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Compiled Cretaceous oxygen and carbon isotope data for benthic and planktonic foraminifera from Sites 257 and 258 in the southern Indian Ocean and Sites 327, 511, 689 and 690 in the southern South Atlantic (this study) shown relative to (left to right): (1) proxy estimates for Cretaceous pCO 2 including the following: blue squares with crosses: Frenelopolis conifer estimates with ± 1σ around the mean pCO 2 level (Barral et al., 2017); green triangles: liverwort δ 13 C, red circles: pedogenic carbonate δ 13 C, and crosses: leaf stomata shown with LOESS best fit line through all but conifer data (from Foster et al., 2017 compilation); (2); (3) Sr isotope seawater curve (McArthur et al., 2012); (4) regional and global Oceanic Anoxic Events (Takashima et al., 2006); (5) global large igneous province (LIP) magma flux estimated by Coffin et al. (2006); and (6) global mid-ocean ridge magma flux estimated by Müller et al. (2016). Global subduction zone length estimated by van der Meer et al. (2014). Strongly negative δ 18 O and δ 13 C values across the Aptian-Albian boundary interval are considered an artifact of a more restricted and shallower depositional basin compared to later periods (see text). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.) 

Compiled Cretaceous oxygen and carbon isotope data for benthic and planktonic foraminifera from Sites 257 and 258 in the southern Indian Ocean and Sites 327, 511, 689 and 690 in the southern South Atlantic (this study) shown relative to (left to right): (1) proxy estimates for Cretaceous pCO 2 including the following: blue squares with crosses: Frenelopolis conifer estimates with ± 1σ around the mean pCO 2 level (Barral et al., 2017); green triangles: liverwort δ 13 C, red circles: pedogenic carbonate δ 13 C, and crosses: leaf stomata shown with LOESS best fit line through all but conifer data (from Foster et al., 2017 compilation); (2); (3) Sr isotope seawater curve (McArthur et al., 2012); (4) regional and global Oceanic Anoxic Events (Takashima et al., 2006); (5) global large igneous province (LIP) magma flux estimated by Coffin et al. (2006); and (6) global mid-ocean ridge magma flux estimated by Müller et al. (2016). Global subduction zone length estimated by van der Meer et al. (2014). Strongly negative δ 18 O and δ 13 C values across the Aptian-Albian boundary interval are considered an artifact of a more restricted and shallower depositional basin compared to later periods (see text). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.) 

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A compilation of foraminiferal stable isotope measurements from southern high latitude (SHL) deep-sea sites provides a novel perspective important for understanding Earth's paleotemperature and paleoceanographic changes across the rise and fall of the Cretaceous Hot Greenhouse climate and the subsequent Paleogene climatic optimum. Both new and prev...

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... δ 13 C and δ 18 O data from Site 511 presented in Table 8 and Fig. 6 were generated mostly from the uppermost Aptian through upper Albian interval. Additional late Aptian-late Campanian data were compiled from several sources ( Huber et al., 1995;Fassell and Bralower, 1999;Bice et al., ...
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... and carbon isotope data from ODP Sites 689 and 690 in Fig. 6 are compiled from Barrera and Huber (1990), , Barrera and Savin (1999), Huber et al. (2002), and Friedrich et al. (2006). They are plotted using the revised age models discussed ...
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... isotope ratios for planktonic foraminifera from samples at Hole 327A and Site 511 are notable for their remarkably low values of −4.2‰ to −4.7‰ that suggest temperatures of 29-32 °C during the Turonian, as discussed above (Fig. 6). Prior to this interval, a gap spans the entire Cenomanian, and, below that, Albian planktonics exhibit The trends summarized above indicate large changes in surface to seafloor isotopic gradients. A vertical δ 18 O and δ 13 C gradient is essen- tially absent within the AABI excursion interval as the extremes of benthic and planktonic ...
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... Campanian-Maastrichtian samples from Maud Rise (Fig. 6 (Fig. 7). High latitude Paleogene isotopic trends have been discussed in many well cited studies (e.g., Kennett and Stott, 1990;Bohaty and Zachos, 2003;Mackensen and Ehrmann, 1992;Thomas and Shackleton, 1996); we include those data for SHL here to facilitate comparison of the two most recent and arguably best studied greenhouse ...
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... and depths for planktonic foraminiferal (F) and calcareous nannofossil (N) species and base magnetic polarity chron boundaries from ODP Site 690 used to constrain the line of correlation in Fig. 6b. Plot code refers to genus-species abbreviations; FAD = first appearance datum. ...
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... temperature simulations at the same CO 2 levels (e.g., Lunt et al., 2016), but ~30 °C temperatures at 60°S is warm regardless. In a more recent study, at 2800 ppmv CO 2 , which is above the upper limit of pCO 2 estimated by paleobarometric proxies for the Late Cretaceous (e.g., Hong and Lee, 2012;Foster et al., 2017;Barral et al., 2017; see Fig. 6), modeled SHL sea surface tem- peratures are between 10 and 20 °C ( Zhou et al., 2012). Sea surface temperatures < 20 °C at 60° paleolatitude are also predicted from lati- tudinal gradients estimated from Cenomanian and Turonian fish teeth δ 18 O values ( Pucéat et al., 2007;Martin et al., ...
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... in the timing of δ 18 O data trends among different locations, and lack of sedimentological evidence of con- temporary glaciation cast doubt on the conclusions (MacLeod and Huber, 2001). Unlike the mid-Cenomanian and early Turonian, though, bathyal temperatures inferred from early Maastrichtian benthic δ 18 O are at a Late Cretaceous minimum (Figs. 6, 7), climate model simula- tions exist that favor perennial ice accumulation (Ladant and Donnadieu, 2016), and seasonal sea ice has been proposed to have existed based on sedimentological and paleontological evidence from the Arctic ( Davies et al., 2009;Ladant and Donnadieu, 2016) and Antarctic (Bowman et al., 2013). Still, no glacial ...
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... of the oxygen isotope paleotemperature compilation for the Cretaceous southern high latitudes with estimates of variation in pCO 2 , crustal production at mid-ocean ridges and (LIPs), and lengths of continental and island arcs (Fig. 6) demonstrates that links between Cretaceous temperature variations and any individual primary forcing factors are not well established. In particular, the proxies used to esti- mate Cretaceous pCO 2 show significant inconsistencies among methods and variable degrees of correlation to long-term paleotemperature trends. For example, the ...
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... ( Bralower et al., 1997;Jones et al., 1994;Larson, 1991a,b;McArthur et al., 2012). Higher 87 Sr/ 86 Sr ratios are taken to indicate dominance of continental mountain building and chemical weathering while lower ratios indicate increased oceanic magmatism and/or erosion of large basaltic provinces on land. The mean seawater 87 Sr/ 86 Sr curve (Fig. 6) shows generally good agreement with the paleotemperature compilation, with lowest values occurring during the hot greenhouse temperatures in the Turonian-Santonian and ratios that increase minimum values during the Maastrichtian when temperatures were ...
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... in rates of Cretaceous oceanic crust production (Müller et al., 2016), in the timing of formation of LIP ( Coffin et al., 2006), and in lengths of subduction-related volcanic arcs (van der Meer et al., 2014), correspond to variations in Cretaceous SHL temperatures in some time intervals, but not in others (Fig. 6). The largest short-term variation in oceanic crustal production rate in the ~115 to ~66 Ma interval, a decrease of ~8 km 3 /yr between ~100 and ~98 Ma ( Matthews et al., 2012), occurred during a gap in the SHL δ 18 O record. However, the decrease in δ 18 O values between ~102 and ~94 Ma ar- gues for increased temperatures, the opposite ...
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... warming of ~3 °C on Seymour Island, Antarctica (Tobin et al., 2012). This event is well-delineated between 66.5 and 66.2 Ma with an ~4 °C warming of SHL at ODP Site 690 ( Barrera and Savin, 1999), as shown in Figs. 6 and 7. Inconsistencies between the timing of elevated rates of tectonic sources of CO 2 and evidence for increased SHL warming (Fig. 6) can be attributed to multiple factors, including uncertainties in: (1) the rates of lithosphere production and consumption (van der Meer et al., 2014), (2) the amount of carbon concentrated in descending slabs and dec- arbonation efficiency at subduction zones (Johnston et al., 2011), (3) estimates of the timing and volume of magma ...

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... Results are reported on the Vienna-PDB scale. Paleotemperature estimates at Site U1513 are calculated from the foraminiferal δ 18 O values, assume seawater δ 18 O of −1‰ SMOW and use the paleotemperature equation of Kim & O'Neil (1997) reformulated by Bemis et al. (1998) following the assumptions detailed in Huber et al. (2018). ...
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International Ocean Discovery Program (IODP) Expedition 369 recovered a stratigraphically complete, Santonian-aged, pelagic sedimentary sequence at Site U1513 on the Naturaliste Plateau (33°47.6084′S, 112°29.1338′E) at ~2800 m water depth. The site was located at ~60°S paleolatitude in Cretaceous times. A total of 59 samples studied from Holes 1513A and 1513B yielded 140 calcareous and nine agglutinated benthic foraminiferal taxa. Gavelinellids (Notoplanulina, Gavelinella, Anomalinoides) and Gyroidinoides spp., are the most abundant taxa overall and also dominate most samples. We also report 382 foraminiferal stable isotopic measurements and infer cooling [of ~5 °C in surface waters and 2 °C at the seafloor] based on decreases in δ¹⁸O values for benthic and planktonic taxa through the Santonian at Site U1513. Inferred cooling was accompanied by changes in planktonic and benthic foraminiferal assemblages. At Site U1513, benthic foraminiferal taxonomic richness declines by 15 species (from 20 to 25 to 5–10 taxa) towards the upper Santonian. Within this interval, epifaunal, oxic foraminifera diminish in absolute and relative abundance, and there is a parallel relative increase in opportunistic taxa and infaunal foraminifera. Changes in benthic foraminiferal communities indicate a shift from oligotrophic to eutrophic conditions in the bottom water despite the general cooling trend. This paper explores the causes and effects of the paleoenvironmental changes in the bottom waters.
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Upper Jurassic-Lower Cretaceous carbonate successions are widely exposed throughout the Sakarya Zone, Northern Turkey. The carbonates are considered as invaluable archives of palaeoceanographic and palaeoclimatic conditions of Tethys Ocean. We, here, present new micropaleontological, microfacies, and stable isotope data (δ 18 O and δ 13 C) of Lower Cretaceous car-bonate succession of the eastern part of Sakarya Zone, NE Turkey. The studied Lower Cretaceous carbonates are characterized by, from bottom to top, microfacies associations within Unit 1 and Unit 2 that were deposited in the Barremian-Albian. Unit 1 (Barremian-Aptian) is represented by the predominance of benthic foraminiferal associations and shallow marine organisms. Unit 1 shows alternations of different microfacies, including non-laminated mudstone (MF-1), autochthonous bioclastic-foraminiferal grainstone and packstone (MF-2) and intraclastic grainstone/packstone (MF-3). An abrupt palaeoenvironment change is represented by Unit 2 with deeper water microfacies associations consisting of dark grey limestones with chert nodules, mud-rich texture, reworked skeletal fragments, sponge spiculitics and a presence of the planktonic organism. Unit 2 (Albian) displays two different microfacies: reworked bioclastic packstone/wackestone microfacies (MF-4) and gradually overlying sponge spiculitic wackestone-mudstone (MF-5). The integration of microfacies and micropaleontological data implies that the Barremian-Albian interval represents the inner platform to the slope palaeoceanographic conditions revealing an overall transgressive trend, which is consistent with a significant rise in the sea level throughout Tethys margin during the Albian. The Albian sea-level rise is likely triggered by the sedimentary evolution of the basin due to the extensional tectonic regime in NE Turkey. Besides, the palaeo-temperatures are measured by the δ 18 O data that were obtained from well-preserved belemnite samples in Unit 2. Palaeotemperature analysis presents a range of 18.30-26.77 °C with an average of 23.13 °C during the Albian. Our palaeo-temperature data are conformable with the warm Cretaceous climatic conditions, which are recorded in the different parts of the Tethys margin. Therefore, this contribution provides the first insight into the palaeoclimatic conditions of the Tethys ocean for the eastern part of the Sakarya, NE Turkey.
... Climate, i.e., temperature variability during the Cretaceous, has been extensively investigated through oxygen isotopes of planktonic and benthic foraminifera (Huber et al., 1995(Huber et al., , 2002(Huber et al., , 2018Friedrich et al., 2012;O'Brien et al., 2017), bulk carbonates (Clarke and Jenkyns, 1999) and more recent clumped isotope methods (e.g., Fernandez et al., 2017;de Winter et al., 2021), biomarker TEX 86 (TetraEther indeX of 86 carbon atoms, O'Brien et al., 2017;O'Connor et al., 2019), plant fossils (Wan et al., 2011), palaeo-CO 2 concentration proxies (Bice et al., 2006;Wang et al., 2014;Barral et al., 2017) and climate model simulations Tabor et al., 2016), which reveal that peak warming of the world's oceans was in the late Cenomanian through Santonian (the Cretaceous Thermal Maximum) that gradually returned to cooler conditions by the Maastrichtian principally in response to pCO 2 drawdown. Maximum warmth of the mid-Cretaceous claimed the release of high pCO 2 by massive volcanic activity and magma fluxes, including oceanic crust formation, increased subduction-related arc volcanism, and eruption of LIPs (Eldholm and Coffin, 2000;Jones and Jenkyns, 2001;Coffin et al., 2006;Bryan and Ernst, 2008). ...
... This led to distinguishing between OC-poor from OC-rich facies. Mainly OC-poor white to grey, and red limestone and chalk, including CORBs, are reported from areas of the whole Tethys (Fig. 3, Tables 1-2, Arthur and Fischer, 1977;Hu, 2002;Wang et al., 2005;Neuhuber and Wagreich, 2009;Wendler et al., 2009Wendler et al., , 2011Mansour et al., 2020), Boreal (Jarvis et al., 2006;Voigt et al., 2010;Thibault et al., 2016;Eldrett et al., 2021), Indian (Petrizzo et al., 2017;O'Connor et al., 2019), Pacific (Friedrich et al., 2012), North Atlantic (De Graciansky et al., 1986;Stein et al., 1989;Tucholke et al., 2004;Luft de Souza et al., 2018) and South Atlantic (De Graciansky et al., 1986;Stein et al., 1989;Huber et al., 1995Huber et al., , 2018. OC-rich facies, mainly black shales and dark grey carbonates, are located in the WIS (Pratt et al., 1993;Locklair et al., 2011;Schröder-Adams et al., 2012;Joo and Sageman, 2014;Tessin et al., 2015Tessin et al., , 2019, Canadian Arctic (Pugh et al., 2014;Davies et al., 2020), west Greenland (Dam et al., 1998), northwest South America (Erlich et al., 1996;Perez-Infante et al., 1996;Alberdi-Genolet and Tocco, 1999;Rey et al., 2004;Machado et al., 2016), equatorial Atlantic (Wagner et al., 2004;Beckmann et al., 2005a;Böttcher et al., 2006;Bice et al., 2006;Beckmann et al., 2008;Flögel et al., 2008;Friedrich et al., 2008;März et al., 2009;Hofmann and Wagner, 2011;Sachse et al., 2014;Prauss, 2015;Junium et al., 2018), extreme east Asia in Songliao Basin from northeastern China (Wang et al., 2016;Jones et al., 2018;Xi et al., 2018), and specific parts of the Australo-Antarctic Gulf especially in areas of delta plains (Gallagher et al., 2005;MacLeod et al., 2020). ...
... To better understand the climate trends that were prevalent during the C-S, the sea-surface temperature and intermediate-to bottom-water temperatures are reconstructed. For this purpose, we compiled published oxygen isotope values from well-preserved benthic (δ 18 O benth ) and planktonic foraminifers (δ 18 O plank ), compiled from various oceanic sites, including the southern Indian Ocean (DSDP Site 258, Huber et al., 1995Huber et al., , 2018ODP sites 1135and 1138, O'Connor et al., 2019ODP Site 762C, Falzoni et al., 2016), southern South Atlantic (DSDP Site 511 and 327, Stein et al., 1989;Huber et al., 1995Huber et al., , 2018, North Atlantic (ODP Site 1050, Huber et al., 2002), western equatorial Atlantic (ODP Site 1259, Bornemann et al., 2008;Friedrich et al., 2008), and the equatorial Pacific (DSDP Site 463, Friedrich et al., 2012; IODP Site U1348, Ando et al., 2013). However, δ 18 O data from C-S planktonic and benthic foraminifers in the northern, northwestern and southeastern Tethys, the Boreal, and the WIS are lacking; consequently, only bulk carbonate δ 18 O values was available from these regions. ...
Article
The Coniacian-Santonian (C-S) was a time of differentiation in marine sedimentation, characterized by organic carbon (OC)-rich black shales and carbonates interpreted as the last oceanic anoxic event, OAE3, versus OC-poor white/reddish limestones, chalk, and claystones known as Cretaceous Oceanic Red Beds (CORBs). Based on compiled geochemical and isotope proxy data of more than 95 study sites and sections, two high-resolution global carbon isotope curves for C-S carbonate and organic matter (OM) were reconstructed based on statistical analysis and discriminated three main levels of short amplitude (around 0.5‰), yet globally recognizable, carbon isotope excursions. These excursions, each some 0.4 to 0.7 Ma in duration, are characterized by regionally restricted benthic anoxia and sea-level highstands that best explain the OM accumulation during the OAE3 subevents defined herein as OAE3a (late mid-Coniacian, ca. 86.9 Ma, Kingsdown Event), OAE3b (late mid-Santonian, ca. 85.0 Ma, Horseshoe Bay Event), and OAE3c (late Santonian to Santonian-Campanian Boundary Event, ca. 83.5 Ma). For a better understanding of the C-S climate evolution on a regional to global scale, a global compilation of δ18O from benthic and planktonic foraminifers and bulk carbonate was conducted and tested for pCO2 trends based on Δ13C curves. Thus, the C-S palaeoclimate can be divided into (1) a steady state phase of warm greenhouse climate during the Coniacian, followed by (2) a hot greenhouse during the early Santonian that might be consistent with the activation of the Central Kerguelen large igneous province (LIP), and (3) a longer-term cooling of the warm greenhouse climate from the mid-Santonian onwards. The mechanism controlling OC-poor versus OC-rich deposition can be attributed mainly to palaeoceanographic conditions such as water column oxygenation and circulation pattern changes during the C-S. OM-rich deposition is largely restricted to the low-latitude Atlantic and adjacent epeiric and shelf seas. Areas of enhanced oceanic circulation systems with a westward-directed Tethyan current, and regional eddies of water mass flow had negative feedback on OM accumulation and preservation during the C-S, which resulted in well-developed water column oxygen content. This, in turn, oxidized OM and led to deposition of OM-poor facies and CORBs in large parts of the Late Cretaceous oceans.