. However, this seems unlikely as the horizons do not record concomitant spikes in ¹⁹²Os contents and no previously published marine Osᵢ chemostratigraphic records (n>15) document global-scale LIP volcanic activity in the upper OAE2 or post-OAE2 intervals¹⁷,¹⁹,²⁰,³⁰ (**Extended Data Fig. 2**). Thus, it seems more likely these upper unradiogenic Osᵢ points reflect either seafloor reworking of continental slope sediments or record a localized source of volcanic activity, such as the Kerguelen Plateau, with a smaller Os flux that did not mix on a global scale³¹.

127
**Global carbonate sedimentation during OAE2.** Trends in marine carbonate sedimentation during OAE2 from a global compilation fall into several categories (Fig. 1b). Eight sites from abyssal, continental rise, or lower continental slope settings preserve less than 20 wt.% CaCO₃ before, during, and after the OAE2 interval, indicating they remained consistently below or near the CCD. Thirteen sites almost exclusively from epicontinental seas preserve 40-95% CaCO₃ prior to and during OAE2 and are therefore interpreted to have remained consistently above the CCD. Of those sites, seven record increases in carbonate content during OAE2. Finally and notably, carbonate contents at thirteen sites mainly from continental slope settings decrease to sub-CCD wt.% CaCO₃ values during OAE2 indicating a pronounced shoaling of the CCD in multiple basins globally. The OAE2 interval is relatively condensed at a majority of these shoaled pelagic sites, reflecting reduced carbonate accumulation³².

139

**Ocean Acidification during OAE2.** The observations from our osmium and compiled sedimentological data for OAE2 suggest that ocean acidification was an additional fundamental paleoceanographic response to this episode of LIP volcanism. Acidification resulted in a shoaling of the CCD recorded at continental slope sites globally. In shallower settings through OAE2 though, carbonate deposition progressed and many carbonate platforms persisted ³³ despite partial drowning and stress from nutrient loading³⁴ (Fig. 1b). Increases in carbonate content were particularly substantial in epicontinental seaways like the North American Western Interior Seaway, where a pronounced global transgression decreased the relative contribution of siliciclastic sedimentation³⁵ (Fig. 1a). Arthur et al. (1987) attributed carbonate barren deep sea intervals during OAE2 to a shift in carbonate sedimentation from the pelagic realm to the vastly expanded epicontinental seaways. Models invoking sea level rise have successfully recreated prolonged shifts in the CCD through the Middle Eocene
Climatic Optimum “MECO”, a rare geologic example of a long-lasting (~500 kiloyears) carbonate compensation event similar to OAE2³⁶. Increased shallow marine calcification and consequent decreased alkalinity would have worked in concert with LIP CO₂ emissions to promote shoaling of the CCD during OAE2 in slope marine settings like Site U1516.

The ~600-kiloyear shoaling of the CCD observed during OAE2 at Site U1516 is much longer than predictions of conventional models for marine carbonate chemistry tailored to other geological episodes of voluminous CO₂ release⁴. During the PETM for example, the delivery of alkalinity from seafloor carbonate dissolution and silicate weathering ends surface ocean acidification events within 10s of kiloyears⁴. To reconcile model results and observational data for OAE2, we hypothesize that the remineralization of ample organic carbon sinking from highly productive surfaces waters produced a significant flux of dissolved CO₂ to marine bottom waters. In the lower OAE2 interval at Site U1516, geochemical proxies—such as biogenic Si³⁷ and Si enrichment factor (SiEF) values—increase near three thin laminated black shale beds (4-12 cm-thick beds labeled “a” “b” “c” in Fig. 2), indicating intermittently enhanced primary productivity. In the upper OAE2 interval, more sustained increases in biogenic silica³⁷, SiEF, and BaEF, and abundant radiolaria²⁷, indicate consistently higher productivity levels, which would have led to higher heterotrophic respiration rates and increased dissolved CO₂ in the water column as has been noted previously at OAE2 sites with high total organic carbon (TOC)³⁸,³⁹. An intense CO₂ flux from LIP volcanism appears to have triggered ocean acidification, but marine biogeochemical feedbacks that enhanced primary productivity appear to have sustained a shallower CCD for hundreds of kiloyears during OAE2. Differences between model and observation may also reflect that models employ a combination of burndown and continental chemical weathering (chemical compensation) as the dampening mechanism to ocean acidification, whereas Ca isotope data for OAE2⁹
and for the older OAE1a⁴⁰ point to additional influence from reduced carbonate production (biological compensation)⁴¹.

These marine biogeochemical factors, along with epicontinental carbonate deposition discussed above, would explain the longer duration CCD anomaly during OAE2 and possibly other OAEs², especially in comparison to the PETM, an event without pervasive black shale deposition and a positive CIE¹.

Notably, the Site U1516 carbonate record also exhibits a recovery of carbonate deposition following OAE2 (Fig. 2). This final deepening of the CCD was likely linked to the factors that ultimately terminated the event, such as alkalinity buildup from reduced carbonate production, increased continental weathering, and/or a decline in global primary productivity, as is observed locally at Site U1516 in waning biogenic silica and BaEF values above ~467.3 m rCCSF.

**Os flux relation with CO₂ release.** The relationship between an increasing flux of Os and the emission of CO₂ as LIP activity progressed is poorly constrained and to some degree depends upon the lithology of the intruded host rock⁴². However, if the volcanic Os and C fluxes scaled linearly, a >30x increase in CO₂ emission above background mantle/volcanic degassing rates would begin to approach modern anthropogenic CO₂ emission rates, which are ~80-270x above background volcanic fluxes and are rapidly altering marine carbonate chemistry¹²,⁴³. In the mid-Cretaceous, a scenario of even lower CO₂ release, though of similar order of magnitude and rate, would have strongly affected ocean chemistry due to lower carbonate ion concentrations¹⁰, and would have likely decreased pH and carbonate mineral saturation, as is inferred from Ca isotope records⁹ and the shoaled CCD.
Typically, negative CIEs occur during candidate ocean acidification events in the geologic record (e.g., PETM or OAE1a) due to the release of CO₂ with low δ¹³C values from volcanogenic carbon reservoirs (~-5‰ PDB or biogenic (<-20‰ PDB - Peedee Belemnite scale)) ¹,⁴⁴. Near the onset of OAE2, some sites record a negative CIE, possibly reflecting the emission of CO₂ that shoaled the CCD during the event ¹⁸,²¹. These records are somewhat rare though and could alternatively be explained by local environmental changes in depositional settings that also influence δ¹³C values. Somewhat puzzlingly and despite exhibiting a shoaling of the CCD and one of the most severe volcanogenic Osᵢ excursions of the Phanerozoic, OAE2 is dominantly characterized by a positive CIE ²⁵. Recently, Mason et al. (2017)⁴⁵ showed that volatile-rich volcanic settings which assimilate crustal carbonate also release CO₂ with high δ¹³C values (~ -1‰ PDB). Thus, if the volcanism associated with OAE2 erupted in a carbonate-rich setting like the Caribbean (Fig. 1a), we hypothesize that CO₂ would have been released from an isotopically heavy pool of C, generating the observed carbonate compensation event independent of a pronounced negative CIE.

**Biotic effects of acidification.** Marine deoxygenation is hypothesized to have accelerated marine macro- and micro-faunal turnover rates during OAE2⁴⁶,⁴⁷. Suppressed calcite and aragonite saturation states due to LIP volcanism would also have been a compounding stressor to marine ecosystems. Laboratory studies of modern calcifying organisms record significant declines in calcification rates at Ωaragonite less than three⁴⁸. Such conditions, where surface waters were not undersaturated and would not experience outright dissolution, but where Ω had declined, are consistent with the epicontinental carbonate sedimentation trends of OAE2. Biotic indicators of ocean acidification, such as dwarfing, thinning, and fragmentation of foraminifera and calcareous nannofossils, occur strikingly during the PETM¹¹. Some, but not all, of these features occur during OAE2, in particular, the dwarfing of calcitic
218 microfossils<sup>49,50</sup>, which correlates with Ca isotope evidence for reduced precipitation rates<sup>51</sup>. These 219 morphological shifts, along with the extinction of the deep-dwelling, keeled foraminifera *Rotalipora* 220 *cushmani* early in OAE2, are consistent with a scenario of LIP volcanic CO<sub>2</sub> emission, decreased Ω, 221 global warming, and biotic extinctions<sup>52</sup>. Similarly, records of demise in the mid-Cretaceous rudist 222 reefs<sup>53</sup> along with turnover of aragonitic ammonite<sup>46</sup> and calcitic microfossil<sup>47</sup> taxa may in part record 223 selection following changes in marine carbonate chemistry during OAE2. 224

225 **Implications and unresolved aspects of mid-Cretaceous ocean acidification.** In summary, osmium 226 chemostratigraphic data and compiled sedimentological observations of carbonate content from OAE2 227 indicate that CO₂ emitted from intense LIP volcanic activity triggered a shoaling of the CCD, which 228 was likely prolonged by subsequent marine biogeochemical feedbacks and sea level transgression. We 229 conclude that ocean acidification represented a notable component of environmental deterioration 230 during OAE2, and perhaps other Mesozoic OAEs, with volcanogenic and biogeochemical fluxes of C 231 to the global ocean that exceeded thresholds for buffering from silicate weathering. In particular, these 232 results raise questions about the temporal relation between shifts in surface water carbonate chemistry 233 and longer duration changes in the CCD during LIP episodes in Earth history. More quantitative depth 234 constraints on the shift of the CCD during OAE2 may help refine the total mass of pelagic carbonate 235 dissolved and, in turn, the amount of CO₂ emitted from LIP sources. However, the relative roles of 236 dissolution versus lower calcification rates<sup>40</sup> in driving the CCD shallower during OAEs must still be 237 quantitatively discerned. As to the source of the CO₂, observations of carbon isotope profiles may 238 fingerprint which volcanic province or provinces were linked to the event, yet the geochronology and 239 eruptive tempo of many mid-Cretaceous LIPs is still coarsely resolved<sup>54</sup> compared with LIPs from 240 other intensely studied intervals, such as the Deccan Traps at the Cretaceous-Paleogene Boundary<sup>55</sup>. In-
241 situ verification of the precise setting and timing of volcanic flows associated with ocean acidification
242 during OAE2 remains a grand challenge in paleoceanography.

243
244 These deep-time findings highlight the role that marine nutrient loading, along with elevated productivity and subsequent anoxia, play in exacerbating and prolonging carbonate mineral undersaturation in marine settings. Continued study of how benthic and planktic communities responded to a decrease in calcite saturation state during OAE2 should provide predictive insights into the vulnerabilities and resiliencies of modern ecosystems as marine pH continues to decline. Additional investigations employing existing ocean acidification-sensitive isotope proxies, such as B⁴, and emerging ones, including Ca⁹ and stable Sr⁴⁰, may further critically test for surface water fluctuations in pH, calcite/aragonite saturation, and calcification stress.

252

253 Acknowledgements

254 We are grateful for the dedication of the IODP Expedition 369 Science Party, technicians, and crew in coring Site U1516; analytical support of Antonia Hofmann, Chris Ottley and Geoff Nowell at Durham University; and assistance of Brian Levay with XRF core scanning. This research was funded by a U.S. Science Support Program (USSSP) Post-Expedition Award to MMJ and box modeling and carbonate compilations represent doctoral research of MMJ, partly funded by National Science Foundation Grant 1338312 to BBS. MMJ acknowledges support from the Smithsonian Peter Buck Fellowship; DS acknowledges the TOTAL endowment fund; LR acknowledges support from IODP-France.

261

262 **Author Contributions**
Study conceptualization – M.M.J., B.B.S., A.D.J.; geochemical analyses – M.M.J., D.S., S.J.B., L.R., K.G.M., K.A.B.; field work and core collection – B.T.H., R.W.H., K.A.B., M.L.G.T., J.K., S.J.B., L.R., K.G.M., M.M.J.; original manuscript draft – M.M.J., B.B.S., D.S., A.D.J., K.G.M., S.J.B., L.R.; modeling analyses – M.M.J., S.J.B., M.L.G.T., J.K.; manuscript editing – all.

267

**Competing Interests** The authors declare no competing interests.

269

## 270 FIGURE CAPTIONS

**Figure 1:** A – Late Cretaceous paleomap of sites compiled for carbonate sedimentation through Oceanic Anoxic Event 2 (OAE2). See **Extended Data Table 3** for data sources. Marker colors correspond to paleobathymetric setting and red rings denote sites where a shoaling of the calcite compensation depth (CCD) is observed. Yellow star denotes IODP Site U1516. Purple regions represent Caribbean (CLIP), High Arctic (HALIP), and Kerguelen (KLIP) large igneous provinces. Map adapted from²⁰. WIS – Western Interior Seaway. B – A global compilation of trends in sedimentary carbonate contents (wt.%) prior to and during Oceanic Anoxic Event 2 (OAE2). Text color corresponds to paleobathymetric setting. Sites with a precipitous decrease in carbonate during the event record a shoaling CCD (bottom center). Shallow water sites (bottom right) remained above the CCD throughout OAE2 with some increasing in carbonate contents in contrast to deeper water sites (see text for discussion).

**Figure 2:** Chemostratigraphy and core photos of the Oceanic Anoxic Event 2 (OAE2) interval at IODP Site U1516 (eastern Indian Ocean). LIP volcanism is recorded by a decrease in initial osmium isotope ratios (Os; red, panel D) and spike in ¹⁹²Os contents (ppt=parts per trillion; purple, panel E) near the
base of OAE2. A positive carbonate carbon isotope ($\delta^{13}C_{carb}$, panel C) excursion and nannofossil biostratigraphy<sup>27</sup> confirm this correlation (see **Extended Data Figure 4** for more details). A ~600-kiloyear shoaling of the calcite compensation depth (CCD) is preserved by the near-complete absence of carbonate minerals from ~471-467.3 m rCCSF (revised core composite depth below sea floor) based on shipboard carbonate content data<sup>26</sup> (black dots –panel A) and Ca 10 kV XRF scanning data (blue, panel A). XRF scanning data plotted from cores 3R-5R in Hole D. The bandpass filter of Fe 10 kV XRF scanning data records the short eccentricity orbital cycle of ~100 kyr and forms a floating timescale (panel B). Elemental proxies at right ($Si_{(EF)}$ - silicon enrichment factors, panel F; biogenic Si – biogenic silica<sup>37</sup>, panel F; $Ba_{(EF)}$ - barium enrichment factor, panel G) show evidence for increased productivity during OAE2, especially in the upper interval of the event. Together, these data are consistent with a volcanic trigger for ocean acidification and the shoaling of the CCD, which was prolonged by elevated productivity during OAE2.

**Figure 3.** Inverse modeling results of large igneous province (LIP) volcanic activity during Oceanic Anoxic Event 2 (OAE2; gray shading) based on $Os_i$ (red) and $^{192}$Os (purple) data from IODP Site U1516 (circles, panel B). At peak intensity in the OAE2 onset interval, input fluxes of Os from LIP sources (panel A) must be >30 times greater than the global background level of the mid-Cretaceous mantle/volcanic Os flux. The Site U1516 $Os_i$ record is consistent with smaller, yet still significant volcanic activity preceding OAE2 by several hundred kiloyears (kyr). Also plotted is a coeval Os time-series from the Angus Core in the Western Interior Basin of North America<sup>17</sup> (triangles) for a shallow marine comparison (panel B). See Methods for full modeling details.
307 1. Zachos, J. C. Rapid Acidification of the Ocean During the Paleocene-Eocene Thermal Maximum. *Science* **308**, 1611–1615 (2005).

308
309 2. Erba, E., Bottini, C., Weissert, H. J. & Keller, C. E. Calcareous Nannoplankton Response to Surface-Water Acidification Around Oceanic Anoxic Event 1a. *Science* **329**, 428–432 (2010).

310
311 3. Payne, J. L. et al. Calcium isotope constraints on the end-Permian mass extinction. *Proc. Natl. Acad. Sci.* **107**, 8543–8548 (2010).

312
313 4. Honisch, B. et al. The Geological Record of Ocean Acidification. *Science* **335**, 1058–1063 (2012).

314
315 5. Greene, S. E. et al. Recognising ocean acidification in deep time: An evaluation of the evidence for acidification across the Triassic-Jurassic boundary. *Earth-Sci. Rev.* **113**, 72–93 (2012).

316
317 6. Naafs, B. D. A. et al. Gradual and sustained carbon dioxide release during Aptian Oceanic Anoxic Event 1a. *Nat. Geosci.* **9**, 135–139 (2016).

318
319 7. Kerr, A. C. Oceanic plateau formation: a cause of mass extinction and black shale deposition around the Cenomanian–Turonian boundary? *J. Geol. Soc.* **155**, 619–626 (1998).

320
321 8. Jenkyns, H. C. Geochemistry of oceanic anoxic events: REVIEW. *Geochem. Geophys. Geosystems* **11**, n/a-n/a (2010).

322
323 9. Du Vivier, A. D. C. et al. Ca isotope stratigraphy across the Cenomanian–Turonian OAE 2: Links between volcanism, seawater geochemistry, and the carbonate fractionation factor. *Earth Planet. Sci. Lett.* **416**, 121–131 (2015).

324
325 10. Zeebe, R. E. & Tyrrell, T. History of carbonate ion concentration over the last 100 million years II: Revised calculations and new data. *Geochim. Cosmochim. Acta* **257**, 373–392 (2019).

326
327 11. Bralower, T. J. et al. Impact of dissolution on the sedimentary record of the Paleocene–Eocene thermal maximum. *Earth Planet. Sci. Lett.* **401**, 70–82 (2014).

328
329 12. Sulpis, O. et al. Current CaCO3 dissolution at the seafloor caused by anthropogenic CO2. *Proc. Natl. Acad. Sci.* (2018) doi:10.1073/pnas.1804250115.

331 13. Pälike, H. et al. A Cenozoic record of the equatorial Pacific carbonate compensation depth. *Nature* **488**, 609–614 (2012).

333 14. Schlanger, S. O. & Jenkyns, H. C. Cretaceous Oceanic Anoxic Events: Causes and Consequences. *Geol. En Mijnb.* **55**, 6 (1976).

335 15. Ostrander, C. M., Owens, J. D. & Nielsen, S. G. Constraining the rate of oceanic deoxygenation leading up to a Cretaceous Oceanic Anoxic Event (OAE-2: ~94 Ma). *Sci. Adv.* **3**, e1701020 (2017).

337 16. Sageman, B. B., Meyers, S. R. & Arthur, M. A. Orbital time scale and new C-isotope record for Cenomanian-Turonian boundary stratotype. *Geology* **34**, 125 (2006).

339 17. Jones, M. M. et al. Regional chronostratigraphic synthesis of the Cenomanian-Turonian Oceanic Anoxic Event 2 (OAE2) interval, Western Interior Basin (USA): New Re-Os chemostratigraphy and 40Ar/39Ar geochronology. *GSA Bull.* **133**, 1090–1104 (2021).

342 18. Kuroda, J. et al. Contemporaneous massive subaerial volcanism and late cretaceous Oceanic Anoxic Event 2. *Earth Planet. Sci. Lett.* **256**, 211–223 (2007).

344 19. Turgeon, S. C. & Creaser, R. A. Cretaceous oceanic anoxic event 2 triggered by a massive magmatic episode. *Nature* **454**, 323–326 (2008).

346 20. Du Vivier, A. D. C. et al. Marine 187Os/188Os isotope stratigraphy reveals the interaction of volcanism and ocean circulation during Oceanic Anoxic Event 2. *Earth Planet. Sci. Lett.* **389**, 23–33 (2014).

349 21. Papadomanolaki, N. M., van Helmond, N. A. G. M., Pälike, H., Sluijs, A. & Slomp, C. P. Quantifying volcanism and organic carbon burial across Oceanic Anoxic Event 2. *Geology* **50**, 511–515 (2022).
22. De Graciansky, P. C. et al. The Goban Spur transect: Geologic evolution of a sediment-starved passive continental margin. *GSA Bull.* **96**, 58–76 (1985).

23. Arthur, M. A. & Dean, W. E. Cretaceous paleoceanography of the western North Atlantic Ocean. in *The Western North Atlantic Region* (eds. Vogt, P. R. & Tucholke, B. E.) (Geological Society of America, 1986).

24. Herbin, J. P. et al. Organic-rich sedimentation at the Cenomanian-Turonian boundary in oceanic and coastal basins in the North Atlantic and Tethys. *Geol. Soc. Lond. Spec. Publ.* **21**, 389–422 (1986).

25. Arthur, M. A., Schlanger, S. O. & Jenkyns, H. C. The Cenomanian-Turonian Oceanic Anoxic Event, II. Palaeoceanographic controls on organic-matter production and preservation. *Geol. Soc. Lond. Spec. Publ.* **26**, 401–420 (1987).

26. Huber, B. T., Hobbs, R. W., Bogus, K. A. & the Expedition 369 Scientists. Site U1516. in *Australia Cretaceous Climate and Tectonics* (eds. Hobbs, R. W., Huber, B. T. & Bogus, K. A.) vol. 369 (International Ocean Discovery Program, 2019).

27. Petrizzo, M. R. et al. Exploring the paleoceanographic changes registered by planktonic foraminifera across the Cenomanian-Turonian boundary interval and Oceanic Anoxic Event 2 at southern high latitudes in the Mentelle Basin (SE Indian Ocean). *Glob. Planet. Change* **206**, 103595 (2021).

28. Lee, E. Y. et al. Early Cretaceous subsidence of the Naturaliste Plateau defined by a new record of volcaniclastic-rich sequence at IODP Site U1513. *Gondwana Res.* **82**, 1–11 (2020).

29. Nana Yobo, L., Brandon, A. D., Holmden, C., Lau, K. V. & Eldrett, J. Changing inputs of continental and submarine weathering sources of Sr to the oceans during OAE 2. *Geochim. Cosmochim. Acta* **303**, 205–222 (2021).
30. Li, Y.-X. et al. Enhanced ocean connectivity and volcanism instigated global onset of Cretaceous Oceanic Anoxic Event 2 (OAE2) ~94.5 million years ago. *Earth Planet. Sci. Lett.* **578**, 117331 (2022).

31. Jiang, Q., Jourdan, F., Olierook, H. K. H., Merle, R. E. & Whittaker, J. M. Longest continuously erupting large igneous province driven by plume-ridge interaction. *Geology* **49**, 206–210 (2021).

32. de Graciansky, P. C. et al. Ocean-wide stagnation episode in the late Cretaceous. *Nature* **308**, 346 (1984).

33. Elrick, M., Molina-Garza, R., Duncan, R. & Snow, L. C-isotope stratigraphy and paleoenvironmental changes across OAE2 (mid-Cretaceous) from shallow-water platform carbonates of southern Mexico. *Earth Planet. Sci. Lett.* **277**, 295–306 (2009).

34. Parente, M. et al. Stepwise extinction of larger foraminifers at the Cenomanian-Turonian boundary: A shallow-water perspective on nutrient fluctuations during Oceanic Anoxic Event 2 (Bonarelli Event). *Geology* **36**, 715 (2008).

35. Sageman, B. B., Rich, J., Arthur, M. A., Birchfield, G. E. & Dean, W. E. Evidence for Milankovitch periodicities in Cenomanian-Turonian lithologic and geochemical cycles, western interior USA. *J. Sediment. Res.* **67**, 286–302 (1997).

36. Sluijs, A., Zeebe, R. E., Bijl, P. K. & Bohaty, S. M. A middle Eocene carbon cycle conundrum. *Nat. Geosci.* **6**, 429–434 (2013).

37. Chen, H. et al. Enhanced hydrological cycle during Oceanic Anoxic Event 2 at southern high latitudes: New insights from IODP Site U1516. *Glob. Planet. Change* **209**, 103735 (2022).

38. Vogt, P. R. Volcanogenic upwelling of anoxic, nutrient-rich water: A possible factor in carbonate-bank/reef demise and benthic faunal extinctions? *GSA Bull.* **101**, 1225–1245 (1989).
397 39. Scopelliti, G., Bellanca, A., Neri, R., Baudin, F. & Coccioni, R. Comparative high-resolution chemostratigraphy of the Bonarelli Level from the reference Bottaccione section (Umbria–Marche Apennines) and from an equivalent section in NW Sicily: Consistent and contrasting responses to the OAE2. *Chem. Geol.* **228**, 266–285 (2006).

401 40. Wang, J., Jacobson, A. D., Sageman, B. B. & Hurtgen, M. T. Stable Ca and Sr isotopes support volcanically triggered biocalcification crisis during Oceanic Anoxic Event 1a. *Geology* (2021) doi:10.1130/G47945.1.

404 41. Boudreau, B. P., Middelburg, J. J. & Luo, Y. The role of calcification in carbonate compensation. *Nat. Geosci.* **11**, 894–900 (2018).

406 42. Ganino, C. & Arndt, N. T. Climate changes caused by degassing of sediments during the emplacement of large igneous provinces. *Geology* **37**, 323–326 (2009).

408 43. Gerlach, T. Volcanic versus anthropogenic carbon dioxide. *Eos Trans. Am. Geophys. Union* **92**, 201–202 (2011).

410 44. Bauer, K. W., Zeebe, R. E. & Wortmann, U. G. Quantifying the volcanic emissions which triggered Oceanic Anoxic Event 1a and their effect on ocean acidification. *Sedimentology* **64**, 204–214 (2017).

413 45. Mason, E., Edmonds, M. & Turchyn, A. V. Remobilization of crustal carbon may dominate volcanic arc emissions. *Science* **357**, 290–294 (2017).

415 46. Elder, W. P. Molluscan extinction patterns across the Cenomanian–Turonian Stage boundary in the Western Interior of the United States. *Paleobiology* **15**, 299–320 (1989).

417 47. Leckie, R. M., Bralower, T. J. & Cashman, R. Oceanic anoxic events and plankton evolution: Biotic response to tectonic forcing during the mid-Cretaceous. *Paleoceanography* **17**, (2002).
419 48. Ries, J. B., Cohen, A. L. & McCorkle, D. C. Marine calcifiers exhibit mixed responses to CO2-induced ocean acidification. *Geology* **37**, 1131–1134 (2009).

421 49. Coccioni, R. & Luciani, V. PLANKTONIC FORAMINIFERA AND ENVIRONMENTAL CHANGES ACROSS THE BONARELLI EVENT (OAE2, LATEST CENOMANIAN) IN ITS TYPE AREA: A HIGH-RESOLUTION STUDY FROM THE TETHYAN REFERENCE BOTTACCIONE SECTION (GUBBIO, CENTRAL ITALY). *J. Foraminifer. Res.* **34**, 109–129 (2004).

426 50. Faucher, G., Erba, E., Bottini, C. & Gambacorta, G. Calcareous nannoplankton response to the latest Cenomanian Oceanic Anoxic Event 2 perturbation. *Riv. Ital. Paleontol. E Stratigr. Res. Paleontol. Stratigr.* **123**, (2017).

429 51. Kitch, G. D. *et al.*, Calcium isotope ratios of malformed foraminifera reveal biocalcification stress preceded Oceanic Anoxic Event 2. *Communications Earth & Environment* (in press).

431 52. Coccioni, R., Sideri, M., Frontalini, F. & Montanari, A. The Rotalipora cushmani extinction at Gubbio (Italy): Planktonic foraminiferal testimonial of the onset of the Caribbean large igneous province emplacement? *Geol Soc Am Spec Pap* **524**, 79–96 (2016).

434 53. Philip, J. M. & Airaud-Crumiere, C. The demise of the rudist-bearing carbonate platforms at the Cenomanian/Turonian boundary: a global control. *Coral Reefs* **10**, 115–125 (1991).

436 54. Kasbohm, J., Schoene, B. & Burgess, S. Radiometric Constraints on the Timing, Tempo, and Effects of Large Igneous Province Emplacement. in *Large Igneous Provinces* 27–82 (American Geophysical Union (AGU), 2021). doi:10.1002/9781119507444.ch2.

439 55. Sprain, C. J. *et al*. The eruptive tempo of Deccan volcanism in relation to the Cretaceous-Paleogene boundary. *Science* **363**, 866–870 (2019).
442 **Astrochronologic time scale construction**

443 The Cenomanian–Turonian sediments at IODP Site U1516 display a rhythmic banding pattern of dark
444 and light lithologies (see Core Photos) that is particularly well expressed around 477 m². Here, the
445 darker bands occur at a scale of ~20 cm, and vary in darkness on a scale of 80 cm to 1 m. This may
446 reflect the influence of eccentricity-modulated precession, with individual (~20 cm scale) alternations
447 corresponding to precession cycles and the bundling corresponding to the influence of the short
eccentricity (~100 kyr) cycle.
448

449 REDFIT power spectra⁵⁶ of the Fe core-scanning XRF data of three stratigraphic intervals (the
450 carbonate-free interval and the intervals above and below) confirm the detection of these periodicities
451 **(Extended Data Figure 3)**. Additionally, they exhibit a decrease in sedimentation rate over the
452 carbonate-free interval attributed to the shoaling of the CCD. Dominant periodicities were extracted by
453 band-pass filtering in AnalySeries⁵⁷, with a frequency of 1.24 cycles/m and a bandwidth of 0.41 for the
454 intervals above and below the carbonate-free interval, and with a frequency of 1.90 cycles/m and a
455 bandwidth of 0.63 for the carbonate-free interval.
456

456

457 **Osmium isotope modeling of volcanic activity**

458 *Principles*: To quantitatively assess the magnitude and rate of LIP volcanic activity during OAE2, we
459 implemented an osmium isotope box model following the approach of Tejada et al. (2009)⁵⁸. The
460 marine Os isotopic reservoir is highly sensitive to changes in unradiogenic volcanic fluxes (¹⁸⁷/¹⁸⁸Os =
461 0.126) relative to radiogenic fluxes of continental weathering (¹⁸⁷/¹⁸⁸Os = ~1.4, global average)⁵⁹.
Moreover, the marine osmium reservoir has a short residence time (10's of kyr) and was relatively well-mixed during Mesozoic OAEs. As a result, the Osᵢ chemostratigraphic records of OAE2 from marine sites globally exhibit rapid perturbations from mantle-sourced Os originating from LIPs¹⁷,¹⁹,²⁰,³⁰,⁶⁰⁻⁶².

We use a simple inverse model to reconstruct the first-order trends in the OAE2 Osᵢ excursion and ¹⁹²Os concentration time-series at IODP Site U1516 (Figure 3). The Site U1516 Osᵢ and ¹⁹²Os time-series were generated from the site's short-eccentricity bandpass astronomical time scale, using the 'tune' function in R-package 'Astrochron' ⁶³. In addition we plot the densely sampled, conformable, and astronomically-tuned Os time series from the Angus Core in the Western Interior Seaway of North America¹⁷ for comparison. Model calculations were performed using MATLAB and R scripts available upon request.

*Parameterization:* The box model calculates the mass (M₀) and isotopic composition (R₀) of the Late Cretaceous global oceanic osmium reservoir through OAE2 based on input and output fluxes of Os with different isotopic compositions, represented in the following equations:

$$dM_o = F_{riv} + F_{mantle} + F_{cosmic} + F_{LIP} - F_{sed}$$

$$dR_o = \frac{[(F_{riv}*(R_{riv}-R_o))+(F_{mantle}*(R_{mantle}-R_o))+(F_{LIP}*(R_{LIP}-R_o))+(F_{cosmic}(R_{cosmic}-R_o))]}{M_o}$$

Guided by values in⁵⁸, the model establishes background input of Os to the ocean from fluxes of unradiogenic cosmic (Fcosmic), unradiogenic volcanic/mantle (Fmantle), and radiogenic continental weathering/riverine (Friv), along with an output sedimentary flux (Fsed) (Extended Data Table 1).

Osmium isotopic compositions are from ⁶⁴. To reconstruct the pre-OAE2 Osᵢ steady-state value of ~0.75 at Site U1516 in the mid-Cretaceous, Fmantle is set at 352 tonnes/kyr (15% less than ⁵⁸ and Friv is
set at 339 tonnes/kyr (15% higher than Tejada et al., 2009⁵⁸). Given likely changes in marine Os cycling between the modeled Early Cretaceous data⁵⁸ and the Late Cretaceous data modeled in this study, the minor offsets of our flux values with those in Tejada et al. (2009)⁵⁸ are considered reasonable. The model runs for 3000 kyr at a step of 1 kyr. For the purpose of comparing the model results to the measured sedimentary Osᵢ datasets, time zero (t=0) is aligned to the datapoint in each chemostratigraphic record that is prior to the earliest extremely unradiogenic Osᵢ value (Osᵢ < 0.3) indicative of intense LIP activity. The model is permitted to run for several hundred kiloyears prior to the oldest measured Os datapoint to reach a steady state.

Large igneous province perturbation: The model also incorporates a perturbation flux representing the flux of unradiogenic Os to the global ocean from LIP volcanic activity (FLIP) during OAE2, presumably from a submarine setting. In an inverse approach and holding all background Os input fluxes constant (Friv, F_mantle, F_cosmic), the FLIP was iteratively adjusted at 11 control points through time to generate a marine ¹⁸⁷/¹⁸⁸Os composition that approximated the general trends in U1516 Os time-series (Fig. 3; Extended Data Table 2). FLIP values were linearly interpolated between these control points using the ‘ interp1’ function in MATLAB to create a continuous perturbation scenario, varying through time. Highly unradiogenic Osᵢ datapoints in the upper OAE2 to post-OAE2 intervals at Site U1516 were interpreted as likely unrepresentative of a globally correlative increase in FLIP for several reasons (see main text). Therefore, iterative modeling of LIP volcanic activity did not attempt to recreate these anomalies.
Samples from IODP Site U1516 were analyzed for rhenium-osmium geochemistry (n=28; 463.00 – 480.04 m rCCSF) to establish chemostratigraphic markers of LIP volcanic activity associated with OAE2. Sampling depth resolution ranged from 17-205 cm/sample spanning the Cenomanian-Turonian boundary stratigraphic interval, with higher sampling density in the OAE2 interval. Data are reported in **Supplementary Table 1**.

Prior to analysis at Durham University, samples were powdered in ceramic containers using high-purity crushing techniques at Northwestern University and the University of Michigan. The ceramic containers were cleaned using Ottawa sand, then washed and finally rinsed with ethanol. The prepared powders were analyzed in the Source Rock and Sulfide Geochronology and Geochemistry Laboratory at Durham University utilizing isotope dilution negative ion mass spectrometry⁶⁵. In brief, sample powders (0.3–1.0 g) were spiked with a mixed ¹⁸⁵Re+¹⁹⁰Os tracer solution and digested in sealed Carius tubes with 8 mL of 0.25 g/g CrO₃ in 4N H₂SO₄ for ~48 hours at 220°C, principally leaching hydrogenous Re and Os (i.e., carbonates and organic matter). The Os fraction was isolated and purified via chloroform extraction with back reduction into HBr and CrO₃·H₂SO₄-HBr microdistillation. The Re fraction was isolated via NaOH-acetone extraction and anion chromatography. Isotopic ratios of samples and solution standards (Re STD and DROs) were measured on a Thermo Triton thermal ionization mass spectrometer (TIMS) in negative ionization mode in the Arthur Holmes Laboratory at Durham University. Running average values for ¹⁸⁷Os/¹⁸⁸Os and ¹⁸⁷Re/¹⁸⁵Re solution standards to the time of these analyses (10/2019) were 0.16094±0.00050 (1 σ; n = 700) and 0.59861±0.00159 (1 σ; n = 506), respectively. Total procedural blanks during this study were 15.5 ± 3.5 pg and 0.08 ± 0.03 pg (1 σ; n = 3) for Re and Os, respectively, with an average ¹⁸⁷Os/¹⁸⁸Os value of 0.21 ± 0.01 (n = 3).
Present-day measured ¹⁸⁷Os/¹⁸⁸Os values of samples were corrected to initial osmium ratios (Osᵢ) by
accounting for post-depositional beta decay of <sup>187</sup>Re (λ = 1.666×10<sup>-11</sup>yr<sup>-1</sup> <sup>66</sup>) using an age of 94.0 Ma for
the Cenomanian-Turonian stage boundary<sup>17</sup>.

528

529 **XRF core scanning methods**

530 Split section halves of cores from Site U1516 were analyzed on two Avaatech X-ray fluorescence
531 (XRF) scanners in the XRF Core Scanning Facility at the IODP Gulf Coast Repository in College
532 Station, Texas in 2018. Section halves spanning the OAE2 stratigraphic interval were scanned from
533 Hole C and Hole D (Cores 2R-5R). Emission of secondary fluoresced X-rays characteristic to a suite of
534 elements were generated from scanning at excitation energy levels of 10, 30, and 50 kV. These scans
535 measured chemostratigraphic records of semi-quantitative elemental concentrations in units of counts
536 per second at a median depth sampling resolution of 1.5 cm. Counts of Ca, Fe, and Si are reported from
537 the 10 kV scan (no filter, 0.160 mA). Counts of Ba are reported from the 50 kV scan (Cu filter, 0.75
538 mA). Full details on XRF scanning methods are presented in Bogus et al. (2019)<sup>67</sup>. Processed scanning
539 data can be accessed at the IODP online repository (https://web.iodp.tamu.edu/LORE/). We interpret
540 Ca XRF scanning data to primarily reflect the relative carbonate mineral (calcite) contents of cores
541 given its correlation with shipboard wt. percent carbonate content data (Figure 2).

542

543 **Major and minor element analytical methods**

544 Prior to analysis, bulk-rock samples were powdered to ~10 µm using a ring and puck mill at Sorbonne
545 University's ISTEP laboratory (Paris, France). Concentration of Si and Ba were determined respectively
with an ICP-OES (Thermo Fisher iCap 6500) and an ICP-MS (7700X Agilent) at the spectrochemical laboratory of the Service d'Analyse des Roches et des Minéraux (SARM) of the Centre National de la Recherche Scientifique (CNRS; Vandoeuvre-les-Nancy, France). Sample powder preparation consisted of fusion with LiBO₂, followed by acid digestion in HNO₃ (2%) ⁶⁸. Precision and accuracy are both better than 1% for Si and 5% for Ba, respectively.

Sedimentary Si and Ba contents are considered paleoproductivity proxies in many marine settings³⁷,⁶⁹. To account for the influence of carbonate non-deposition on elemental abundances during OAE2, enrichment factors (EF) were calculated for the elements Si and Ba at Site U1516. For a given element of interest, enrichment factors calculate the ratio of a sample's elemental content to its Al content and normalize that ratio to a reference value, using the following equation⁶⁹:

$$EF = \frac{\left(\frac{\text{element}}{\text{Al}}\right)_{\text{sample}}}{\left(\frac{\text{element}}{\text{Al}}\right)_{\text{average shale}}} \text{Various reference datasets exist. For this work, the EFs were calculated using the Upper Continental Crust (UCC) reference shale}^{70}.$$

## Global OAE2 carbonate content compilation

Values for sedimentary weight percent carbonate contents in the pre-OAE2 and OAE2 intervals reported in the global compilation are estimated averages from studies with quantitative data ²²,²⁶,⁷¹⁻¹⁰⁷.

**Code Availability Statement:** Matlab and R codes for inverse box modeling of isotopic records are available as a Zenodo data repository item (10.5281/zenodo.7182186).
564 **Data Availability Statement:** All geochemical data measured for this study are available as a Zenodo
565 data repository item (10.5281/zenodo.7182186). Core scanning X-ray fluorescence (XRF) data for Site
566 U1516 are available through the International Ocean Discovery Program (IODP) at
567 web.iodp.tamu.edu/LORE/.

568 Correspondence and request for materials should be directed to M.M.J.

569

570 **<u>Methods-only references</u>**

571 56. Schulz, M. & Mudelsee, M. REDFIT: estimating red-noise spectra directly from unevenly spaced
572 paleoclimatic time series. *Comput. Geosci.* **28**, 421–426 (2002).

573 57. Paillard, D., Labeyrie, L. & Yiou, P. Macintosh Program performs time-series analysis. *Eos Trans.*
574 Am. Geophys. Union **77**, 379–379 (1996).

575 58. Tejada, M. L. G. et al. Ontong Java Plateau eruption as a trigger for the early Aptian oceanic anoxic
576 event. *Geology* **37**, 855–858 (2009).

577 59. Peucker-Ehrenbrink, B. & Ravizza, G. The marine osmium isotope record. *Terra Nova* **12**, 205–
578 219 (2000).

579 60. Du Vivier, A. D. C., Selby, D., Condon, D. J., Takashima, R. & Nishi, H. Pacific 187 Os/ 188 Os
580 isotope chemistry and U–Pb geochronology: Synchroneity of global Os isotope change across OAE
581 2. *Earth Planet. Sci. Lett.* **428**, 204–216 (2015).

582 61. Schröder-Adams, C. J., Herrle, J. O., Selby, D., Quesnel, A. & Froude, G. Influence of the High
583 Arctic Igneous Province on the Cenomanian/Turonian boundary interval, Sverdrup Basin, High
584 Canadian Arctic. *Earth Planet. Sci. Lett.* **511**, 76–88 (2019).
62. Sullivan, D. L. et al. High resolution osmium data record three distinct pulses of magmatic activity during cretaceous Oceanic Anoxic Event 2 (OAE-2). *Geochim. Cosmochim. Acta* **285**, 257–273 (2020).

63. Meyers, S. R. Astrochron: An R Package for Astrochronology. http://cran.r-project.org/package=astrochron (2014).

64. Levasseur, S., Birck, J. L. & Allegre, C. J. The osmium riverine flux and the oceanic mass balance of osmium. *Earth Planet. Sci. Lett.* **174**, 7–23 (1999).

65. Selby, D. & Creaser, R. A. Re-Os geochronology of organic rich sediments: an evaluation of organic matter analysis methods. *Chem. Geol.* **200**, 225–240 (2003).

66. Smoliar, M. I., Walker, R. J. & Morgan, J. W. Re-Os ages of group IIA, IIIA, IVA, and IVB iron meteorites. *Science* **271**, 1099–1102 (1996).

67. Bogus, K. A. et al. Shore-based X-ray fluorescence core scanning of IODP Expedition 369 (Australia Cretaceous Climate and Tectonics) material. in *Australia Cretaceous Climate and Tectonics* vol. 369 (2019).

68. Carignan, J., Hild, P., Mevelle, G., Morel, J. & Yeghicheyan, D. Routine Analyses of Trace Elements in Geological Samples using Flow Injection and Low Pressure On-Line Liquid Chromatography Coupled to ICP-MS: A Study of Geochemical Reference Materials BR, DR-N, UB-N, AN-G and GH. *Geostand. Geoanalytical Res.* **25**, 187–198 (2001).

69. Brumsack, H.-J. The trace metal content of recent organic carbon-rich sediments: Implications for Cretaceous black shale formation. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **232**, 344–361 (2006).

70. McLennan, S. M. Relationships between the trace element composition of sedimentary rocks and upper continental crust. *Geochem. Geophys. Geosystems* **2**, (2001).
607 71. Tucholke, B. E. et al. Site 387; Cretaceous to Recent sedimentary evolution of the western
608 Bermuda Rise. in *Initial Reports of the Deep Sea Drilling Project* vol. 43 (Texas A & M
609 University, Ocean Drilling Program, College Station, TX, United States, 1979).

610 72. Arthur, M. A. North Atlantic Cretaceous Black Shales: The Record at Site 398 and a Brief
611 Comparison with Other Occurrences. in *Initial Reports of the Deep Sea Drilling Project* (eds.
612 Sibuet, J.-C. & Ryan, W. B. F.) vol. 47 719–751 (U.S. Government Printing Office, 1979).

613 73. Meyers, P. A. Appendix II. Organic carbon and calcium carbonate analyses, Deep Sea Drilling
614 Project Leg 93, North American continental rise. in *Initial Reports, DSDP, 93* (eds. van Hinte, J. E.
615 & Wise, S. W.) vol. 93 465–469 (U.S. Government Printing Office, 1987).

616 74. Dean, W. E. & Arthur, M. A. 46. Inorganic and organic geochemistry of Eocene to Cretaceous
617 strata recovered from the lower continental rise, North American Basin, Site 603, Deep Sea Drilling
618 Project Leg 93. in *Initial Reports of the Deep Sea Drilling Project* vol. 93 1093–1137 (1987).

619 75. Herbin, J. P., Masure, E. & Roucache, J. Cretaceous Formations from the Lower Continental Rise
620 off Cape Hatteras: Organic Geochemistry, Dinoflagellate Cysts, and the Cenomanian/Turonian
621 Boundary Event at Sites 603 (Leg 93) and 105 (Leg 11). in *Initial Reports of the Deep Sea Drilling
622 Project* (eds. van Hinte, J. E. & Wise, S. W.) vol. 93 1139–1162 (U.S. Government Printing Office,
623 1987).

624 76. Cameron, D. H. Grain-Size and Carbon/Carbonate Analyses, Leg43. in *Initial Reports of the Deep
625 Sea Drilling Project* vol. 43 1043–1047 (U.S. Government Printing Office, 1979).

626 77. van Helmond, N. A. G. M., Ruvalcaba Baroni, I., Sluijs, A., Sinninghe Damsté, J. S. & Slomp, C.
627 P. Spatial extent and degree of oxygen depletion in the deep proto-North Atlantic basin during
628 Oceanic Anoxic Event 2. *Geochem. Geophys. Geosystems* **15**, 4254–4266 (2014).
629 78. Forster, A. et al. The Cenomanian/Turonian oceanic anoxic event in the South Atlantic: New 630 insights from a geochemical study of DSDP Site 530A. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **267**, 256–283 (2008). 631

632 79. Huber, B. T. Middle–Late Cretaceous climate of the southern high latitudes: Stable isotopic 633 evidence for minimal equator-to-pole thermal gradients. *Geol. Soc. Am. Bull.* 28 (1995).

634 80. Uchman, A., Bąk, K. & Rodríguez-Tovar, F. J. Ichnological record of deep-sea 635 palaeoenvironmental changes around the Oceanic Anoxic Event 2 (Cenomanian–Turonian 636 boundary): An example from the Barnasiówka section, Polish Outer Carpathians. *Palaeogeogr.* 637 *Palaeoclimatol. Palaeoecol.* **262**, 61–71 (2008).

638 81. Huber, B. T., Hobbs, R. W., Bogus, K. A. & others, the Expedition 369 Scientists. Site U1513. in 639 *Proceedings of the International Ocean Discovery Program* vol. 369 (2019).

640 82. Thurow, J., Brumsack, H. J., Rullkotter, J., Littke, R. & Meyers, P. A. The Cenomanian-Turonian 641 Boundary event in the Indian Ocean - a key to understand the global picture. in *Synthesis of Results* 642 from *Scientific Drilling in the Indian Ocean* (eds. Duncan, R. A., Rea, D. K., Kidd, R. B., von Rad, 643 U. & Weissel, J. K.) vol. 70 253–274 (American Geophysical Union, 1992).

644 83. Rullkötter, J. et al. Petrography and geochemistry of organic matter in Triassic and Cretaceous 645 deep-sea sediments from the Wombat and Exmouth plateaus and nearby abyssal plains off 646 northwest Australia. in *Proceedings of the Ocean Drilling Program, Scientific Results* (eds. Von 647 Rad, U. & Haq, B. U.) vol. 122 317–333 (Ocean Drilling Program, 1992).

648 84. Thurow, J. et al. The Cenomanian/Turonian Boundary Event (CTBE) at Hole 641A, ODP Leg 103 649 (Compared with the CTBE interval at Site 398). in *Proceeding of the Ocean Drilling Program, 650 Scientific Results* (eds. Boillot, G. & Winterer, E. L.) vol. 103 587–634 (Ocean Drilling Program, 651 1988).
652 85. Shipboard Scientific Party. Leg 207 summary. in *Proc. ODP, Init. Repts.* (eds. Erbacher, J., 653 Mosher, D. C. & Malone, M. J.) 1–89 (2004).

654 86. Sinninghe Damsté, J. S., van Bentum, E. C., Reichart, G.-J., Pross, J. & Schouten, S. A CO2 655 decrease-driven cooling and increased latitudinal temperature gradient during the mid-Cretaceous 656 Oceanic Anoxic Event 2. *Earth Planet. Sci. Lett.* **293**, 97–103 (2010).

657 87. Arnaboldi, M. & Meyers, P. A. Data report: multiproxy geochemical characterization of OAE-658 related black shales at Site 1276, Newfoundland Basin. in *Proc. ODP, Sci. Results* (eds. Tucholke, 659 B. E., Sibuet, J.-C. & Klaus, A.) vol. 210 1–18 (2006).

660 88. Hetzel, A., Boettcher, M. E., Wortmann, U. G. & Brumsack, H.-J. Paleo-redox conditions during 661 OAE 2 reflected in Demerara Rise sediment geochemistry (ODP Leg 207). *Palaeogeogr. 662 Palaeoclimatol. Palaeoecol.* **273**, 302–328 (2009).

663 89. Waples, D. W. & Cunningham, R. Leg 80 Shipboard Organic Geochemistry. in *Initial Reports* 664 *Deep Sea Drilling Program* (eds. de Graciansky, P. C. & Poag, C. W.) vol. 80 949–968 (U.S. 665 Government Printing Office, 1985).

666 90. Wagreich, M., Bojar, A.-V., Sachsenhofer, R. F., Neuhuber, S. & Egger, H. Calcareous 667 nannoplankton, planktonic foraminiferal, and carbonate carbon isotope stratigraphy of the 668 Cenomanian–Turonian boundary section in the Ultrahelvetic Zone (Eastern Alps, Upper Austria). 669 *Cretac. Res.* **29**, 965–975 (2008).

670 91. Turgeon, S. & Brumsack, H.-J. Anoxic vs dysoxic events reflected in sediment geochemistry 671 during the Cenomanian–Turonian Boundary Event (Cretaceous) in the Umbria–Marche Basin of 672 central Italy. *Chem. Geol.* **234**, 321–339 (2006).
92. Linnert, C., Mutterlose, J. & Herrle, J. O. Late Cretaceous (Cenomanian–Maastrichtian) calcareous nannofossils from Goban Spur (DSDP Sites 549, 551): Implications for the palaeoceanography of the proto North Atlantic. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **299**, 507–528 (2011).

93. Exon, N. F., Borella, P. E. & Ito, M. Sedimentology of marine Cretaceous sequences in the central Exmouth Plateau (northwest Australia). in *Proceedings of the Ocean Drilling Program* (eds. von Rad, U. & Haq, B. U.) vol. 122 233–257 (U.S. Government Printing Office, 1992).

94. Meyers, P. A., Yum, J.-G. & Wise, S. W. Origins and maturity of organic matter in mid-Cretaceous black shales from ODP Site 1138 on the Kerguelen Plateau. *Mar. Pet. Geol.* **26**, 909–915 (2009).

95. Dickson, A. J. et al. A Southern Hemisphere record of global trace-metal drawdown and orbital modulation of organic-matter burial across the Cenomanian–Turonian boundary (Ocean Drilling Program Site 1138, Kerguelen Plateau). *Sedimentology* **64**, 186–203 (2017).

96. Huber, B. T., Leckie, R. M., Norris, R. D., Bralower, T. J. & CoBabe, E. Foraminiferal assemblage and stable isotopic change across the Cenomanian–Turonian boundary in the subtropical North Atlantic. *J. Foraminifer. Res.* **29**, 392–417 (1999).

97. Bowman, A. R. & Bralower, T. J. Paleooceanographic significance of high-resolution carbon isotope records across the Cenomanian–Turonian boundary in the Western Interior and New Jersey coastal plain, USA. *Mar. Geol.* **217**, 305–321 (2005).

98. Joo, Y. J. & Sageman, B. B. Cenomanian To Campanian Carbon Isotope Chemostratigraphy from the Western Interior Basin, U.S.A. *J. Sediment. Res.* **84**, 529–542 (2014).

99. Bomou, B. et al. The expression of the Cenomanian–Turonian oceanic anoxic event in Tibet. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **369**, 466–481 (2013).
694 100. Sageman, B. B., Lyons, T. W. & Joo, Y. Ji. Geochemistry of Fine-Grained, Organic Carbon-Rich Facies. in *Treatise on Geochemistry* 141–179 (Elsevier, 2014). doi:10.1016/B978-0-08-095975-7.00706-3.

697 101. Wendler, J. E., Lehmann, J. & Kuss, J. Orbital time scale, intra-platform basin correlation, carbon isotope stratigraphy and sea-level history of the Cenomanian–Turonian Eastern Levant platform, Jordan. *Geol. Soc. Lond. Spec. Publ.* **341**, 171–186 (2010).

700 102. Paez-Reyes, M. et al. Assessing the contribution of the La Luna Sea to the global sink of organic carbon during the Cenomanian-Turonian Oceanic Anoxic Event 2 (OAE2). *Glob. Planet. Change* **199**, 103424 (2021).

703 103. Nederbragt, A. J. & Fiorentino, A. Stratigraphy and palaeoceanography of the Cenomanian-Turonian Boundary Event in Oued Mellegue, north-western Tunisia. *Cretac. Res.* **20**, 47–62 (1999).

706 104. Voigt, S., Gale, A. S. & Voigt, T. Sea-level change, carbon cycling and palaeoclimate during the Late Cenomanian of northwest Europe; an integrated palaeoenvironmental analysis. *Cretac. Res.* **27**, 836–858 (2006).

709 105. Beil, S. et al. New insights into Cenomanian paleoceanography and climate evolution from the Tarfaya Basin, southern Morocco. *Cretac. Res.* **84**, 451–473 (2018).

711 106. Takashima, R. et al. Litho-, bio- and chemostratigraphy across the Cenomanian/Turonian boundary (OAE 2) in the Vocontian Basin of southeastern France. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **273**, 61–74 (2009).

714 107. Peryt, D. & Wyrwicka, K. The Cenomanian/Turonian boundary event in Central Poland. *Palaeogeogr. Palaeoclimatol. Palaeoecol.* **104**, 185–197 (1993).
716 108. Huber, B. T., Hobbs, R. W., Bogus, K. A. & the Expedition 369 Science Party. Expedition 369

717 summary. in *Australia Cretaceous Climate and Tectonics* (eds. Hobbs, R. W., Huber, B. T., Bogus,

718 K. A. & the Expedition 369 Scientists) (2019).

719 109. Bralower, T. J. Calcareous nannofossil biostratigraphy and assemblages of the Cenomanian-

720 Turonian boundary interval: Implications for the origin and timing of oceanic anoxia.

721 *Paleoceanography* **3**, 275–316 (1988).

722 **Extended Data Figure 1. Map of International Ocean Discovery Program (IODP) Site U1516**

723 **(yellow circle)**. Additional sites cored during IODP Expedition 369 in the Mentelle Basin offshore

724 southwest Australia are displayed (modified from <sup>108</sup>). Bathymetric contour interval of 500 meters.

725 DSDP = Deep Sea Drilling Program.

726 **Extended Data Figure 2. Stratigraphic correlation of the Cenomanian-Turonian boundary**

727 **interval at IODP Site U1516 to the Portland Core near the base Turonian Global Stratotype**

728 **Section and Point (GSSP) in Colorado, USA.** Oceanic Anoxic Event 2 (OAE2) is defined by a
729 positive carbon isotope (δ¹³C) excursion (grey shading). The correlation is based on: 1) the base and
730 termination of the δ¹³C excursion²⁷, 2) the anomalously high ¹⁹²Os abundance interval (light purple
731 dashed line) linked to large igneous province (LIP) volcanic activity¹⁹, 3) the base of the initial osmium
732 isotope (¹⁸⁷/¹⁸⁸Os) excursion associated with LIP volcanism at the base of OAE2¹⁹, and 4) bandpassed
733 ~100 kyr-short eccentricity cycles. Calcareous nannofossil biostratigraphy independently supports the

734 correlation as the CC10a/CC10b subzone boundary, approximated by the last occurrence (LO) of

735 *Axopodorhabdus albianus*, falls in the carbonate barren interval at Site U1516²⁷ and lower OAE2
736 interval in Portland Core¹⁰⁹. Site U1516 data sources: δ¹³C data²⁷, Os data and cyclostratigraphy
737 (this study). Portland Core data sources: Os data²⁰, δ¹³C data¹⁶ (figure adapted from¹⁷). kcps =
738 kilocounts per second; VPDB = Vienna Peedee Belemnite scale; ppt = parts per trillion; CCD = calcite
739 compensation depth.

740 **Extended Data Figure 3. REDFIT power spectra of Fe XRF scanning data from IODP Site U1516.** Plots span intervals of A: 455.2–467.4 m rCCSF; B: 467.4–470.7 m rCCSF and C: 470.7–481.8 m rCCSF, using a Welch window with 80%, 90%, 95% and 99% confidence levels, and main periodicities indicated.
741
742
743

744 **Extended Data Figure 4. Floating astrochronology for the Cenomanian-Turonian boundary interval at IODP Site U1516.** This is based on a short eccentricity bandpass ~100-kyr filter of the cyclic Fe XRF scanning data from Hole D. Squares in right pane show sedimentation rate estimates from the bandpass filter and dashed line shows an average sedimentation rate of 0.8 cm/kyr for the general interval from shipboard nannofossil biostratigraphy²⁶. The unradiogenic Osᵢ excursion marks the onset of LIP volcanism several tens of kiloyear before the initiation of the OAE2 carbon isotope excursion in most conformable sites¹⁷,¹⁹,²⁰,⁶⁰,⁶¹. The base of the Osᵢ excursion has been dated to 94.55 ± 0.1 Ma¹⁷ and ~94.9 Ma⁶², providing scenarios for absolute numerical age tie points for the Site U1516 record.
745
746
747
748
749
750
751
752

753 **Extended Data Figure 5. Additional plot of box modeling results for marine $^{187/188}$Os.** This includes the scenario of large igneous province (LIP) volcanism (top panel, A) and a second scenario (bottom panel, B) where fluxes of radiogenic osmium from continental weathering ($F_{riv}$) are set to increase by 80% for 500 kiloyears through Ocean Anoxic Event 2 (OAE2)²⁹. An increase in $F_{riv}$ necessitates an even larger increase in the flux of unradiogenic osmium from large igneous province sources ($F_{LIP}$) by as much as 50x larger than background mantle/volcanic fluxes. Model outputs are plotted against initial osmium isotope ratio data (Osᵢ) through OAE2 from IODP Site U1516 in the
760 Indian Ocean (circle symbols and dotted line, this study) and the Angus Core in the Western Interior
761 Basin of North America¹⁷.

762 **Extended Data Table 1.** Model parameters for Late Cretaceous marine Os fluxes and model
763 initial conditions. See **Extended Data Table 2** for final modeled F<sub>LIP</sub> values through time.

764 **Extended Data Table 2.** Modeled large igneous province Os flux (F<sub>LIP</sub>) values through time used
765 to perturb the marine Os reservoir during OAE2. F<sub>LIP</sub> values are reported as a ratio to the
766 background global volcanic/mantle flux (F<sub>mantle</sub>) (i.e., F<sub>LIP</sub>/F<sub>mantle</sub>). For example, a F<sub>LIP</sub> value of 1.0
767 represents a large igneous province Os flux as large as the pre-event background volcanic/mantle Os
768 flux.

769 **Extended Data Table 3.** A global compilation of carbonate contents (wt.% CaCO3) pre-OAE2
770 and during OAE2 from 46 sites. Sites in compilation are mapped in **Figure 1a**.
A

~94 Ma ocean

<figure><img src="image_3.png" /></figure>

**Sites**

● epicontinental sea

● LIP or continental slope

● continental rise/abyssal

● shoaling of CCD during OAE2

B

global marine carbonate content trends through OAE2

<figure><img src="image_4.png" /></figure>
<figure><img src="image_5.png" /><figcaption>Proxies for local primary productivity</figcaption></figure>
A

<figure><img src="image_6.png" /></figure>

LIP volcanic activity

(LIP Os flux / background mantle Os flux)

B

$^{192}$Os (ppt)

<figure><img src="image_7.png" /></figure>

$^{187/188}$Os$_{ocean}$

OAE2
interval

<figure><img src="image_8.png" /></figure>

Age (Ma)
32°

34°

36°

<figure><img src="image_9.png" /></figure>

Perth

Western Australia

U1514

Mentelle Basin

U1513

O

U1516

Naturaliste Plateau

DSDP Site 258
# IODP Site U1516

## Mentelle Basin, southeast Indian Ocean

<figure><img src="image_10.png" /></figure>

### Portland Core, Colorado, USA
(base Turonian GSSP)

<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>
<figure><img src="image_13.png" /><figcaption>B</figcaption></figure>

<figure><img src="image_14.png" /><figcaption>C</figcaption></figure>

<figure><img src="image_15.png" /></figure>
Depth m (rCCSF)

Core Photos

<figure><img src="image_16.png" /></figure>

XRF Fe counts
(kilocounts/sec)

<figure><img src="image_17.png" /></figure>

%Carbonate

<figure><img src="image_18.png" /></figure>

CCD shoaled

Osᵢ

<figure><img src="image_19.png" /></figure>

LIP Volcanism

lower sed. rates
in OAE2 interval

<figure><img src="image_20.png" /></figure>

XRF Fe counts
(kilocounts/sec)

XRF Ca counts
(kilocounts/sec)
A

SCENARIO 1: LIP VOLCANISM

Time (kiloyears)

<figure><img src="image_21.png" /></figure>

Osᵢ

<figure><img src="image_22.png" /></figure>

LIP Os flux / baseline mantle

<figure><img src="image_23.png" /></figure>

Riverine Os flux / baseline

B

SCENARIO 2: LIP VOLCANISM + INCREASED CONTINENTAL WEATHERING

Time (kiloyears)

<figure><img src="image_24.png" /></figure>

Osᵢ

<figure><img src="image_25.png" /></figure>

LIP Os flux / baseline mantle

<figure><img src="image_26.png" /></figure>

Riverine Os flux / baseline
<table><thead><tr><th>Marine Os Source/Sink</th><th><sup>187/188</sup>Os<br>(isotopic composition)</th><th>Flux<br>(tonnes Os/kyr)</th></tr></thead><tbody><tr><td>Riverine (F<sub>riv</sub>)</td><td>1.4</td><td>339</td></tr><tr><td>Mantle/volcanic (F<sub>mantle</sub>)</td><td>0.126</td><td>352</td></tr><tr><td>Cosmic (F<sub>cosmic</sub>)</td><td>0.126</td><td>17.6</td></tr><tr><td>Large igneous province (F<sub>LIP</sub>)</td><td>0.126</td><td>iteratively varied*</td></tr><tr><td>Sedimentation (F<sub>sed</sub>)</td><td>Ro(t)</td><td>M<sub>o</sub>(t)/τ</td></tr><tr><td>Model Initial Conditions (t=0)</td><td>Symbol</td><td>Value</td></tr><tr><td>marine <sup>187/188</sup>Os ratio</td><td>R<sub>o</sub>(0)</td><td>0.9</td></tr><tr><td>mass of marine Os reservoir</td><td>M<sub>o</sub>(0)</td><td>13,000 tonnes</td></tr><tr><td>marine Os residence time</td><td>τ</td><td>20,000 years</td></tr></tbody></table>
<table><thead><tr><td>Floating time (kiloyears)</td><td>LIP Os flux (F<sub>LIP</sub>) (ratio: F<sub>LIP</sub>/F<sub>mantle</sub>)</td></tr></thead><tbody><tr><td>1500</td><td>0.75</td></tr><tr><td>800</td><td>0.75</td></tr><tr><td>550</td><td>1</td></tr><tr><td>280</td><td>2</td></tr><tr><td>170</td><td>3.75</td></tr><tr><td>50</td><td>35.75</td></tr><tr><td>-10</td><td>35.75</td></tr><tr><td>-20</td><td>2.5</td></tr><tr><td>-200</td><td>0.75</td></tr><tr><td>-440</td><td>0</td></tr><tr><td>-1500</td><td>0</td></tr></tbody></table>
<table><thead><tr><th>Site</th><th>Basin</th><th>Setting</th><th>pre-OAE2<br>%CaCO3</th><th>OAE2<br>%CaCO3</th><th>% change</th><th>CCD<br>classification</th><th>Resolution of<br>%carb record</th><th>References</th></tr></thead><tbody><tr><td>W Bermuda Rise<br>(DSDP Site 387)</td><td>proto-North Atlantic</td><td>Abyssal Plain</td><td>0</td><td>0</td><td>0%</td><td>sub-CCD</td><td>coarse</td><td>Tucholke et al. (1979)</td></tr><tr><td>Vigo Seamount<br>(DSDP Site 398)</td><td>proto-North Atlantic</td><td>Seamount</td><td>0</td><td>0</td><td>0%</td><td>sub-CCD</td><td>coarse</td><td>Arthur (1979)</td></tr><tr><td>eastern N American Rise<br>(DSDP Site 603)</td><td>proto-North Atlantic</td><td>Abyssal Plain</td><td>0</td><td>0</td><td>0%</td><td>sub-CCD</td><td>coarse</td><td>Meyers (1987)<br>Dean and Arthur (1987)</td></tr><tr><td>New Jersey Margin<br>(DSDP Site 105)</td><td>proto-North Atlantic</td><td>Continental Rise</td><td>7</td><td>7</td><td>0%</td><td>sub-CCD</td><td>high</td><td>Herbin et al. (1987)</td></tr><tr><td>Central Bermuda Rise<br>(DSDP Site 386)</td><td>proto-North Atlantic</td><td>Continental Rise</td><td>15</td><td>9</td><td>-40%</td><td>sub-CCD</td><td>coarse</td><td>Cameron (1979)<br>van Helmond et al. (2014)</td></tr><tr><td>Walvis Ridge<br>(DSDP Site 530)</td><td>proto-South Atlantic</td><td>Abyssal Plain</td><td>15</td><td>20</td><td>33%</td><td>sub-CCD</td><td>high</td><td>Forster et al. (2008)</td></tr><tr><td>Falkland Plateau<br>(DSDP Site 511)</td><td>proto-South Atlantic</td><td>Continental Slope</td><td>~0</td><td>~0</td><td>0%</td><td>sub-CCD</td><td>coarse</td><td>Huber et al. (1995)</td></tr><tr><td>Poland - Barnasiowka<br>Section</td><td>Tethys?</td><td>Continental slope?</td><td>~0</td><td>~0</td><td>0%</td><td>sub-CCD</td><td>n/a</td><td>Uchman et al. (2008)</td></tr><tr><td>Mentelle Basin<br>(IODP Site U1516)</td><td>proto-Indian Ocean</td><td>Continental Slope</td><td>25</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>Huber et al. (2019) Site U1516</td></tr><tr><td>Mentelle Basin<br>(IODP Site U1513)</td><td>proto-Indian Ocean</td><td>Continental Slope</td><td>25</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>Huber et al. (2019) Site U1513</td></tr><tr><td>Exmouth Plateau<br>(ODP Site 763C)</td><td>Tethys</td><td>Continental Slope</td><td>38</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>Thurow et al. (1992)<br>Rullkötter et al. (1992)</td></tr><tr><td>Galicia Margin<br>(ODP Site 641)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>40</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>Thurow et al. (1988)<br>van Helmond et al. (2014)</td></tr><tr><td>Newfoundland Margin<br>(ODP Site 1276)</td><td>proto-North Atlantic</td><td>Continental Rise</td><td>42</td><td>9</td><td>-79%</td><td>shoaled CCD</td><td>high</td><td>Shipboard Sci. Party (2004)<br>Sinninghé Damste et al (2010)<br>Arnaboldi and Meyers (2006)</td></tr><tr><td>Demerara Rise<br>(ODP Site 1258)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>52</td><td>20</td><td>-62%</td><td>shoaled CCD?</td><td>high</td><td>Hetzel et al. (2009)</td></tr><tr><td>Austria - Rehkogelgraben<br>Section</td><td>Tethys</td><td>Continental Slope</td><td>60</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>Wagreich et al. (2008)</td></tr><tr><td>Italy - Furlo section</td><td>Tethys</td><td>Continental Slope?</td><td>67</td><td>3</td><td>-97%</td><td>shoaled CCD</td><td>high</td><td>Turgeon & Burmsack (2006)</td></tr><tr><td>Goban Spur<br>(DSDP Site 549)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>70</td><td>6</td><td>-91%</td><td>shoaled CCD</td><td>high</td><td>de Graciansky et al. (1985)<br>Waples & Cunningham (1985)<br>Linnert et al. (2011)</td></tr><tr><td>Goban Spur<br>(DSDP Site 551)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>75</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>high</td><td>de Graciansky et al. (1985)<br>Waples & Cunningham (1985)<br>Linnert et al. (2011)</td></tr><tr><td>Exmouth Plateau<br>(ODP Site 762)</td><td>Tethys</td><td>Continental Slope</td><td>76</td><td>0</td><td>-100%</td><td>shoaled CCD</td><td>coarse</td><td>Exon et al. (1992)</td></tr><tr><td>Kerguelen Plateau<br>(ODP Site 1138)</td><td>proto-Indian Ocean</td><td>LIP Slope</td><td>77</td><td>4</td><td>-95%</td><td>shoaled CCD</td><td>high</td><td>Meyers et al. (2009)<br>Dickson et al. (2017)</td></tr><tr><td>Blake Nose<br>(ODP Site 1050)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>85</td><td>14</td><td>-84%</td><td>shoaled CCD</td><td>coarse</td><td>Huber et al. (1999)</td></tr><tr><td>Bass River, New Jersey<br>(ODP Leg 174x/)</td><td>proto-North Atlantic</td><td>Epicontinental</td><td>10</td><td>25</td><td>150%</td><td>above CCD?<br>Siliciclastic-<br>rich, proximal</td><td>high</td><td>Bowman and Bralower (2005)</td></tr><tr><td>Colorado, USA<br>Angus Core, Denver Basin</td><td>Western Interior<br>Seaway</td><td>Epicontinental</td><td>45</td><td>65</td><td>44%</td><td>above CCD</td><td>high</td><td>Joo & Sageman (2014)</td></tr><tr><td>northwest Europe<br>(Grobern borehole)</td><td>European Shelf Seas</td><td>Epicontinental</td><td>45</td><td>55</td><td>22%</td><td>above CCD</td><td>high</td><td>Voigt et al. (2006)</td></tr><tr><td>Gongzha, Tingri Tibet</td><td>Tethys</td><td>Epicontinental</td><td>50</td><td>70</td><td>40%</td><td>above CCD</td><td>high</td><td>Bomou et al. (2013)</td></tr><tr><td>Colorado, USA<br>(Rock Canyon Anticline)</td><td>Western Interior<br>Seaway</td><td>Epicontinental</td><td>50</td><td>65</td><td>30%</td><td>above CCD</td><td>high</td><td>Sageman et al. (2014)</td></tr><tr><td>Jordan<br>(GM3 Section)</td><td>Tethys</td><td>Epicontinental</td><td>60</td><td>75</td><td>25%</td><td>above CCD</td><td>high</td><td>Wendler et al. (2010)</td></tr><tr><td>Columbia<br>Olini & Paipa sections</td><td>La Luna Sea</td><td>Epicontinental</td><td>60</td><td>50</td><td>-17%</td><td>above CCD</td><td>high</td><td>Paez-Reyes et al. (2021)</td></tr><tr><td>Demerara Rise<br>(ODP Site 1260)</td><td>proto-North Atlantic</td><td>Continental Slope</td><td>63</td><td>48</td><td>-24%</td><td>above CCD</td><td>high</td><td>Hetzel et al. (2009)</td></tr><tr><td>Tunisia<br>Oued Mellegue area</td><td>Tethys</td><td>Epicontinental</td><td>70</td><td>80</td><td>14%</td><td>above CCD</td><td>high</td><td>Nederbragt & Fiorentino (1999)</td></tr><tr><td>Morocco<br>(SN4 Core, Tarfaya Basin)</td><td>proto-North Atlantic</td><td>Epicontinental</td><td>70</td><td>75</td><td>7%</td><td>above CCD</td><td>high</td><td>Beil et al. (2018)</td></tr><tr><td>Kansas, USA<br>(Cuba section)</td><td>Western Interior<br>Seaway</td><td>Epicontinental</td><td>75</td><td>80</td><td>7%</td><td>above CCD</td><td>high</td><td>Bowman and Bralower (2005)</td></tr><tr><td>France<br>(Lambruisse Section)</td><td>European Shelf Seas</td><td>Epicontinental</td><td>85</td><td>45</td><td>-47%</td><td>above CCD</td><td>high</td><td>Takashima et al. (2009)</td></tr><tr><td>Poland<br>(Pulawy Borehole)</td><td>European Shelf Seas</td><td>Epicontinental</td><td>85</td><td>60</td><td>-29%</td><td>above CCD</td><td>coarse</td><td>Peryt & Wyrwicka (1993)</td></tr><tr><td>Guerrero state, Mexico<br>carbonate platforms</td><td>Epicontinental Sea</td><td>Epicontinental</td><td>95</td><td>80</td><td>-16%</td><td>above CCD</td><td>high</td><td>Elrick et al. (2009)</td></tr></tbody></table>