Stable Isotope Constraints on Marine Productivity Across the Cretaceous-Paleogene Mass Extinction

Stable Isotope Constraints on Marine Productivity Across the Cretaceous#Paleogene Mass Extinction

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Citation Sepúlveda, Julio et. al., "Stable Isotope Constraints on Marine Productivity Across the Cretaceous#Paleogene Mass Extinction." Paleoceanography and Paleoclimatology 34, 7 (July 2019): 1195-1217 doi. 10.1029/2018PA003442 ©2019 Authors

As Published 10.1029/2018PA003442

Publisher American Geophysical Union (AGU)

Version Final published version

Citable link https://hdl.handle.net/1721.1/125278

Terms of Use Article is made available in accordance with the publisher's policy and may be subject to US copyright law. Please refer to the publisher's site for terms of use. RESEARCH ARTICLE Stable Isotope Constraints on Marine Productivity Across 10.1029/2018PA003442 the Cretaceous‐Paleogene Mass Extinction Special Section: Julio Sepúlveda1 , Laia Alegret2, Ellen Thomas3,4 , Emily Haddad5, Changqun Cao6, Climatic and Biotic Events of 7 the Paleogene: Earth Systems and Roger E. Summons and Planetary Boundaries in a 1 Greenhouse World Organic Geochemistry Laboratory, Department of Geological Sciences and Institute of Arctic and Alpine Research (INSTAAR), University of Colorado Boulder, Boulder, CO, USA, 2Departamento de Ciencias de la Tierra and Instituto Universitario de Ciencias Ambientales de Aragón, Universidad de Zaragoza, Zaragoza, Spain, 3Department of Geology Key Points: and Geophysics, Yale University, New Haven, CT, USA, 4Department of Earth and Environmental Sciences, Wesleyan • Carbonate and organic matter stable University, Middletown, CT, USA, 5Department of Earth Sciences, University of California, Riverside, CA, USA, 6State isotope records were obtained from eight neritic and upper bathyal Key Laboratory of Palaeobiology and Stratigraphy, Nanjing Institute of Geology and Palaeontology, Chinese Academy of sections Sciences, Nanjing, China, 7Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of • Isotopic signatures are spatially and Technology, Cambridge, MA, USA temporally heterogeneous, and the organic and inorganic carbon pools are decoupled The effects of the Cretaceous‐Paleogene (K/Pg) mass extinction (~66 Ma) on marine primary • Carbon cycling and marine Abstract productivity may have recovered and export productivity remain debated. We studied changes in carbon and nitrogen cycling in eight faster than pelagic sites and within neritic and upper bathyal sections with expanded K/Pg boundary clay layers in the western Tethys and the deposition of the boundary clay northeastern Atlantic Ocean, by measuring stable carbon isotopes of bulk carbonate (δ13C ) and organic layer carb 13 15 matter (δ Corg), nitrogen isotopes in bulk organic matter (δ N), and selected compound‐specific carbon 13 13 13 13 Supporting Information: isotopic records (δ Clipid). Negative carbon isotope excursions (CIEs) in δ Ccarb, δ Corg, and δ Clipid are • Supporting Information S1 temporally and spatially heterogeneous as well as decoupled from each other, suggesting that factors affecting the δ13C of dissolved inorganic carbon, as well as isotopic fractionation during carbon fixation 13 13 across the K/Pg, are more complex than commonly assumed. The negative CIEs in δ Corg and δ Clipid at Correspondence to: each site are smaller in amplitude and shorter in duration than those in δ13C , but in most sections both J. Sepúlveda, carb [email protected] carbon pools recovered to preboundary conditions within the time of deposition of the boundary clay layer (<103–104 Kyr) or shortly thereafter. This rapid recovery is supported by limited δ15N data, which mostly suggests moderate or minor changes in redox conditions (except in Denmark), marine productivity, and Citation: Sepúlveda, J., Alegret, L., Thomas, E., phytoplanktonic nitrate utilization in the earliest Danian. Our results indicate that carbon cycling and Haddad, E., Cao, C., & Summons, R. E. primary productivity in neritic and upper bathyal regions recovered to preboundary levels faster (<104Kyr) (2019). Stable isotope constraints on than in oceanic regions (105–106 years), likely sustained by resilient noncalcifying phytoplankton with marine productivity across the Cretaceous‐Paleogene mass extinction. resting stages, consistent with modeling and proxy studies. Paleoceanography and Paleoclimatology, 34. https://doi.org/ Plain Language Summary Sixty‐six million years ago, at the boundary between the Cretaceous 10.1029/2018PA003442 and Paleogene Periods (K/Pg boundary), a meteorite impacted the Earth during a time of active volcanism causing the mass extinction of marine and terrestrial species. Despite decades of research, the consequences Received 17 JUL 2018 of the mass extinction to marine photosynthesis and the cycling of carbon in the ocean remain contentious. Accepted 31 MAY 2019 Accepted article online 19 JUN 2019 We investigated the light and heavy stable isotopes of carbon preserved in rocks of ancient seafloor along continental margins of the western Tethys and northeastern Atlantic Oceans to establish for how long the cycling of carbon may have been disrupted. Our results indicate that the response of marine productivity and carbon cycling following the impact was heterogenous for thousands and tens of thousands of years but that it recovered to pre‐K/Pg boundary levels hundreds of thousands of years earlier than open ocean regions. We suggest that resilient phytoplankton without carbonate skeletons living along continental margins may have recolonized surface waters relatively quickly after this mass extinction event.

1. Introduction The Cretaceous‐Paleogene (K/Pg) boundary (~66 Ma) is the most recent of the mass extinction events that punctuate the Phanerozoic Eon and possibly the most intensively studied. The hypothesis that an asteroid impact on the Yucatan was the proximate cause of the extinction has been widely (e.g., Alvarez et al.,

©2019. American Geophysical Union. 1980; Schulte et al., 2010; Morgan et al., 2016) but not universally accepted (Keller et al., 2004; Renne All Rights Reserved. et al., 2015). However, recent high‐precision geochronologic studies indicate that the most highly abrupt,

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eruptive phases of the Deccan Traps Large Igneous Province were not synchronous with the mass extinction (Burgess, 2019; Schoene et al., 2019; Sprain et al., 2019). In addition, the specific effects on ecosystems remain vigorously debated (e.g., Keller et al., 2004; Kring, 2007; Schulte et al., 2010). Pelagic calcareous microfossils (e.g., foraminifera and nannoplankton) collapsed in abundance and diversity at the K/Pg boundary (Bown, 2005; MacLeod et al., 2007; Olsson et al., 1999; Pospichal, 1994), simultaneously with an abrupt, global negative carbon isotopic excursion (~1–3‰; CIE) 13 5 6 in planktonic foraminifera and bulk marine carbonates (δ Ccarb), which persisted for ~10 –10 years (e.g., Birch et al., 2016; Coxall et al., 2006; D'Hondt, 2005; Hsü et al., 1982; Hsü & Mckenzie, 1985; Zachos & Arthur, 1986). Benthic foraminifera, however, did not experience substantial extinction (e.g., Alegret et al., 2001; Alegret & Thomas, 2005; Culver, 2003; Thomas, 1990), and their tests (carbonate formed in ocean bottom waters) show a more gradual, opposite (positive) CIE, indicating that the ocean's surface bottom δ13C gradient may have disappeared or even reversed (e.g., Alegret et al., 2012; Birch et al., 2016; Coxall et al., 2006; D'Hondt, 2005; D'Hondt et al., 1998; Hsü et al., 1982; Hsü & Mckenzie, 1985; Zachos & Arthur, 1986). This breakdown of the vertical gradient has been interpreted as evidence for a global disruption of the export of organic matter into the deep sea due to an extreme decline in marine biological productivity (“Strangelove Ocean”; Hsü et al., 1982, Hsü & Mckenzie, 1985; Kump, 1991; Zachos & Arthur, 1986). Alternatively, it has been proposed that the fraction of productivity exported to the deep sea was reduced globally, with the reduction in export ranging from a few percent to a “collapse” (Living Ocean; Adams et al., 2004; D'Hondt et al., 1998; D'Hondt, 2005; Coxall et al., 2006; Ridgwell et al., 2010). A severe and prolonged, globally uniform decline in primary and/or export productivity conflicts, however, with microfossil (Alegret et al., 2012; Alegret & Thomas, 2009; Brinkhuis et al., 1998; Brinkhuis & Leereveld, 1988; Brinkhuis & Zachariasse, 1988; Hollis et al., 1995; Jiang et al., 2010; Vellekoop et al., 2015, 2017), macrofossil (Sogot et al., 2013), molecular (Mita & Shimoyama, 1999; Sepúlveda, Wendler, Summons, et al., 2009; Sepúlveda et al., 2012; Shimoyama et al., 2001), and bulk geochemical evidence (Hull & Norris, 2011; Kaiho et al., 1999). There are further complications with the hypothesis of a global, prolonged decline in export productivity 13 13 based on the nature of δ Ccarb records. First, many assume that changes in δ Ccarb dominantly reflect changes in the carbon isotopic composition of the oceanic‐atmospheric reservoir (i.e., dissolved inorganic carbon in the oceans). However, changes in the relative contribution of different carbonate sources with dis- tinctive isotopic signatures across the K/Pg complicate this assumption, because of the mass extinction of the carbonate‐producing organisms. For instance, the extinction of calcareous nannoplankton and planktonic foraminifera at the K/Pg removed the main sources of pelagic biogenic carbonate, leading to common abundance of isotopically depleted diagenetic carbonate, including dolomitic crystals (Minoletti et al., 2005). The extinction of large, symbiont‐bearing planktonic foraminifera caused a decline in the δ13C in planktonic foraminiferal tests (Birch et al., 2012), and the living counterparts of small biserial and triserial planktic foraminifera, which thrived after the extinction, generally have a light stable carbon isotope signal (Kimoto et al., 2009; Smart & Thomas, 2006). The mass extinction of calcareous nannoplankton led to a larger relative contribution of fine‐grained carbonate produced by the isotopically light calcareous dinocyst Thoracosphaera (e.g., Alegret et al., 2012; Gardin & Monechi, 1998; Minoletti et al., 2005). Taken together, 13 this evidence highlights that δ Ccarb signatures can be related to the cycling of carbon on a global scale but influenced by changes in carbonate sources (“vital effects”) in the aftermath of the extinction of pelagic calcifiers at the K/Pg boundary. In addition, a decline in organic carbon export to the seafloor does not in itself result in an inverted (rather than a collapsed) isotopic gradient, as observed in some locations (e.g., Alegret et al., 2012; Hsü & Mckenzie, 1985). Such a reversed gradient has been attributed to the existence of a “respiring ocean,” where respiration (including remineralization of organic matter) may have been more important in surface waters than photosynthesis (Hsü & Mckenzie, 1985), or to the ocean's “solubility pump,” which imprints an inverted isotopic gradient than that generated by the biological pump (Thomas et al., 2009). Alternatively, isotopically light carbon may have been added to surface oceanic and atmospheric reservoirs due to biomass burning of up to ~25% of above ground biomass (Arinobu et al., 1999; Ivany & Salawitch, 1993) or to dissociation of gas hydrates as a result of seismic activity due to the impact (Bralower et al., 1998; Max et al., 1998; Alegret et al., 2003; Day & Maslin, 2005). Such a release of isotopically light carbon, however, should have resulted in a negative CIE throughout the system, including terrestrial biomass and deep‐sea benthic foraminifera. A

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CIE ranging from −1.0 to −2.8‰ has been reported in terrestrial organic matter (e.g., Arens & Jahren, 2000; Maruoka et al., 2007; Schimmelmann & DeNiro, 1984) and plant biomarkers (Arinobu et al., 1999), although the statistical signiﬁcance of the CIE in bulk terrestrial records has been challenged (Grandpre et al., 2013). It is difﬁcult to argue that a large‐scale release of isotopically light carbon persisted for the time‐scale of the CIE in sedimentary records (>105 years), i.e., much longer than the mixing time of the oceans, without a similarly sized CIE in deep‐sea benthic foraminifera. Indeed, benthic stable isotope records show an opposite trend to that of the plankton. Finally, the magnitude and duration of the CIEs in pelagic sections (Alegret et al., 2012; Coxall et al., 2006; D'Hondt, 2005; D'Hondt et al., 1998; Hsü et al., 1982; Zachos & Arthur, 1986), and neritic and upper bathyal sections (e.g., Kaiho et al., 1999; Keller & Lindinger, 1989; Sepúlveda, Wendler, Summons, et al., 2009) are markedly different. This disagreement by itself argues against hypotheses explaining global carbon isotope records by one cause, whether a reduction in export productivity or the emission of isotopically light carbon. For this study, it is not of great importance whether the proximate cause of extinctions was the Chicxulub asteroid impact (on limestone) or the Deccan Trap volcanism (through degassing): potential carbon emitted as the results of these processes probably did not affect the carbon isotope records directly, because of the short time scale and the fact that the difference in carbon isotopic signatures of lime-

stone and volcanogenic CO2 is insufﬁcient (e.g., Kump & Arthur, 1999; Mason et al., 2017). Alternative means to evaluate past changes in carbon cycling and primary productivity include the δ13Cof 13 13 13 total organic carbon (TOC; δ Corg) and the δ C in lipid biomarkers (δ Clipid) derived from known sources (e.g., Freeman & Hayes, 1992; Hayes, 1993; Hayes et al., 1999; Sinninghe Damsté et al., 2008). Across mass

extinction events, the concentrations of marine dissolved CO2 would have changed markedly, affecting iso- 13 13 topic fractionation and, therefore, δ Corg and δ Clipid in paleorecords (e.g., Grice et al., 2005; Hollander et al., 1993; Wang et al., 1994; Ward et al., 2001). Due to its specificity to known biological sources, 13 δ Clipid is the most reliable and can be used to constrain changes in carbon cycling and marine productivity across time intervals of environmental and biological perturbations (e.g., Sinninghe Damsté et al., 2008; van 13 Bentum et al., 2012) and mass extinction (Grice et al., 2005). However, acquiring δ Clipid is challenging due to the low organic content and biomarker yield of most sedimentary rocks across the K/Pg boundary. When 13 13 analysis of δ Clipid is not feasible, δ Corg can serve as a partial alternative, at least if the bulk sources of organic matter are known (e.g., Hayes, 1993; Meyers, 1992, 1997). 15 Finally, the stable isotope composition of nitrogen in bulk organic matter (δ Norg) is a valuable tool to gain information on past changes in marine nitrogen biogeochemistry, including denitrification, nitrogen fixation, and phytoplanktonic nutrient utilization (e.g., Rau et al., 1987; Robinson et al., 2012; Sigman et al., 15 2009). Records of δ Norg have been used, for instance, to estimate changes in nitrogen cycling during periods of widespread oxygen deficiency (e.g., Junium & Arthur, 2007; Kuypers et al., 2004; Luo et al., 2011; Meyers et al., 2009) and across glacial‐interglacial transitions (Altabet et al., 1995; Galbraith et al., 2013; Ganeshram et al., 1995). We present a multiisotopic investigation of carbon and nutrient cycling across expanded K/Pg boundary clay 13 13 layers in eight sections from northern Africa to northern Europe. We use the bulk δ Ccarb, δ Corganic, and 15 13 δ Norganic data, combined with some δ Cbiomarker values, to infer changes in marine productivity after the K/Pg extinction.

2. Materials and Methods 2.1. Study Areas and Sampling Samples were collected across the boundary clay layer in eight outcrop sections in four regions along a latitudinal and bathymetric gradient (Figures 1 and 2) during ﬁeld campaigns in Tunisia (sections El Kef and Aïn Settara; Peryt et al., 2002; Alegret, 2003, 2008), Spain (sections Caravaca and Agost; Alegret et al., 2003; Alegret, 2007), France (sections Bidart and Loya; Alegret et al., 2004; Alegret, 2007), and Denmark (sections Kulstirenden and Højerup; Hart et al., 2004; Sepúlveda, Wendler, Summons, et al., 2009). We provide a summary of each location below, with more detailed descriptions available in Figure 2 and the references cited above. Sample resolution varied, but the same sampling scheme was used for all sections (below, within, and above the clay layer). Within the clay layer, sampling resolution varied between 5 and 20 cm at El Kef and Aïn Settara (with relatively high sedimentation rates), 1 cm at Caravaca and Agost, 2 cm at Bidart

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Figure 1. Paleogeography of the Cretaceous‐Paleogene boundary (66 Ma). (a) Global paleogeographic reconstruction. White start denotes the location of the Chicxulub impact and the white box depicts the area shown in (b). (b) Paleomap showing the study areas in Tunisia (El Kef and Ain Settara), Spain (Caravaca and Agost), France (Bidart and Loya), and Denmark (Kulstirenden and Højerup). Green, light blue, and dark blue depict emerged land, shallow seas, and open oceans, respectively. © 2016 Colorado Plateau Geosystems Inc.

and Loya, and 1–5 cm at Højerup and Kulstirenden. Fresh unweathered rock samples were collected after removal of the upper surface layer. The base of the boundary clay layer coincides with the K/Pg boundary (Molina et al., 2006) and has been interpreted to be synchronous worldwide (Arinobu et al., 1999; Giron, 2013; Giron et al., 2012; Mita & Shimoyama, 1999; Mukhopadhyay et al., 2001). We plotted geochemical and stable isotope data with respect to distance from the K/Pg boundary, with positive and negative numbers indicating a position above or below the base of the impact layer, respectively. In order to facilitate the comparison of data between different localities on the same temporal scale, the vertical axis of each section was normalized to the thickness of the clay layer, where 0 represents the lowermost limit of the clay layer, that is, the K/Pg boundary, and 1 indicates the uppermost limit of the clay layer. We deﬁned the thickness of the clay layer as previously described in the studies cited above based on lithological, micropaleontological, and geochemical features (e.g., Molina et al., 2006).

In most cases, the boundary clay layer is characterized by dark shales or marly shales with low CaCO3%. 2.1.1. Sections in Tunisia The El Kef and Aïn Settara sections have the most expanded K/Pg boundary layers, with ~100‐ and 55‐cm‐ thick clay layers, respectively (Figure 2; e.g., Dupuis et al., 2001; Keller & Lindinger, 1989; Molina et al., 2006;

Figure 2. Biostratigraphy and lithology of the studied sections. The light gray area with connecting lines indicates the Guembelitria cretacea zone across the eight sections.

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Perch–Nielsen et al., 1982). The K/Pg boundary lies in the upper Maastrichtian‐Paleocene El Haria Formation, deposited in a narrow strait separating the Gafsa‐Metlaoui Basin from the Tunisian platform (Dupuis et al., 2001). Benthic foraminiferal assemblages indicate an outer shelf‐uppermost bathyal setting (~200–300 m; Keller, 1988; Speijer & Van der Zwaan, 1996; Peryt et al., 2002; Culver, 2003; Alegret, 2008). The El Kef section (Kalaat Senan Region) is the Global Boundary Stratotype Section and Point for the K/Pg boundary (Molina et al., 2006). The Aïn Settara section is located ~50 km south of the town of El Kef, in the region between Kalaat Senan, Tajerouine, and Kalaa Khasba (Arenillas et al., 2000; Dupuis et al., 2001; Peryt et al., 2002) 2.1.2. Sections in Spain The Caravaca and Agost sections are located in the Betic Cordillera in southeastern Spain (Murcia and Alicante provinces, respectively; Figure 1) and are among the most continuous K/Pg boundary outcrop sections (e.g., Smit, 1977, 1990; Molina et al., 1996; Kaiho et al., 1999; Mukhopadhay et al., 2001). Uppermost Cretaceous and lowermost Paleogene hemipelagic gray marls were deposited at middle bathyal depths (~600–1,000 m; Coccioni et al., 1993; Coccioni & Galeotti, 1994; Alegret et al., 2003). The K/Pg boundary is at the contact between Maastrichtian marls and an ∼10‐cm‐thick layer of black clays, with a 2‐ to 3‐ mm‐thick, ferruginous level at its base (Figure 2). The clay layer contains an iridium anomaly, has a low 13 CaCO3 content, a negative δ C excursion, and trace elements indicative of low oxygen conditions (Arinobu et al., 1999; Kaiho et al., 1999; Kaiho & Lamolda, 1999; Martínez‐Ruiz et al., 1999; Ruiz et al., 1992; Smit, 1990). 2.1.3. Sections in France The Bidart and Loya sections are located in the Basque‐Cantabrian Basin of Southwestern France, near the towns of Bidart and Hendaya, respectively (Figure 1). Upper Maastrichtian gray marls and marly limestones with evidence of bioturbation are overlain by a red ferruginous layer that characterizes the K/Pg boundary (Figure 2; Smit, 2004; Rodríguez‐Tovar et al., 2010; Alegret et al., 2015). This layer is ~2 mm thick at Bidart, 2 cm thick at Loya, and overlain by a dark gray clay layer (Bonté et al., 1984). At Bidart, the clay layer is about 13 cm thick and overlain by marls and marly limestones (Apellaniz et al., 1997; Renard et al., 1982; Rodríguez‐Tovar et al., 2010). At Loya, the clay layer is 18 cm thick and overlain by a 30‐cm‐thick limestone. Both sections were deposited at upper‐middle bathyal depths (~500–700 m) in well‐oxygenated bottom waters (Alegret, 2007; Alegret et al., 2004; Rodríguez‐Tovar et al., 2011). 2.1.4. Sections in Denmark The Kulstirenden and Højerup sections (Hart et al., 2004; Sepúlveda, Wendler, Summons, et al., 2009) are located in the northern and southern areas of the Stevns Klint region near Store Heddinge, respectively (Figure 1). The Upper Cretaceous to lower Paleogene succession is composed mainly of chalk and shallow water carbonates, deposited in a wide epeiric sea (e.g., Surlyk, 1997; Surlyk et al., 2006). The K/Pg boundary is at the base of the Fiskeler (“ﬁsh clay”), a 2‐mm‐thick red clay overlain by a dark and clay‐rich unit of variable thickness (Figure 2). At Kulstirenden, the Fiskeler reaches a thickness of 37.5 cm, at Højerup 5 cm. 2.2. Elemental and Bulk Isotopic Analyses

The calcium carbonate content (% CaCO3) was calculated by the weight difference between rock powders before and after acid treatment. Rock powders were mixed with 10% HCl for 24 hr, until no further reaction was observed. After decalcification of rock powders by HCl treatment, samples were mixed with 48% HF for 72 hr to dissolve silicates then mixed again with 6N HCl to remove fluoride precipitates. Residues were rinsed with Milli‐Q water until pH neutral, dried at 60°C, and ground to powder in an agate mortar. Samples were packed in tin boats and analyzed in triplicate for TOC, total nitrogen, and the stable isotope 13 15 composition of carbon and nitrogen in kerogen (δ Corg and δ Norg) using a Fisons (Carlo Erba) NA 1500 elemental analyzer fitted with a Costech Zero Blank Autosampler and coupled to a Thermo Finnigan Delta Plus XP isotope ratio mass spectrometer at the Massachusetts Institute of Technology. Analytical reprodu- 13 cibility (1σ) was calculated as better than 0.1% for TOC and total nitrogen and 0.1‰ and 0.5‰ for δ Corg 15 and δ Norg, respectively, by comparison with in‐house standards (acetanilide #1 and urea #1 from Arndt Schimmelmann, Indiana University, and IAEA‐CH‐6 sucrose and NBS‐22 oil). The carbon isotopic compo- 13 sition of bulk carbonates (δ Ccarbonate) was analyzed using a Finnigan MAT 253 Mass Spectrometer at the Nanjing Institute of Geology and Paleontology. About 5–10 mg of rock powder was reacted with 100% phos-

phoric acid at 25 °C, and the evolved CO2 was cryogenically puriﬁed before entering the instrument using a dual inlet system. The analytical reproducibility (1σ) was better than 0.04‰ and 0.09‰ for δ13C and δ18O,

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respectively. Isotope values were checked using a standard of known isotopic composition (GBW04405; δ13C = 0.57‰, δ18O=–8.49‰). The results are reported in conventional delta notation as (‰) relative to VPDB.

2.3. Compound‐Specific Carbon Isotope Analysis Detailed methodology for biomarker extraction and identification follows Sepúlveda, Wendler, Leider, et al. (2009) and Sepúlveda, Wendler, Summons, et al. (2009). Briefly, rock powders were extracted (three times) with DCM:MeOH (9:1 v:v) using a Dionex ASE 200 accelerated solvent extractor. Total lipid extracts were concentrated in the presence of acid‐activated copper to remove elemental sulfur, and mal- tenes were separated after precipitation of asphaltenes in hexane. Saturated and aromatic hydrocarbons and polar compounds were separated by silica gel chromatography using n‐hexane, n‐hexane:dichloromethane (4:1 v/v), and dichloromethane:ethyl acetate (5:5 v/v), respectively. The compound‐specific isotope analysis of the aliphatic hydrocarbon fraction, particularly the acyclic isoprenoid phytane 13 (δ Cphytane), was performed by gas chromatography‐combustion‐isotope ratio‐mass spectrometry (GC‐ C‐IR‐MS) at the Massachusetts Institute of Technology, using a Trace GC fitted with a PTV injector and equipped with a Varian DB‐1 (60‐m length, 0.32‐mm inner diameter, and 0.25‐mm film thickness) fused‐silica capillary column. The GC was coupled to a Thermo Finnigan Deltaplus XL IR‐MS via a combustion interface at 850 °C. The GC oven temperature program was 60 °C (2 min) to 150 °C at 10 °C/min, to 315 °C at 4 °C/min, and then held isothermal for 20 min. Carbon isotope ratios were determined rela-

tive to an external CO2 standard, regularly calibrated relative to a reference mixture of n‐alkanes (Mixture B; Arndt Schimmelmann, Indiana University). The analytical reproducibility (1σ) was better than 0.2‰.

3. Results 3.1. Sections in Tunisia

The CaCO3 content at El Kef and Aïn Settara varied between 33% and 50% in the uppermost Maastrichtian, with a sudden decrease to ~13–20% in the lower 10 cm of the boundary clay layer (Figure 3a,a′). The TOC content generally ranged between 0.2% and 1.0%, except for the uppermost sample in Aïn Settara 13 (Figure 3b,b′). The δ Ccarb varied between −2.5‰ and +1.4‰ (Figure 3c,c′), with the most isotopically enriched values in the Maastrichtian and a negative CIE of about 3‰ in the lowermost clay layer. In both 13 sections, the bulk δ Ccarb did not recover to preboundary values within the studied interval. Despite the low sampling resolution, both localities displayed similar temporal trends, except for a slightly more pro- 13 nounced negative CIE in Aïn Settara. The δ Corg varied between −27.2‰ and −25.3‰ at El Kef and between −26.9‰ and −26.1‰ at Aïn Settara (Figure 3d,d′). A negative CIE excursion occurs in the lowermost clay layer in both sections, although with different magnitudes (~2‰ in El Kef and ~0.6‰ at Aïn 13 Settara). At El Kef, δ Corg did not return to preboundary values within the ﬁrst 1.2 m of the Danian, whereas 13 at Aïn Settara preboundary values returned at about +10 cm. The δ Cphytane varied between −33.8‰ and −30.6‰ in El Kef, with a pronounced although brief negative CIE in the lowermost clay layer 15 (Figure 3d). Finally, the δ Norg varied between +3.5‰ and +5.7‰ at El Kef and between +5.8‰ and +7.7‰ at Aïn Settara (Figure 3e,e′), with both sections displaying a 1–2‰ isotopic enrichment in the immediate aftermath of the K/Pg boundary and above.

3.2. Sections in Spain

The CaCO3 content varied between 31% and 87% and between 41% and 87% at Agost and Caravaca, respectively (Figure 4a,a′). The lowest CaCO3 content was in the lowermost 5 cm of the clay layer, and preboundary values were reached within the upper half of the clay layer. The TOC content varied between 0.2% and 13 0.5% in Agost and between 0.3% and 0.4% in Caravaca (Figure 4b,b′). The δ Ccarb varied between −0.27‰ and +1.85‰ at Agost and between +0.05‰ and +1.98‰ at Caravaca (Figure 4c,c′). In both sections, 13C‐ enriched values were observed in the uppermost Cretaceous, followed by a negative CIE in the lowermost 5 cm of the clay layer. A return to more isotopically enriched values was observed in the upper half of the clay layer in both sections, although preboundary values were not fully reached within the studied interval. 13 The δ Corg varied between −24.2‰ and −22.5‰ in Agost and between −25.3‰ and −22.8‰ in Caravaca (Figure 4d,d′). A negative CIE at the K/Pg boundary was followed by a smooth recovery to preboundary 13 values within the clay layer in Caravaca, in contrast to a positive δ Corg excursion observed in Agost. The

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Figure 3. Elemental and isotopic composition of the El Kef (upper panels) and Aïn Settara (lower panels) sections in Tunisia. (a,a′) Calcium carbonate content (white symbols this study and dark symbols from Keller & Lindinger, 1989). (b,b′) Total organic carbon content (TOC; white symbols this study and dark symbols from Keller & Lindinger, 1989). (c,c′) 13 Carbon isotopic composition of bulk carbonate (δ Ccarb; white symbols this study and dark symbols from Keller & 13 Lindinger, 1989). (d,d′) Carbon isotopic composition of total organic carbon (δ Corg; upper X axis) and phytane 13 15 (δ Cphytane; lower X axis). (e,e′) Nitrogen isotopic composition of total organic nitrogen (δ N). Gray area represents the position of the boundary clay layer at each location.

13 δ Cphytane varied between −31.7‰ and −30.3‰ in Agost, with a positive CIE across the K/Pg boundary and 15 throughout the clay layer (Figure 4d). No reliable δ Norg measurements were obtained for Agost and Caravaca, and we use the Agost record from Martínez‐Ruiz (1994), which shows stable values between +4‰ and +4.5‰ throughout the K/Pg boundary (Figure 4e).

3.3. Sections in France

The CaCO3 content varied between 35.8% and 85.6% at Bidart and between 50.3% and 66.0% at Loya (Figure 5a,a′). At Bidart, CaCO3 content fell to 35% in the lower 5 cm of the clay layer, followed by a rapid return to preboundary values. At Loya, the %CaCO3 dropped to 30% at about 4 cm above the K/Pg boundary and remained at ~50–60% within the clay layer. The TOC content ranged between 0.2% and 0.3% in both sec- 13 tions with little variation (Figure 5b,b′). The δ Ccarb varied between −1.5‰ and +1.8‰ at Bidart and 13 between +0.3‰ and +1.5‰ at Loya (Figure 5c,c′). A 3‰ negative CIE in δ Ccarb was observed at Bidart in the lowermost 5 cm of the clay layer above the K/Pg boundary, followed by a rapid enrichment toward preboundary values at ~30 cm above the boundary. At Loya, a single negative CIE (~2.5‰) was observed 13 at ~4 cm, followed by rather stable values around 0.5–1‰ in the remaining of the section. The δ Corg varied between −24.2‰ and −23.8‰ at Bidart and between −24.6‰ and −24.0‰ at Loya (Figure 5d,d′). Both

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Figure 4. Elemental and isotopic composition of the Agost (upper panels) and Caravaca (lower panels) sections in Spain. (a,a′) Calcium carbonate content. (b,b′) Total organic carbon content (TOC). (c,c′) Carbon isotopic composition of bulk 13 13 carbonate (δ Ccarb). (d,d′) Carbon isotopic composition of total organic carbon (δ Corg; upper X axis) and phytane 13 15 (δ Cphytane; lower X axis). (e,e′) Nitrogen isotopic composition of total organic nitrogen (δ N). Gray area represents the position of the boundary clay layer at each location.

sections displayed minor variability across the K‐Pg, except for a moderate (~0.7‰) negative isotopic 15 excursion in the lowermost 5 cm of the clay layer at Loya. No reliable δ Norg measurements were obtained.

3.4. Sections in Denmark The bulk elemental and isotopic composition of the Kulstirenden section has been reported elsewhere

(Sepúlveda, Wendler, Summons, et al., 2009). The CaCO3 content varied between 12% and 98% in both Kulstirenden and Højerup (Figure 6a,a′), with the most dramatic decrease at the K/Pg boundary, followed by a smooth recovery to preboundary values toward the upper part of the clay layer. The TOC content varied between 0.03% and 1.55% at Kulstirenden and between 0.02% and 0.20% at Højerup (Figure 6b,b′). At Kulstirenden, the maximum TOC% was in the lower half of the clay layer, whereas in Højerup no major var- 13 iations where observed. The δ Ccarb varied between +1.3‰ and +1.8‰ in Kulstirenden and between +1.2‰ and +1.8‰ in Højerup (Figure 6c,c′). A small ~0.5‰ positive isotopic excursion is recorded at the K/Pg boundary at Kulstirenden and a brief ~0.5‰ negative isotopic excursion in the lowermost clay layer 13 at Højerup. The δ Corg at Kulstirenden was highly variable, ranging between −27.0‰ and −24.4‰, with 15 the most depleted values immediately above the K/Pg boundary (Figure 6d). The δ Norg at Kulstirenden varied between −2.2‰ and +8.8‰, with a sharp negative isotopic excursion at the K/Pg boundary and mini- 13 15 mum values in the lowermost 5 cm of the clay layer (Figure 6e). No δ Corg and δ Norg data were obtained for the Højerup section due to low organic content and sample limitation.

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Figure 5. Elemental and isotopic composition of the Bidart (upper panels) and Loya (lower panels) sections in France. (a,a ′) Calcium carbonate content. (b,b′) Total organic carbon content (TOC), (c,c′) Carbon isotopic composition of bulk car- 13 13 bonate (δ Ccarb); (d,d′) Carbon isotopic composition of total organic carbon (δ Corg). Gray area represents the position of the boundary clay layer at each location.

4. Discussion 4.1. Elemental Composition and Carbonate Sedimentation The TOC and carbonate contents in our sections are within the range of error reported in previous studies for El Kef (Keller & Lindinger, 1989), Aïn Settara (Dupuis et al., 2001), Agost (Martínez‐Ruiz et al., 1992), Caravaca (Kaiho & Lamolda, 1999), Bidart (Bonté et al., 1984), Kulstirenden (Hart et al., 2004; Sepúlveda, Wendler, Summons, et al., 2009), and Højerup (Khristensen et al., 1973). To the best of our knowledge, no elemental and isotopic data are available for Loya. Slight differences with previous reports are likely due to natural variability in outcrop material and different analytical procedures. Although subject to changes in preservation and dilution by clastic material, the carbonate content of sedimentary rocks can be used as a rough proxy for the bulk production and preservation of calcareous organisms (e.g., D'Hondt, 2005), particularly for open ocean sites deposited above the calcite compensation depth, where the low carbonate content above the K/Pg is likely not related to dissolution (Zachos et al., 1989; Zachos & Arthur, 1986). However, dilution by clays caused by the bolide impact itself (short‐term

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Figure 6. Elemental and isotopic composition of the Kulstirenden (upper panels; Sepúlveda, Wendler, Summons, et al., 2009) and Højerup (lower panels) sections in Denmark. (a,a′) Calcium carbonate content; (b,b′) total organic carbon 13 content (TOC). (c,c′) Carbon isotopic composition of bulk carbonate (δ Ccarb). (d) Carbon isotopic composition of total 13 15 organic carbon (δ Corg). (e) Nitrogen isotopic composition of total organic nitrogen (δ N). Gray area represents the position of the boundary clay layer at each location.

influence) or by clastic input impacts the carbonate content of sedimentary rocks deposited along continental margins (short‐ and long‐term influence). The abrupt reduction in carbonate content at the K/Pg in all studied sections indicates disruption of calcification due to the global extinction of pelagic calcifiers, consistent with the observed extinction of calcareous nannoplankton and planktonic foraminifera at El Kef (Arenillas et al., 2002; Pospichal, 1994), Ain Settara (Dupuis et al., 2001), Caravaca (Kaiho & Lamolda, 1999; Romein, 1977; Smit, 2004), Agost (Molina et al., 1996), Bidart and Loya (Apellaniz et al., 1997; Peyberne's et al., 1996), and Kulstirenden and Hojerup (Hart et al., 2005). Above the extinction horizon, carbonate sedimentation returned to preboundary values in the uppermost portion of the clay layer at most localities, except for the sections in Tunisia (Figures 3–6 and 8), where full carbonate sedimentation returned about 3 m above the K/Pg boundary (Dupuis et al., 2001; Keller and Lindingen, 1989). Thus, the bulk of carbonate sedimentation in the western Tethys and northeastern Atlantic basins recovered fairly rapidly to preboundary levels, that is, mostly within the deposition of the clay layer (<10 Kyr; Mukhopadhyay et al., 2001; Giron et al., 2012; Giron, 2013). The TOC content, in contrast, experienced little variability across the K/Pg boundary at multiple localities, except for slight increases in El Kef (Figure 3b) and Spain (Figures 4b,b′) and a marked increase in Kulstirenden (Figure 6b) right above the impact layer, which may be related to changes in organic matter preservation due to the potential occurrence of low oxygen conditions at the water sediment interphase during the earliest Danian (Kaiho et al., 1999; Keller & Lindinger, 1989; Sepúlveda, Wendler, Summons, et al., 2009). At Caravaca and Agost, lithology, variations in sulfur isotope ratios, bulk geochemical parameters, foraminiferal indices, and redox‐sensitive elements all imply a decrease in dissolved oxygen in the early

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Danian (Martínez‐Ruiz et al., 1992; Kaiho et al., 1999; Alegret et al., 2003; Vellekoop et al., 2018). Similarly, at Kulstirenden, nitrogen isotopes (Sepúlveda, Wendler, Summons, et al., 2009), redox‐sensitive elements, and micropaleontological records (Vellekoop et al., 2018) are consistent with reduced oxygenation in the early Danian. At El Kef (Figure 2b), evidence for low oxygen conditions is not consistent: published data on benthic foraminifera show low diversity assemblages but the dominant taxon (Cibicidoides pseudoacutus, Alegret, 2003, 2008) is not indicative of low oxygen conditions, although ostracod faunas have been interpreted as potentially indicative of low oxygen (Peypouquet et al., 1986). Enrichment in δ15N values could also be seen as indicative of low oxygen (see section 4.4 below). 4.2. Carbon Isotopes in Bulk Carbonate 13 The δ Ccarb, either as bulk carbonate or its fine fraction, has been commonly used to provide a rough, inte- 13 grated signal of δ CDIC in surface waters and to assess changes in the exogenic global carbon cycle across the K/Pg boundary (e.g., Coxall et al., 2006; D'Hondt, 2005; D'Hondt et al., 1998;Hsü et al., 1982 ; Zachos 13 & Arthur, 1986). However, bulk δ Ccarb values integrate the isotopic signature of multiple carbonate 13 sources and, compounded by diagenetic processes, can mask the original δ Ccarb, particularly during the very lowermost Danian when calcareous plankton is very rare or absent (Minoletti et al., 2005). 13 Consequently, we emphasize that δ Ccarb should be carefully interpreted when inferring variations 13 in δ CDIC. 13 13 Our δ Ccarb data are generally consistent with published δ C records based on bulk and fossil‐specific carbonate from Tunisia (Keller & Lindinger, 1989; Stüben et al., 2002), Spain (Kaiho et al., 1999; Kaiho & Lamolda, 1999; Martínez‐Ruiz et al., 1992; Smit, 1990), France (Bonté et al., 1984), Denmark (Hart et al., 2004; Schmitz et al., 1992), Brazos River, Texas, USA (Barrera & Keller, 1990), and the eastern Tethys (Vellekoop et al., 2017). Except for Kulstirenden (Denmark), all our records show the characteristic shift to negative δ13C values at the K/Pg boundary (Figures 3–6 and 8), which has been traditionally used as evidence for a global perturbation of the carbon cycle due to an extreme drop in productivity or export productivity (Hsü et al., 1982; Zachos & Arthur, 1986). However, both the magnitude and duration of the negative 13 CIEs in δ Ccarb differ among the sections (Figure 8), and the recovery to preboundary values also varies regionally. This observation in itself is inconsistent with a dominantly globally driven, synchronous perturbation of the carbon cycle. For instance, our southernmost sections (Tunisia) do not show a full return to preboundary values within the span of our record (i.e., up to planktonic foraminiferal Zone Pα), but the 13 Spanish and French sections further to the North show a rebound of δ Ccarb close to preboundary values within the limits of the boundary clay layer (Figures 3–6 and 8). Geochemical and micropaleontological records from Caravaca, Spain, indicate that normal primary productivity commenced ~13 Kyr after the K/Pg boundary (Kaiho et al., 1999). The northernmost Danish sections, on the other hand, indicate either a brief (1 data point) and small (~0.5‰) CIE in Højerup or a slight (~0.5‰) positive isotopic excursion at Kulstirenden, within the deposition of the Fiskeler. The small variability and opposite trend to the global patterns seen in the Danish sections, located in a wide and shallow epeiric sea, could reflect local rather than global processes, such as enhanced marine productivity (e.g., Surlyk, 1997; Surlyk et al., 2006). Biomarker data show that marine productivity at these sites was reduced only briefly (~102–103 years; Sepúlveda, Wendler, Summons, et al., 2009). Additionally, local processes involving lithofacies controls cannot be ruled out (Cao et al., 2018; Richardson et al., 2019), although all sections transition to a clay‐rich boundary layer in the earliest Danian. In general, the more rapid recovery of the CIEs toward more northern latitudes in our sites appears to be consistent with the possibility of an overall reduced ecological perturbation at higher latitudes (Alegret & Thomas, 2013; Barrera & Keller, 1994; Hollis et al., 1995; Jiang et al., 2010; Keller et al., 1993), although this is not universally accepted (Witts et al., 2016). 13 13 Assuming that our δ Ccarb records trace variations in δ CDIC resulting from changes in marine primary and export productivity, the period of maximum disruption following the K/Pg boundary (except for the tropical Tunisian sections) must have been similar to or shorter than the period of deposition of the boundary clay layer, estimated as ~10 Kyr (Arinobu et al., 1999; Giron, 2013; Giron et al., 2012; Mita & Shimoyama, 1999; Mukhopadhyay et al., 2001). This observation contradicts the common view that the negative CIEs represent a global collapse in marine primary productivity lasting for 105–106 years, as reconstructed from deep‐sea sites (e.g., Coxall et al., 2006; D'Hondt, 2005; D'Hondt et al., 1998; Hsü et al., 1982; Kump, 1991; 13 Zachos et al., 1985; Zachos et al., 1989; Zachos & Arthur, 1986). Our δ Ccarb records in neritic and upper

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13 bathyal settings show changes in the carbon cycle and δ CDIC that are more complex, heterogenous, and 13 different from what is typically seen in the δ Ccarb of deep‐sea carbonate (e.g., Coxall et al., 2006; D'Hondt, 2005; D'Hondt et al., 1998; Hsü et al., 1982; Zachos & Arthur, 1986), with marine productivity and carbon export in neritic settings and along continental margins possibly resurging earlier than in open ocean areas. 13 Several mechanisms may explain the observed heterogeneity in δ Ccarb records. First, factors controlling 13 δ Ccarb may have varied widely geographically—that is, there was regional heterogeneity in the oceans— and collapsed or reduced primary or export productivity per se cannot account for the entire global variabil- 13 ity in δ Ccarb records. This is consistent with geochemical, stable isotope, and micropalentological records (Alegret et al., 2012; Esmeray‐Senlet et al., 2015; Hull & Norris, 2011; Kaiho et al., 1999; Sepúlveda, Wendler, Summons, et al., 2009; Sibert et al., 2014) documenting post extinction geographic heterogeneity in ecosystem response and primary or export productivity. The most obvious explanation of the shorter duration of the negative CIEs in neritic and upper bathyal regions than in open ocean sites is that primary productivity and export may have resumed faster in the former environments, at least in the western Tethys and northeastern Atlantic Ocean. 13 Second, the bulk δ Ccarb is influenced by the vectors of the isotopic signal, thus by changes in carbonate 13 sources due to the mass extinction of calcifiers. For instance, the upper Maastrichtian δ Ccarb was dominated by calcareous nannoplankton and large Cretaceous planktonic foraminifera, whereas the lowermost Danian was dominated by an admixture of 13C‐depleted calcareous dinoflagellates and small, nonsymbiont bearing bitriserial planktonic foraminifera (Alegret et al., 2012; Birch et al., 2016). This biotic replacement was global (Apellaniz et al., 1997; Arenillas et al., 2002; Brinkhuis & Zachariasse, 1988; Perch‐Nielsen et al., 1982; Peybernés et al., 1996) and may explain at least part of the signal, that is, the more 13C‐depleted 13 13 values observed in δ Ccarb than in the δ C of planktonic foraminifera (e.g., Birch et al., 2016; Esmeray‐ 13 Senlet et al., 2015; Kaiho & Lamolda, 1999). For instance, the δ Ccarb of ODP Site 1262 (southeast Atlantic Ocean) is ~0.5–1.0‰ more depleted than the δ13C of planktonic foraminifera for ~700–800 Kyr after 13 the K/Pg boundary (Birch et al., 2016), whereas in the shallow‐water Bass River section, NJ δ Ccarb values are ~1–2‰ more depleted than in planktonic foraminifera (Esmeray‐Senlet et al., 2015). Additionally, the extinction of calcareous nannoplankton and planktonic foraminifera in the early Danian resulted in a change in the proportion of biogenic and nonbiogenic carbonates with different isotopic signatures across the K/Pg boundary. For instance, in Bidart around 50–80% of the bulk carbonate in the lowermost Danian sediments appears to be largely nonbiogenic, 13C‐depleted carbonates (Minoletti et al., 2005). Preliminary analysis of a small set of samples from El Kef shows that diagenetic carbonate (e.g., dolomite crystals) represents a significant fraction of the bulk carbonate above the K/Pg boundary (F. Minoletti, 2012, personal communication). We thus argue that at least part of the globally observed negative CIE in 13 13 δ Ccarb may be due diagenetic components, as supported by the strong correlation between the δ Ccarb and CaCO3 content in our sites, particularly in the Spanish and French sections (Figure 7). Third, we cannot rule out post depositional diagenetic alteration in the studied sections. For instance, there 13 18 is a variable degree of correlation between δ Ccarb and δ Ocarb, especially in Caravaca, Agost, Aïn Settara, and Bidart (Figure 7), largely driven by samples within the negative CIE. A significant positive relationship 18 13 between δ Ocarb and δ Ccarb can result from the interaction of carbonate and meteoric ground waters or varying carbonate sources (Mitchell et al., 1997; Sakai & Kano, 2001; Schrag et al., 1995; Swart & Eberli, 18 2005). To test the potential influence of diagenetic overprint, we used δ Ocarb values to reconstruct paleo- 18 temperatures following Bemis et al. (1998) and assuming a δ O of Cretaceous seawater (δW)of−1‰ (Shackleton, 1975). This exercise yields temperatures of 28.4–29.4 ± 2.8 °C for Tunisia, 23.9–25.4 ± 3.2 °C for Spain, 19.1–26.6 ± 4.3 °C for France, and 16.0–17.4 ± 1.8 °C for Denmark. These reconstructed temperatures are higher than those derived from species‐specific planktonic foraminifera and fish debris from similar latitudes (e.g., Huber et al., 1995; Macleod et al., 2018; Pucéat et al., 2003, 2007) and suggest that 18 diagenesis influenced at least δ Ocarb values. 4.3. Bulk Organic and Compound‐Specific δ13C 13 13 Further insights into marine productivity and carbon cycling can be obtained from δ Corg and δ Cbiomarker 13 records. The δ Corg documents the average C‐isotope composition of organic carbon preserved in sediments

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Figure 7. Cross‐plots of elemental and isotopic data. (ﬁrst row) Carbon versus oxygen isotopes of bulk carbonate. (second row) Carbon isotopes of bulk carbonate versus calcium carbonate content. (third row) Carbon isotopes of bulk carbonate versus carbon isotopes of total organic carbon. (fourth row) Carbon isotopes of total organic carbon versus total organic carbon content.

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13 (e.g., Hayes et al., 1999; Meyers, 1997), and the δ Cbiomarker provides more specific information on isotopic fractionation in the photo- trophic community and specific algal groups (e.g., Freeman et al., 1994; Freeman & Hayes, 1992). Our data show spatial and temporal 13 13 heterogeneity in δ Corg records compared to δ Ccarb (Figures 3–6 and 8). A negative CIE occurs in the Tunisian sections, Caravaca, Loya, and Kulstirenden, although the latter shows a brief CIE followed by high variability (Figures 3–6 and 8). In contrast, a positive CIE occurs at Agost (Figures 4 and 8), whereas little to no isotopic change occurred in Bidart (Figures 5 and 8). We highlight that the 13 heterogeneity among δ Corg records could be explained by the influence of local processes, particularly in shallow‐water settings like Kulstirenden and Højerup and in areas where low oxygen conditions may have prevailed. We provide further explanation of the processes 13 13 affecting δ Corg below. Despite this heterogeneity, δ Corg values returned to preboundary levels within the deposition of the clay layer in all studied sections except El Kef (Figures 3–6 and 8). Even if we 13 13 assume that δ Ccarb roughly reflects variations in δ CDIC and calcu- late the bulk isotopic fractionation during carbon fixation (εTOC), our results indicate that εTOC recovers to preboundary values within the deposition of the clay layer in most sections (except in Tunisia; data not shown). 13 For δ Clipid, we focused on the abundant acyclic isoprenoid phytane 13 (δ Cphytane; Figures 3e and 4e), a biomarker derived from the degra- dation of the phytol side chain of chlorophyll‐a produced by most photosynthetic algae and cyanobacteria (Brooks et al., 1969) and commonly assumed to integrate the isotopic signature of marine primary producers (e.g., Bice et al., 2006; Sinninghe Damsté et al., 2008; Witkowski et al., 2018). Alternative sources of phytane include bacteriochlorophyll‐a and bacteriochlorophyll‐b from purple sulfur bacteria (Brooks et al., 1969), although they are unlikely to be 13 quantitatively relevant in our study sites. The δ Cphytane at El Kef and Agost, the only sections where phytane was abundant enough for compound‐specific stable isotope analysis, displays a pattern simi- 13 lar to that of δ Corg (Figures 3 and 4), thus confirming the unusual positive CIE in Agost. The most common interpretation of positive CIEs is that they reflect increases in global organic carbon burial rates (Kump & Arthur, 1999). However, this positive CIE is only seen in 13 the organic fraction in Agost and not in the δ Ccarb. Thus, these records were probably influenced by local factors affecting isotopic fractionation during carbon fixation or by biological sources with 13C‐enriched biomass. For instance, the range of isotopic fractionation among the dominant groups of marine phytoplankton can be as large as 13‰ (Wong & Sackett, 1978), whereas the range of δ13C in particulate organic carbon in coastal settings can be up to ~15‰ Figure 8. Comparison of geochemical and stable isotope signatures among sec- 13 (Rau et al., 2001). Thus, it is possible that the δ Corg and tions. (a) calcium carbonate content. (b) Carbon isotopes of bulk carbonate. (c) 13 13 δ Cphytane in Agost may have been affected by changes in the phyto- Carbon isotopes of total organic carbon (δ Corg). Gray area represents the position of the boundary clay layer at each location. To facilitate the comparison planktonic assemblage in response to local oceanographic variability. between sections with different sedimentation rates, thus thickness of the clay Two important features of the isotopic signature in the organic frac- layer, the Y axis is displayed as a relative distance to the K/Pg boundary. Each section was normalized to the thickness of its clay layer boundary, where 0 tion thus are as follows: (a) the magnitude and duration of the nega- 13 13 represents its lowermost limit at the K‐Pg boundary and 1 indicates the upper- tive CIEs in δ Corg and δ Cphytane in the lowermost Danian are most limit. 13 smaller and shorter than the CIEs of δ Ccarb in all our study sites,

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13 and the δ Ccarb of deep‐sea sediments, and (b) there is spatial and temporal heterogeneity in isotopic values in different areas. Given the number of factors that can influence the isotopic composition of marine particulate organic matter 13 compared to dissolved inorganic carbon (e.g., Rau et al., 2001), the δ Corg is expected to be more variable 13 than δ Ccarb. Multiple processes can control the isotopic signature of the organic carbon pool as preserved in the geologic record (e.g., Hayes, 1993; Hayes et al., 1999; Kump & Arthur, 1999), but their individual and/or collective influence in the negative CIE appears to have been restricted to the deposition of the boundary clay layer, estimated to have lasted ~10 Kyr (Giron, 2013; Giron et al., 2012; Mukhopadhyay et al., 2001). These processes include changes in (a) organic matter provenance (e.g., marine vs. terrestrial, 13 phytoplankton type), (b) δ CDIC, (c) partial pressure of aquatic CO2 (pCO2), and (d) burial rates of organic matter. We highlight that modeling is necessary to quantitatively disentangle the relative influence of differ- 13 13 ent processes on δ Corg and δ Cphytane. Here, we focus on the qualitative direction of isotopic change. 13 Given our current knowledge of these processes across the K/Pg, we consider that changes in δ CDIC and aquatic pCO2 were the two most likely drivers influencing the negative CIE of the organic carbon pool. a The organic matter preserved in the studied sections is mostly of marine origin (Meyers, 1992; Sepúlveda, Wendler, Summons, et al., 2009; Sepúlveda et al., 2012). Thus, changes in the relative proportion of mar- 13 13 ine and terrestrial organic matter are unlikely to have influenced δ Corg and δ Cphytane records, with the possible exception of Caravaca. An increase in plant biomarkers and terrestrial kerogen macerals in the first few centimeters above the K/Pg boundary (Arinobu et al., 2005; Kaiho et al., 1999; Mizukami et al., 13 2013) suggests that δ Corg may have been influenced by changes in organic matter sources. Also, changes in the phytoplanktonic assemblage (Wong & Sackett, 1978) and different mechanisms of carbon transport 13 and assimilation (Rau et al., 2001) can lead to large variations in δ Corg under varying oceanographic conditions. Additionally, samples with low TOC% are more likely to be influenced by the mixtures of organic matter of differing ages and thermal histories, including allochthonous sources of previously fos- silized and thermally mature organic matter entering the system through weathering of local sedimentary rocks (Sepúlveda, Wendler, Summons, et al., 2009). The latter is possibly an important factor explaining 13 13 the low variability of δ Corg in Bidart and Loya compared to δ Ccarb (Figure 6). b The injection of isotopically light CO2 into the atmosphere and its subsequent equilibration within the 13 13 surface ocean could have influenced δ CDIC and the δ C of organic and inorganic carbon produced in surface waters. This process could have affected the δ13C of planktonic calcareous organisms and thus at least partly explained the reduced isotopic gradient between planktonic and benthic foraminifera in the aftermath of the K/Pg extinction. Potential sources of isotopically light carbon include the burning of ~18–24% (Arinobu et al., 1999) or ~25% (Ivany & Salawitch, 1993) of the aboveground terrestrial biomass and/or methane release from impact‐driven slope failure and destabilization of gas hydrates (Alegret et al., 2003; Day & Maslin, 2005). 13 13 c δ Corg and δ Cphytane are influenced by variations in atmospheric pCO2, and thus aquatic pCO2, which can lead to changes in isotopic fractionation during photosynthesis (e.g., Kump & Arthur, 1999). If large amounts of isotopically light carbon were released to the atmosphere either due to the burning of continental biomass (Arinobu et al., 1999; Ivany and Salawitz, 1993) or the destabilization of gas hydrates

(Alegret et al., 2003; Day & Maslin, 2005), CO2 levels must have increased in both the atmosphere and the surface ocean. Additional sources of CO2 in the atmosphere could have derived from the Deccan Traps (e.g., McLean, 1985; Caldeira & Rampino, 1990), although recent high‐precision geochronologic studies indicate that the Deccan Traps most voluminous eruptive phases were not synchronous with the mass extinction (Burgess, 2019; Schoene et al., 2019; Sprain et al., 2019). However, there is no agree-

ment on changes in pCO2 across the K/Pg boundary. Estimates indicate an increase (up to ~2,300 ppm; Beerling et al., 2002; Day & Maslin, 2005) or a decrease (down to ~400 ppm; Huang et al., 2013;

Steinthorsdottir et al., 2016) or highly variable (Nordt et al., 2002) pCO2 values in the earliest Danian. 13 13 In order to assess the effect of changing pCO2 on δ Corg and δ Cphytane across the K/Pg boundary, we calculated changes in isotopic fractionation assuming a conservative doubling of pCO2 from ~500 ppm in the latest Maastrichtian to ~1,000 ppm in the earliest Danian under different scenarios: (a) without changes in algal growth rates and cell geometry (“b” factor; e.g., Pagani, 2014), (b) a doubling of the b factor, and (c) a doubling of the b factor and a cooling of 4 °C. This exercise shows that a doubling of atmo- 13 spheric pCO2 could lead to an isotopic depletion of ~2–3‰ in δ Cphytane resulting from elevated isotopic

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fractionation accompanying carbon fixation. Thus, the effects of this process could either partly or com- 13 13 pletely explain the negative CIE in both δ Corg and δ Cphytane. d Kump (1991) demonstrated that organic carbon burial rates must have continued following the K/Pg boundary, possibly at only slightly reduced rates (10% reduction), while organic matter production may have been important in shallow‐marine environments or even bathyal environments. Changes in carbon 13 burial should be reflected in variations in TOC content and δ Corg. However, except for the Spanish sec- 13 tions, our studied areas displayed no significant relationship between TOC content and δ Corg (Figure 7). Given the low TOC% in all studied sections, this lack of relationship is most likely due to poor organic carbon preservation.

4.4. Marine Productivity and Carbon Cycling Across the K/Pg Boundary An earlier recovery of marine productivity in neritic and upper bathyal regions may have been facilitated by the resilience of noncalcareous phytoplankton and their potential to perform photosynthesis even in the worst‐case scenario of environmental change for phototrophs expected in the aftermath of the K/Pg impact (Perez et al., 2013). Noncalcareous phytoplankton such as diatoms, dinoflagellates, naked haptophytes, and cyanobacteria may have survived through the K/Pg boundary without significant extinction (Brinkhuis & Leereveld, 1988; Brinkhuis et al., 1998; Hollis et al., 1995; Wendler & Willems, 2002; Sepúlveda, Wendler, Summons, et al., 2009; Sepúlveda et al., 2012; Alegret et al., 2012; Vellekoop et al., 2015). Certain diatom species can survive darkness in vegetative stages for 30–60 days (Peters, 1996; Reeves et al., 2011) or up to 9 months (Peters & Thomas, 1996). Additionally, resting stages of dinoflagellates, diatoms, noncalcifying haptophytes, and cyanobacteria can be viable even after 100, 40–100, 25, and 10 years of dormancy, respectively (Härnström et al., 2011; Livingstone & Jaworski, 1980; McQuoid et al., 2002; Ribeiro et al., 2011). Thus, noncalcareous phytoplankton in neritic settings, particularly those with resting stages preserved in sediments, may have endured unfavorable growing conditions such as darkness or cold temperatures for periods far beyond those predicted by climate models (e.g., Bardeen et al., 2017; Pope, 2002; Toon et al., 1982). Indirect evidence for changes in marine productivity can be inferred from our δ15N records. Variations in δ15N at our study sites can be explained either in terms of changes in water column denitrification, nitrogen fixation, and/or phytoplankton nutrient uptake if nitrate was not depleted (e.g., Altabet & Francois, 1994; Robinson et al., 2012; Sigman et al., 2009). The ~1–2‰ 15N enrichment above the K/Pg boundary in the Tunisian sections (Figure 3) may thus represent the combined effect of enhanced bacterial denitrification and phytoplanktonic nitrate assimilation resulting from elevated marine productivity in surface waters. This is consistent with El Kef's location in an outer continental platform/upper bathyal location characterized by upwelling and high nutrient conditions (Coccioni & Marsili, 2007). On the other hand, the rather invariable δ15N values in Agost (Figure 3; Martínez–Ruiz et al., 1994) suggest that nitrogen cycling in this area was not significantly affected by the environmental changes at the K/Pg boundary, despite evidence for oxygen depletion. Finally, the strong 15N depletion (~7–8‰) in the lower Danian at Kulstirenden, Denmark (Figure 6) has been explained as minimal N‐isotopic fractionation due to diminished availability and biological assimilation of nitrate under oxygen‐depleted conditions (Sepúlveda, Wendler, Summons, et al., 2009). The low oxygen conditions have been confirmed by redox‐sensitive elements and microfossil assemblages (Vellekoop et al., 2018). Lastly, limited biomarker data indicate that cyanobacteria biomass, and possibly nitrogen fixation, was not altered across the K/Pg boundary in Danish sections (Sepúlveda, Wendler, Summons, et al., 2009). The δ15N records thus indicate that changes in oxygen depletion and nutrient utilization across the K/Pg were heterogeneous and do not provide evidence for a collapse or severe reduction of marine productivity beyond the deposition of the boundary clay layer. We cannot confidently extrapolate our results beyond neritic and upper bathyal settings, but we argue that the recovery of marine productivity in oceanic regions must have followed that of continental margins, although possibly with a time lag. Furthermore, changes in the vertical δ13C gradient across the K/Pg boundary are not inconsistent with a rapid resurgence of marine productivity in the surface ocean. Indeed, the reduced surface‐to‐bottom δ13C gradient in the early Danian may have been caused by a 10% reduction in carbon burial (Kump, 1991), an increase of ~5% in the fraction of total production degraded in the upper 200 m of the ocean (D'Hondt et al., 1998), and/or a reduction of ~30–40% in organic export (Ridgwell et al., 2010; Thomas et al., 2009).

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5. Conclusions We measured the δ13C of bulk inorganic and organic carbon, in addition to a limited number of compound‐ specific δ13C and bulk organic δ15N records, in eight K/Pg boundary sections in Tunisia, Spain, France, and Denmark that represent neritic and upper bathyal depositional environments of the western Tethys and northeastern Atlantic. Our results allow us to derive the following conclusions: 1. The magnitude and duration of the negative CIE in the δ13C of inorganic and organic carbon exhibit spatial and temporal heterogeneity among the studied sections. We interpret this heterogeneity as the result of complex processes affecting the isotopic signature of both carbon pools, including variations in carbo- 13 nate provenance and diagenesis in δ Ccarb (e.g., Minoletti et al., 2005) and those controlling isotopic 13 13 fractionation during carbon fixation in δ Corg and δ Cphytane. Additional factors may include local seafloor topography and the nature of the hinterland rates of sea level change. 13 13 13 2. The δ Corg and δ Cphytane recovered to preboundary values earlier than δ Ccarb and within the deposition of the boundary clay layer (<10 Kyr), except for El Kef where this occurred above the clay layer. 13 Therefore, the environmental factors affecting δ CDIC, primary productivity, and isotopic fractionation during carbon fixation were most likely restricted to time scales of 103–104 years in our neritic and upper bathyal settings. 15 3. A limited number of δ Norg records indicate that oxygen depletion and phytoplanktonic nitrate utilization either increased in the earliest Danian or remained unchanged across the K/Pg boundary. 4. Carbon cycling and primary productivity in neritic and upper bathyal settings recovered to pre‐K/Pg boundary values earlier than in many open ocean settings, likely supported by noncalcifying phytoplankton with resting stages. This observation is consistent with modeling and proxy studies suggesting a fast recovery of neritic environments and laboratory experiments demonstrating the resilience of noncalcifying phytoplankton to unfavorable environmental conditions for photosynthesis.

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