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A multi-proxy approachto decode the end-Cretaceous mass extinction
Jahnavi Punekar, Gerta Keller, Hassan M. Khozyem, Thierry Adatte,
Eric Font, Jorge Spangenberg
PII: S0031-0182(15)00461-7
DOI: doi: 10.1016/j.palaeo.2015.08.025
Reference: PALAEO 7424
To appear in: Palaeogeography, Palaeoclimatology, Palaeoecology
Received date: 17 January 2015
Revised date: 9 July 2015
Accepted date: 19 August 2015
Please cite this article as: Punekar, Jahnavi, Keller, Gerta, Khozyem, Hassan M., Adatte,
Thierry, Font, Eric, Spangenberg, Jorge, A multi-proxy approach to decode the end-
Cretaceous mass extinction, Palaeogeography, Palaeoclimatology, Palaeoecology (2015),
doi: 10.1016/j.palaeo.2015.08.025
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A multi-proxy approach to decode the end-Cretaceous
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mass extinction
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Jahnavi Punekar1, Gerta Keller1, Hassan M. Khozyem2, Thierry Adatte3, Eric Font4, Jorge
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Spangenberg3
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1Geosciences, Princeton University, Princeton, NJ 08540, USA.
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2Department of Geology, Faculty of Science, Aswan University, Aswan 81528, Egypt.
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3Institute of Earth Sciences, University of Lausanne, 1015 Lausanne, Switzerland.
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4IDL-FCUL, Instituto Dom Luís, Faculdade de Ciências, Universidade de Lisboa, Portugal, Campo
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Grande, 1749-016, Lisbon, Portugal.
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Corresponding Author:
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Jahnavi Punekar
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Phone: +1- 609-258-6482
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Email: jpunekar@princton.edu
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Abstract
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Mass extinctions generally involve a complex array of interrelated causes and are
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best evaluated by a multi-proxy approach as applied here for the end-Cretaceous mass
19
extinction. This study documents and compares the planktic foraminiferal records,
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carbonate dissolution effects, stable isotopes, and magnetic susceptibility in France
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(Bidart), Austria (Gamsbach) and Tunisia (Elles) in order to explore the environmental
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conditions during the uppermost Maastrichtian Plummerita hantkeninoides zone CF1
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leading up to the mass extinction. Planktic foraminiferal assemblages at Bidart and
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Gamsbach appear to be more diverse than those at Elles, with unusually high abundance
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(20-30%) and diversity (~15 species) of globotruncanids in the two deep-water sections
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but lower abundance (<10%) and diversity (<10 species) at the middle shelf Elles section.
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Oxygen isotopes in zone CF1 of Elles record rapid climate warming followed by cooling
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and a possible return to rapid warming prior to the mass extinction.
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The onset of high stress conditions for planktic foraminifera is observed ~50-60
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cm below the KTB at Bidart and Gamsbach, and ~4.5 m below the KTB at Elles due to
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much higher sediment accumulation rates. These intervals at Bidart and Gamsbach record
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low magnetic susceptibility and high planktic foraminiferal fragmentation index (FI) at
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Elles, Bidart and Gamsbach. An increased abundance of species with dissolution-resistant
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morphologies is also observed at Gamsbach. The correlative interval in India records
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significantly stronger carbonate dissolution effects in intertrappean sediments between
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the longest lava flows, ending with the mass extinction. Based on current evidence, this
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widespread dissolution event stratigraphically coincides with the climate cooling that
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follows the Late Maatrichtian global warming and may be linked to ocean acidification
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due to Deccan volcanism. The estimated 12,000–28,000 Gigatons (Gt) of CO2 and 5200–
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13,600 Gt of SO2 introduced into the atmosphere likely triggered the carbonate crisis in
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the oceans resulting in severe stress for marine calcifiers leading to the mass extinction.
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1. INTRODUCTION
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One of the best-known European Cretaceous-Tertiary boundary (KTB) sections,
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also known as Cretaceous-Paleogene (KPB or KPg) sections, is exposed at a beach near
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Bidart, in the Basque-Cantabrian basin of southwestern France (Figs. 1, 2B, Seyve, 1990;
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Haslett, 1994). At this locality about 8 m of uppermost Maastrichtian and ~4 m of basal
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Danian sediments are exposed, including the boundary clay, an Iridium (Ir) anomaly and
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negative 13C excursion that indicate a relatively complete KTB transition (Fig. 2A, C;
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Renard et al., 1982; Bonté et al., 1984; Apellaniz et al., 1997; Font et al., 2014).
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Nevertheless, the Bidart section remained in limbo for nearly two decades because of
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uncertain age control, particularly the reported absence of the latest Maastrichtian
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nannofossil Micula prinsii zone and absence of the planktic foraminiferal zones CF1
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(Plummerita hantkeninoides) and CF2, which together are correlative with paleomagnetic
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chron C29r. This led to the assumption that the latest Maastrichtian is missing (Gallala et
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al., 2009). Subsequent paleomagnetic and microfossil studies revealed that the ~8 m of
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uppermost Maastrichtichtian sediments below the KTB were deposited during the Micula
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prinsii zone (Galbrun and Gardin, 2004) and the recent finding of P. hantkeninoides zone
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CF1 (Font et al., 2014) further confirms deposition in paleomagnetic chron C29r below
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the KTB boundary and hence a substantially complete KTB transition.
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Restudy of the Bidart section is particularly important because of the potential
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connection between the high-stress interval spanning the last 50-cm of the Maastrichtian
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and Deccan volcanism in India (Font et al., 2011, 2014). As early as the 1990s, Apellaniz
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et al. (1997) reported a drop in carbonate content and increased planktic foraminiferal test
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dissolution particularly in the KTB clay and the underlying 28-cm uppermost
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Maastrichtian sediments. This interval depleted in carbonate content is also featured by a
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loss of iron oxides (biogenic and detrital magnetite), interpreted to be the result of
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acidification linked to Deccan acid rains (Font et al., 2014; Font and Abrajevitch, 2014).
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The possible link between this dissolution interval and ocean acidification related to
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Deccan volcanism appears to be more than coincidental and warrants a fresh
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investigation of associated changes in planktic foraminiferal assemblages. Bidart
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therefore provides a unique opportunity to analyze this critical time interval in Earth
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history to understand the environmental changes in the northern mid-latitude Atlantic
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Ocean that may be related to the global effects of Deccan volcanism.
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Preliminary faunal analysis of the Bidart section reveals a planktic foraminiferal
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assemblage remarkably different from those reported for El Kef (GSSP) and Elles,
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Tunisia, and other continental shelf locations (Abramovich et al., 2002; Font et al., 2014).
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To evaluate whether this is due to different depositional settings (open ocean bathyal
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depths for Bidart versus shelf depth for Tunisia), we chose a second bathyal section,
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Gamsbach, Austria, as a control site. Gamsbach is located in the Eastern Alps with a
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palaeogeographic setting and depositional history similar to Bidart (Figs. 1, 3B; Grachev
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et al., 2005). Gamsbach contains planktic foraminiferal assemblages similar to those at
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Bidart, including a pre-KTB dissolution interval that supports the choice of Gamsbach as
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a complementary site.
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Although numerous studies have explored the KTB transition at Bidart and
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Gamsbach over the past three decades (see sections 1 and 2, supplementary material), the
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published microfossil records are generally not quantitative and at very low sample
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resolution yielding little or no information for the critical pre-extinction interval. We
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present comprehensive biostratigraphic, assemblage and stable isotope, geochemical and
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mineralogical data that focus on the rapid climatic and biotic events of zone CF1, which
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globally record the crises that led up to the KTB mass extinction. The primary objective
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of this study is to test the hypothesis that Deccan volcanism may have caused global
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climate changes and ocean acidification that directly resulted in the KTB mass extinction
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recorded in planktic foraminifera.
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We test this hypothesis based on: (1) High-resolution quantitative planktic
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foraminiferal species abundances through the uppermost Maastrichtian zones CF1-CF2 at
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Bidart and Gamsbach. (2) High-resolution biostratigraphic analysis with special emphasis
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on the presence/absence of index species (e.g., Gansserina gansseri and Plummerita
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hantkeninoides) to re-evaluate the conflicting published reports (reviewed below). (3)
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Evaluation of the palaeoclimatic and the paleoenvironmental conditions recorded in
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stable isotopes, geochemical proxies, and associated biotic events. (4) Evaluation of
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carbonate and iron oxide dissolution events based on the quality of foraminiferal test
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preservation (fragmentation index FI) and magnetic susceptibility, respectively. (5)
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Determination of the chronologic sequence of biotic, climatic and geochemical events
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through zone CF1 at Bidart and Gamsbach, as well as their regional and global
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oceanographic significance in the context of environmental perturbations related to
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Deccan volcanism. And (6) comparison with shelf sequences at Elles and El Kef, Tunisia,
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to assess the nature of environmental changes in shallow vs. deep-water environments.
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2. BACKGROUND
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2.1. Bidart and Gamsbach
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Previous studies of the Bidart and Gamsbach sections report sedimentologic,
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geochemical, paleomagnetic and microfossil biostratigraphic data. A brief summary is
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given here (See supplementary material for details).
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Bidart: Planktic foraminifera and nannofossils record a rapid decline at the KTB
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at Bidart (Gorostidi and Lamolda, 1995; Thibault et al., 2004; Apellaniz et al., 1997;
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Gallala et al., 2009), whereas benthic foraminifera switch from infaunal to epifaunal
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dominance across the KTB (Alegret et al., 2004). An Iridium anomaly of 6.3±1.1 ppb,
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enrichment of Co, Cr, Ni, As, Sb, Se and depletion of rare earth elements (REE) are
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reported in the Bidart KTB red clay layer (Delacotte, 1982; Smit and Ten Kate, 1982;
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Bonté et al., 1984). Some studies report the presence of microtektites, microspherules and
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Ni-rich crystals in the KTB red layer in the Basque sections but provide no supporting
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data (Apellaniz et al., 1997; Arz and Arenillas, 1998; Arenillas et al., 2004).
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Gamsbach: Previous studies on Gamsbach show the KTB clay enriched in Ir (6
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ppb), iron hydroxides, Co, Ni, Cr and siderophile elements and sporadic occurrence of
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pure Ni crystals, awaruite (Fe3Ni), Ni-Fe, Ni-Fe-Mo and Ni-Fe-Co alloys, cosmic dust
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and spherules of varied geochemical affinities (Grachev et al., 2005, 2008; Pechersky et
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al., 2006; Egger et al., 2009). Micropaleontological and biostratigraphic studies are
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limited due to poor carbonate preservation throughout the KTB transition (Egger et al.,
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2004, Summesberger et al., 2009, Korchagin and Kollmann in Grachev, 2009).
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Summesberger et al. (2009) reported on the cephalopod, nannofossil and planktic
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foraminiferal biostratigraphy at Gamsbach but provided no quantitative documentation.
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2.2. Deccan Volcanism
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Deccan eruptions resulted in an estimated 1.5 million km3 of lava flooding the
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Indian sub-continent (Raja Rao et al., 1999). Three main phases of eruptions are
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recognized: the initial phase-1 (~6% of the total volume) in the early late Maastrichtian
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recently dated by 40Ar/39Ar at 67.12a at the chron C30n/C29r transition
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(Schoebel et al., 2014); the main phase-2 (~80% of the total lava pile) in chron C29r
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(Subbarao et al., 2000; Jay and Widdowson, 2008; Chenet et al., 2007, 2008; Schoene et
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al., 2014) culminating in the KTB mass extinction (Keller et al., 2011a, 2012); and the
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final phase-3 (~14% of the total volume) in the early Danian chron C29n. The
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environmental effects of the three Deccan phases are determined by the tempo and
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magnitude of eruptions and the amounts of SO2, CO2, Cl and other gases released into the
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atmosphere (Self et al., 2008). A global review of the planktic foraminiferal events
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contemporaneous with Deccan phase-2 and phase-3 can be found in Punekar et al.
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(2014a).
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In the Krishna-Godavari Basin of SW India a rapid succession of four phase-2
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lava mega-flows span C29r below the KTB and mark the CF1–CF2 and Micula prinsii
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(nannofossil) zones; intertrappean sediments reveal rapid extinctions (Keller et al., 2011a,
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2012). The correlative interval in Meghalaya (NE India) is dominated (95%) by
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Guembelitria blooms, and high-stress is also marked by ocean acidification and strong
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carbonate dissolution (Gertsch et al., 2011). Schoene et al. (2015) show that the Phase-2
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volcanism itself lasted ~500-kyr into the Danian. On a global basis paleoclimatic data
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from DSDP 525A (Li and Keller, 1998a), Tunisia, (Stüben et al., 2003) and Texas (Keller
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et al., 2011b; Abramovich et al., 2011) show at least one and possibly multiple
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hyperthermal events during the CF1-CF2 global warming, which indicates complex and
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episodic climate fluctuations in the latest Maastrichtian correlative with Deccan phase-2
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(Punekar et al., 2014a).
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At Bidart and Gubbio Font et al. (2011, 2014) discovered akaganeite, an unusual
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Cl-bearing iron hydroxide preserved in a low magnetic susceptibility (MS) interval below
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the KTB. The origin of this low MS interval is explained by the loss of detrital and
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biogenic magnetites from reductive iron hydroxide dissolution due to acid rains and
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ocean acidification linked to Deccan Phase-2 (Font et al., 2014). They proposed a Deccan
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volcanic origin for akaganeite, formed by interaction of acid aerosols with the high
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atmosphere and potentially transported through the stratosphere at Bidart (Atlantic realm)
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and Gubbio, Italy (Tethys realm). If a volcanic origin for akaganeite is confirmed, this
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can provide a promising new geochemical benchmark for identifying Deccan
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environmental effects across the globe.
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3. GEOLOGIC SETTING AND LITHOLOGY
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3.1. Bidart, France
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The Bidart KTB boundary section outcrops along the Erreteguia beach 2 km north
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of Bidart and can be accessed by the national highway R.N. 10 (W 1°35', N 43°26'; Fig.
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2B). Sediments consist of hemipelagic to pelagic marls and limestones deposited at
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upper-middle bathyal depths in the Aturian Trough during the late Maastrichtian to
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Paleocene (Galbrun and Gardin, 2004; Alegret et al., 2004; Font et al., 2011). Deposition
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occurred in a flysch zone and accumulated at 3-4 cm/ky, resulting in thick marl beds
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(Seyve, 1984; Nelson et al., 1991; Clauser, 1994; Peybernes et al., 1997; Vonhof and
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Smit, 1997). Tectonic disturbance (Pyrenean orogeny) and diapirism resulted in interbed
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sliding, slickensides, and mass-flow deposits (Razin, 1989; Apellaniz et al., 1997).
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About 8 m of uppermost Maastrichtian (C29r) pink to purple marlstones and
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marls with occasional turbidites and cut by local faults are exposed below the KTB at
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Bidart and can be traced throughout the Basque basin (Fig. 2A; Apellaniz et al., 1997). At
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the base of the section analyzed is a 25-cm thick marlstone followed by ~2.5 m of marls
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with common pelecypod shells and fragments. In the ~50-cm below the KTB carbonate
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content decreases and macrofossils are absent, except at the top of this interval where
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burrows are truncated by the overlying boundary clay.
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The KTB is easily recognized by a 2 mm “rusty” layer at the base of an 8-15 cm
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thick clay layer (Fig. 2B, C; Bonté et al., 1984; Apellaniz et al., 1997). The base of the
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clayey interval is a grey to yellow silty clay overlain by red brown siltstone and a thinly
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laminated dark grey siltstone at the top. Carbonate content gradually increases in the
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overlying basal Danian claystones, which are overlain by hemipelagic limestones marked
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by alternating pink and white (occasionally glauconitic) biogenic limestones bioturbated
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near the top of the section (Apellaniz et al., 1997; Font et al., 2011). At the top of the
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outcrop is a mass-flow deposit with an erosive basal surface.
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3.2. Gamsbach, Austria
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The KTB boundary section is located in the Gamsbach valley of the Austrian
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Alps (E 14o51’50; N 47o39’51; Figs. 3B). During the KTB transition the Gamsbach area
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was located in the northwestern Tethys between paleolatitudes 20° to 30°N (Fig. 1;
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Haubold et al. 1999; Pueyo et al. 2007). The basin was formed after the early Cretaceous
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thrusting followed by transtension and subsidence due to subduction (Wagreich, 1993,
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1995). Erosion at the front of the Austro-Alpine microplate resulted in deposition of
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sediments at middle bathyal depths (600-1000 m) during the late Maastrichtian and lower
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bathyal depth (>1000 m) in the early Danian (Egger et al., 2009).
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Sediments consist of hemipelagic pelites interbedded with thin sandy turbidites
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(<15 cm) characteristic of the Nierental Formation of the Gosau group in the northern
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calcareous Alps (Wagreich and Krenmayr, 1993, 2005). The Maastrichtian is composed
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of medium gray marlstones and marly limestones. Truncated burrows mark the top of the
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Maastrichtian below the 2-cm thick clay layer that marks the KT boundary. This KT clay
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layer contains 0.2-0.4 cm thick yellowish clay at the base (Fig. 3A, C).
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4. MATERIAL AND METHODS
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Sampling at Bidart concentrated on the 3.5 m interval below the KTB with
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samples collected at 15-cm intervals for the bottom 3 m and 5-cm intervals for the top 50
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-cm. At Gamsbach, 2 m of the uppermost Maastrichtian below the KTB were sampled at
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5-6 cm intervals. In the laboratory, samples were crushed into small fragments and left
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overnight in 3% hydrogen peroxide solution to oxidize any organic carbon. The
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disaggregated sediment samples were then washed through >63 μm and >38 μm sieves
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(Keller et al., 1995). The washed residues were oven dried at 50°C. Quantitative faunal
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analysis was based on 63-150 μm and >150 μm size fractions. Each size fraction of every
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sample was split with an Otto micro-splitter to obtain approximately 300 specimens of
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planktic foraminifera (for a statistical representation of the species population). These
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were picked, sorted and mounted on micro-slides and identified. The residual sample was
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searched for rare species and index species for biostratigraphy but not included in the
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quantitative dataset. Species identification is based on standard taxonomic concepts (e.g.,
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Robaszynski et al., 1983-1984; Nederbragt, 1991; Olsson et al., 1999).
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For the foraminifera fragmentation index, a microsplitter was used to obtain
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approximately 500-700 foraminifera and fragments from the >63 mm fraction such that at
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least 100 entire tests were counted. Three categories were identified based on the quality
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of preservation: entire (nearly) perfect tests (Plate 1, Plate 3: A-G), partially damaged
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(imperfect) tests (Plate 2: A-L; Plate 3: H-L, O, P) and fragments (Plate 2: M-T; Plate 3:
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M, N, Q-S). Specimens consisting of less than two-third of an entire test were counted as
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fragments (Berger et al., 1982). Planktic foraminifera fragmentation data was obtained
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for Bidart, Gamsbach and Elles sections. Benthic foraminifera fragmentation data was
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also obtained for Bidart to account for mechanical breakage due to post-depositional
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transport and sample processing techniques.
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Stable carbon and oxygen isotope analyses were performed on whole-rock
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samples from Gamsbach for this study. These analyses were conducted using a Thermo
247
Fisher GasBench II preparation device interface with a Thermo Fisher Delta Plus XL
248
continuous flow isotope ratio mass spectrometer at the Institute of Earth Surface
249
Dynamics (IDYST) of the University of Lausanne, Switzerland. The stable carbon and
250
oxygen isotope ratios are reported in delta (δ) notation as permil (‰) deviation relative to
251
the Vienna Pee Dee belemnite (VPDB). Whole rock and clay mineral data were acquired
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from XRD analyses using SCINTAG XRD 2000 Diffractometer at the Geological
253
Institute of the University of Lausanne, Switzerland. The procedure for sample
254
processing was based on Adatte et al. (1996).
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Mass specific magnetic susceptibility (MS) was measured at the Institute Dom
256
Luís (IDL), at the University of Lisbon, Portugal with a MFK-1 (AGICO). Rock
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fragments were crushed by using an agate mortar and filled within typical cubic plastic
258
boxes of 8 cm3 in volume. MS values are reported relative to mass (m3/kg).
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5. BIOSTRATIGRAPHY: HOW COMPLETE IS THE KTB TRANSITION?
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To evaluate the stratigraphic completeness of the KTB transition we apply the
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high-resolution planktic foraminiferal zonal scheme by Li and Keller (1998a,b) and
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Keller et al. (1995, 2002a) (Fig. 4). The KTB is placed at 65.5 Ma (Gradstein et al.,
265
2004). However, the precise age of this boundary event is in flux with more recent
266
geochronologic dating suggesting an age closer to 66.0 Ma (Renne et al., 2013) and
267
additional dating still in progress. Based on cyclostratigraphy the duration for
268
paleomagnetic chron C29r is estimated at 750 ky with the base of C29r at 66.25 Ma
269
(Gradstein et al., 2004; Schoene et al., 2014; Thibault et al., this vol.).
270
Uppermost Maastrichtian Zone CF1: This zone is defined by the total range of
271
the index species P. hantkeninoides. Previous studies concluded that P. hantkeninoides is
272
absent at Bidart (Arz and Molina, 2002; Gallala et al., 2009; Galalla, 2013). However, we
273
observed this species in the 5 m below the KTB (see also Font et al., 2014) and 1.75 m
274
below the KTB at Gamsbach (Supplementary material Section 3). This indicates that in
275
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both localities the uppermost Maastrichtian zone CF1 is present. Zones CF1 and CF2 are
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equivalent to the upper part of the nannofossil M. prinsii zone, which spans the top 8 m
277
of the Bidart section (Galbrun and Gardin, 2004) and corresponds to C29r below the
278
KTB. The sediment accumulation rate for this interval is 3.2 cm/ky (800 cm/250ky) and
279
3.1 cm/ky for zone CF1. Previous studies estimated a sedimentation rate of 4 cm/ky for
280
the Maastrichtian distal sea fan at Bidart (Seyve 1990; Nelson et al., 1991; Vonhof and
281
Smit, 1997) and 2.5 cm/ky for the nearby Sopelana section (Mary et al., 1991). This study
282
suggests that the zone CF1 interval is substantially complete, although truncated burrows
283
at the top of CF1 just below the KT boundary clay suggest some erosion. Compared with
284
the middle bathyal environment at Bidart, the middle shelf depositional environment at
285
Elles, Tunisia, reveals a much higher sediment accumulation rate of 8.6 cm/ky for C29r
286
below the KTB. Based on this section, the duration of zone CF1 is estimated at ~160 ky
287
based on the KTB at 65.5 Ma (Gradstein et al., 2004). Considering the KTB at 66 Ma and
288
the C30n/C29r transition at ~66.288 Ma, zone CF1 at Elles is ~130 ky long (Renne et al.,
289
2013; Schoene et al., 2014)
290
At Gamsbach, P. hantkeninoides was identified in the top ~1.75 m of the
291
Maastrichtian for the first time in this study (Fig. 6B, Plate 1: M). Truncated burrows
292
mark the top of zone CF1 below the boundary clay similar to Bidart. Based on these
293
observations we conclude that the upper part of zone CF1 to the KTB mass extinction at
294
Gamsbach is similar to Bidart and substantially complete. The abrupt negative 13C shift
295
in bulk rock at the KTB and the presence of an erosional surface truncating burrows at
296
both Gamsbach and Bidart suggests some erosion. Biostratigraphy indicates that erosion
297
was primarily of basal Danian sediments.
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KT boundary clay Zone P0: The KTB consists of a “boundary clay” zone P0
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overlying the Maastrichtian mass extinction horizon. The boundary is easily identified on
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the basis of five globally verified criteria: (1) mass extinction of Cretaceous planktic
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foraminifera, (2) appearance of the first five Danian species within a few cm of the
302
boundary clay, (3) KTB clay and red layer, (4) an Ir anomaly and (5) the 13C negative
303
shift (Keller et al., 1995; 2011b). The KTB is also characterized by an abrupt increase in
304
magnetic susceptibility (Font et al., 2011; this study). At Bidart and Gamsbach, the KTB
305
clay is very thin (~5 cm and ~3 cm respectively) and overlies an erosion surface with
306
truncated burrows. The zone P0 clay, which is defined by the interval between the mass
307
extinction horizon and first appearance of Parvularugoglobigerina eugubina, is absent as
308
this species directly overlies the mass extinction horizon. Ir anomalies of 6.3 ppb and
309
~6.0 ppb at Bidart and Gamsbach, respectively (Bonté et al., 1984; Vonhof and Smit,
310
1997; Egger et al., 2009) are concentrated in the thin clay that represents redox conditions
311
above the erosion surface. Similarly, the 13C negative shift of 2.0 to 2.3‰ is abrupt
312
across the erosion surface in both sections (Rocchia et al., 1987; Font et al., 2014). In
313
comparison, at the stratotype El Kef and expanded Elles sections in Tunisia, the P0 clay
314
is 50 to 75 cm thick with an Ir anomaly of 18 ppb at the base and a 4‰ negative carbon
315
isotope excursion (Rocchia et al., 1996; Stuben et al., 2003). The relative time
316
represented by the condensed P0 intervals and hiatuses at Bidart and Gamsbach can be
317
estimated based on the zone P1a planktic foraminiferal assemblages.
318
Zone P1a hiatuses: This zone is defined by the total range of P. eugubina and/or
319
P. longiapertura and can be subdivided into subzones P1a(1) and P1a(2) based on the FO
320
of Parasubbotina pseudobulloides and/or Subbotina triloculinoides (Fig. 4). At the time
321
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of the P0/P1a(1) boundary only about five early Danian species had evolved and all were
322
rare as the assemblages were dominated by the Cretaceous survivor and disaster
323
opportunists Guembelitria species (review in Keller and Pardo, 2004).
324
At Bidart, zone P1a(1) directly overlies the mass extinction with common
325
Parvularugoglobigerina extensa, P. eugubina and P. longiapertura, an assemblage that is
326
known to first appear well into zone P1a(1) about 100 kyr after the KTB mass extinction
327
(Fig. 5A). This indicates that the early evolution of Danian species in P0 and lower part
328
of P1a(1) is missing due to erosion or non-deposition (Fig. 4). About 30-cm above this
329
hiatus subzone P1a(1) ends with another sudden faunal assemblage change marked by
330
dramatically decreased Guembelitria, P. eugubina and P. longiapertura, a sudden
331
appearance of abundant Chiloguembelina morsei and the FO of S. triloculinoides (Fig.
332
5A). This assemblage is indicative of subzone P1a(2) and marks another short hiatus
333
between subzones P1a(1) and P1a(2) (Fig. 4). Hiatuses at the KTB and lower Danian
334
(resulting from condensed sedimentation and/or deep-sea currents) have been
335
documented worldwide in various studies (reviews in MacLeod and Keller, 1991; Keller
336
et al., 2003, 2013).
337
At Gamsbach an early Danian hiatus is also present as evident by the diverse (12
338
species) early Danian assemblage including P. pseudobulloides, the index species for
339
subzone P1a(2) directly overlying the mass extinction horizon (Fig. 6A). This indicates
340
erosion of P0, P1a(1) and at least part of P1a(2) (Fig. 4). Another abrupt faunal change
341
and hiatus occurs at the P1a(2)/P1b boundary about 30-cm above the KTB marked by the
342
extinction of P. eugubina and P. longiapertura (index for top P1a(2)) and terminal
343
abundance decrease in Globigerina edita. Above this hiatus abundant C. morsei and
344
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common P. pseudobulloides followed by abundant Guembelitria spp. indicates zone P1b
345
(Figs. 4, 6A).
346
347
6. STABLE ISOTPES AND FAUAL TURNOVER
348
349
6.1. Stable isotopes
350
Whole-rock stable carbon and oxygen isotope data for Gamsbach (Austria) were
351
obtained for this study (supplementary materials Table 7). Planktic, benthic and bulk
352
stable isotope data for Elles (Tunisia) and whole-rock isotope data for Bidart (France)
353
have already been documented (Stüben et al., 2003; Thibault et al., this vol.; Font et al.,
354
2014). Visual inspection of preservation and degree of recrystallization of individual
355
foraminifera tests indicated that the Gamsbach isotope data were likely to be the most
356
compromised.
357
At Elles, the overall low 18O values (-7.0 to -4.0‰) for planktic as well as
358
benthic (-4.5 to -2.0‰) foraminifera indicate diagenetic effects but long term trends may
359
still be preserved (Stüben et al., 2003; supplementary materials Table 5). At Bidart, the
360
whole-rock 13C values range between -1.7 and 1.8‰ and bulk 18O values range
361
between -3.2 and -0.3‰. Plotting 13C vs. 18O values yields a correlation coefficient
362
R2=0.53, suggesting that diagenetic alteration of the primary signal may not be ruled out
363
(supplementary materials Table 6). A low 13C event is recognized between 3.5 m and
364
0.75 m below the KTB. A similar event is also observed in the planktic 13C values at
365
Elles between 1.75 m and 6.6 m below the KTB boundary indicating that this signal may
366
be real (Fig. 7). A long term increasing trend through zone CF1 is observed in 18O
367
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profiles of Bidart, Gamsbach and Elles, although the values at Gamsbach and Elles
368
record frequent fluctuations (Fig. 8). The whole-rock 13C values for Gamsbach range
369
between 1.41 to 2.42‰ and the bulk 18O range between -2.6 and -1.1‰. The 13C vs.
370
18O correlation coefficient is lower than that for Bidart (R2=0.27) but the poorly
371
preserved fragmented and recrystallized tests suggest considerable overprint on the
372
primary isotopic composition. The low 13C event of Bidart and Elles is not preserved at
373
Gamsbach.
374
375
6.2. Faunal Turnover
376
Bidart (France)
377
Maastrichtian planktic foraminifera at Bidart are recrystallized but relatively well
378
preserved and identification is fairly easy. About 51 species were identified in the 63-150
379
m size fraction and 23 species in the >150 m size fraction (Fig. 5A). The species
380
richness in the 63-150 m fraction gradually drops from 30 to 20 through the analyzed
381
interval of zone CF1. A rapid decline from ~20 species to 3 species is seen 3-cm below
382
the KTB. A brief increase to 13 species occurs 2-cm above the boundary clay is observed
383
in the small and large size fractions and is likely due to erosion and redeposition (Fig.
384
5A). In the >150 m fraction, diversity remains nearly constant through CF1 but declines
385
from ~50 species to 11 species at 3-cm below the KTB.
386
All typical late Maastrichtian globotruncanids, rugoglobigerinids, heterohelicids
387
and pseudoguembelinids are represented, although the biserial heterohelicids and
388
pseudoguembelinids dominate the assemblages in the 63-150 um (Fig. 5B). In addition to
389
common cosmopolitan species, biserial species Hartella harti and Spiroplecta americana
390
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Ehrenberg are also frequent in the assemblage (Plate 1: K, J); these species were first
391
described by Georgescu and Abramovich (2009) from upper Maastrichtian sediments of
392
the Atlantic Ocean. At Bidart H. harti and Heterohelix navarroensis are the most
393
abundant and together constitute 40-60% of the assemblage throughout CF1 (Figs. 5B, 7;
394
Plate 1: K, L). Guembelitria sp. is present in the 63-150 m size fraction but not in the
395
abundance observed in shallow marine KTB transitions (e.g. Egypt, Sinai, Tunisia
396
(Seldja); Keller and Benjamini, 1991; Keller et al., 1997; Keller, 1998b; 2002b; Punekar
397
et al., 2014b; Plate 1: I). The newly evolved Danian assemblage in zone P1a(1) is
398
dominated by Guembelitria sp. and Parvularugoglobigerina sp. (30-40%, Fig. 5A). The
399
P1a(2) assemblage is dominated by Chiloguembelina midwayensis. (See supplementary
400
material Fig. S1 for planktic foraminifera in the >150 m size fraction at Bidart)
401
Gamsbach (Austria)
402
A planktic foraminiferal study by Korchagin in Grachev et al. (2005) identified
403
only 25 Maastrichtian species and placed this assemblage in the Abathomphalus
404
mayaroensis zone, which spans most of the late Maastrichtian (68.72 – 65.5 Ma). The
405
preservation of Maastrichtian planktic foraminifera in the 63-150 m fraction at
406
Gamsbach is very poor. Recrystallization and the difficulty of freeing specimens from
407
surrounding sediments result in specimens with a highly fragmented and abraded
408
appearance and no reliable quantitative data can be obtained for the Maastrichtian
409
(supplementary material Fig. S2). In the >150 m fraction preservation is better and
410
therefore was analyzed quantitatively (Fig. 6B). A total of 46 species were identified in
411
the >63 m size fraction which is likely an underestimate of the total assemblage due to
412
poor preservation. Most species are consistently present in the lower 1.25 m of the
413
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section. But in the uppermost 0.5 m, species are more sporadic and species richness drops
414
from 46 to 21-30 species. In the >150 m fraction, species richness ranges between 30-40
415
species and drops from ~25 to 8 species at the KTB (Fig. 6A).
416
All common Maastrichtian groups such as the globotruncanids, rugoglobigerinids,
417
heterohelicids and pseudoguembelinids are present in the assemblage. Pseudotextularia
418
elegans, Pseudotextularia nuttali, Pseudoguembelina hariaensis, Heterohelix globulosa
419
and Planoglobulina brazoensis dominate the >150 m fraction (Fig. 6B). Guembelitria
420
sp. is almost absent in the 63-150 m fraction. In contrast to Bidart, H. harti and S.
421
americana are not present in the assemblage (Fig. 8C). In the Danian, the diversity in
422
zone P1a(2) is about 10 species which increases to ~20 species in zone P1c. The P1a(2)
423
assemblage overlying the KTB is dominated by G. cretacea, P. eugubina, P.
424
longiapertura, Globigerina edita, Globanomalina archaeocompressa and Praemurica
425
taurica (Fig. 6A).
426
427
6.3. Depth-ranked species
428
Planktic foraminifera species have been classified into surface dwelling
429
opportunistic species Guembelitria, surface-subsurface mixed layer, intermediate or
430
thermocline and deep-water dwellers based on stable oxygen and carbon isotope ranking
431
of well-preserved specimens (Abramovich et al., 2003, 2010). The diversity and
432
abundance changes for each depth group can indicate climatic and environmental effects
433
at different depths of the water column. Figure 9 shows the diversity and abundance of
434
the four groups in small (63-150 m) and larger (>150 m) size fractions analyzed
435
through zone CF1 at Bidart and Gamsbach and compared with Elles, Tunisia
436
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(Supplementary material Section 5, Table 1 show the depth-ranked grouping of species
437
used for this study).
438
At Bidart, the small sized opportunistic Guembelitria (63-150 m fraction) are
439
rare in the lower part of the section and slightly increase in the upper ~1.5 m below the
440
KTB (Fig. 9B). This group is not as rare in Elles but the relative abundance is <10% (Fig.
441
9A). The subsurface mixed layer dwellers (Table 1) constitute 80-90% of the CF1
442
assemblage at both Bidart and Elles. In the small size fraction, this group at Bidart with
443
30-40 species is almost twice as diverse as at Elles, (10-20 species). In the >150 m
444
fraction, the relative abundance of mixed layer dwellers constitute 60-80% at Bidart and
445
~80% at Elles (Fig. 9A, B). Their diversity fluctuates through CF1 and shows a gradual
446
decrease from ~20 to 14 species followed by a rapid decline ~3 cm below the KTB. The
447
thermocline dwelling globotruncanids are rare in the small size fraction at both Bidart
448
and Elles with relative abundance <5%. Two peaks of increased abundance and diversity
449
are noted in the upper 30-cm of late Maastrichtian at Bidart and in the last meter at Elles.
450
In the >150 m fraction, thermocline dwellers are more abundant (20-30%) at Bidart
451
compared to the same group in Elles (<10%). The sub-thermocline deep-water dwellers
452
in both size fractions at Bidart and Elles show consistently low abundances (<5%) and
453
diversity (<3 species) throughout zone CF1 (Fig. 9A, B).
454
At Gamsbach the mixed dwellers dominate in the larger size fraction, with ~70%
455
relative abundance, followed by the thermocline dwellers with 25-30% abundance. The
456
high abundance of thermocline dwelling globotruncanids in Gamsbach is more
457
comparable to the assemblage at Bidart than that at Elles (Fig. 9C). The deep dwellers are
458
represented by 1-3 species and account for <10% throughout CF1.
459
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The faunal assemblage differences between Elles vs. Bidart and Gamsbach appear
460
to be related to deposition in a relatively shallow continental shelf vs. deep middle
461
bathyal environments. This is indicated by the similarity in the faunal compositions
462
between Bidart, Gamsbach and the middle bathyal DSDP Site 525A, but dissimilarity
463
with Elles. For example (1) there is significantly higher relative abundance of P.
464
hariaensis at Bidart (20-25%), Gamsbach (10-20%) and Site 525A (10-20%), compared
465
with Elles (<5%) Abramovich and Keller, 2002; 2003); (2) Heterohelix globulosa is less
466
abundant in the >150 m fraction at Bidart (15-20%), Gamsbach (<10% with acme of
467
20%) than at Elles (40-50%); (3) Planoglobulina brazoensis is more abundant (5-10%) in
468
all three deeper sections, but rare at Elles; and (4) Globotruncana arca is more abundant
469
at Bidart (10%), Gamsbach (10-20%) and Site 525A (20-30%) than at Elles (<5%).
470
471
472
7. DISSOLUTION-BASED PROXIES FOR OCEAN ACIDIFICATION
473
474
7.1 Magnetic Susceptibility (MS)
475
Magnetic susceptibility (MS) of marine deposits depends essentially on their
476
mineralogical composition, and includes contributions (in proportion to their abundance)
477
from all - diamagnetic (e.g., calcite), paramagnetic (e.g., clay) and ferromagnetic (ex:
478
magnetite) - minerals present in the sediment. Since the pristine signal of magnetic
479
susceptibility in marine sediment reflects the balance between detrital input (high MS)
480
and carbonate productivity (low MS), it represents a robust paleoenvironmental indicator.
481
The rock magnetic properties for Elles are published in Stüben et al. (2003) and for
482
Bidart in Font et al. (2011, 2014) and Font and Abrajevitch (2014).
483
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Mass specific magnetic susceptibility values of the Maastrichtian marls from Elles
484
are in the range of 10-7 to 10-6 m3/kg, and comparable to other marine sediments
485
worldwide (Ellwood et al., 2008). The overall MS profile for zones CF3-CF1 shows a
486
positive correlation with percent phyllosilicates and an inverse correlation with carbonate
487
content, indicating a strong relationship between MS, climate (precipitation and runoff)
488
and/or sea-level rise. The KTB is featured by an abrupt shift in MS values, probably
489
resulting from an abrupt change in lithofacies (i.e. the clay layer). The Elles section does
490
not show the typical low MS interval below the KT boundary as in Bidart (Fig. 10A). At
491
Bidart the average MS value is 1.85x10-7 m3/kg for the lower part of zone CF1 and
492
~0.84x10-7 m3/kg for the final ~60 cm that forms the benchmark interval. The
493
characteristic abrupt increase in MS values (to 4.62x10-7 m3/kg, likely due to or a very
494
rapid change in sedimentation) marks the KTB hiatus at Bidart (Fig. 10B). For the
495
Gambsach section, the mass specific magnetic susceptibility of 49 samples excluding
496
turbiditic levels was measured (Fig. 10C). Maastrichtian MS values range between 10-8 to
497
10-7 m3/kg. The average MS value for the lower part of zone CF1 is 6.7x10-8m3/kg. About
498
~60 cm below the KTB, MS values reach a minimum of 4.5 x10-8m3/kg (average of
499
5.1x10-8m3/kg). These low MS values persist over an interval of 36-cm. Across the KTB
500
MS values show the typical increase culminating at 2.3x10-7m3/kg, similar to Bidart and
501
other KTB sections (e.g., Gubbio, Oman: Ellwood et al., 2003; Atlantic ODP 1259:
502
Erbacher et al., 2004; North Atlantic ODP 1049A: Moore et al., 1998).
503
504
7.2. Percent calcium carbonate
505
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Whole-rock percent CaCO3 content of marine sediments is the net result of the
506
local paleaoclimate, calcareous nannoplankton and calcareous dinoflagellate
507
palaeoproductivity, planktic foraminiferal abundance, water column pH/dissolution, pore-
508
water dissolution/re-crystallization and detrital influx. For localities with greater
509
terrigenous influx the ratio of Ca/detritus is a better estimate of biogenic CaCO3 as it
510
accounts for the detrital contribution. At Elles, the calcite/detritus ratio is low (0.5) near
511
the base of zone CF1 (8-10 m below the KTB, Fig. 10A). In this interval, the
512
phyllosilicate content is relatively high (25-35%) with an increasing trend. The
513
Ca/detritus ratio fluctuates but increases to 0.75-1.00 about 5-8 m below the KTB along
514
with an increase in MS values. The Ca/detritus ratio is higher in the top 4 m of zone CF1,
515
albeit with three sharp decreases.
516
The percent CaCO3 at Bidart gradually decreases from 55% at the base of zone
517
CF1 to about 40% 2 m below the KTB. An increase to 50% is observed 1.5 m below
518
KTB followed by values between 40-50% up to the low MS interval where values
519
sharply increase to 60% about 0.25 m below the KTB (Fig. 10B). At Gamsbach, CaCO3
520
ranges between 40-50% and records several abrupt decreases in zone CF1 correlative
521
with abrupt changes in percent quartz and MS values (Fig. 10C). In the 1m below the
522
KTB CaCO3 varies between 55-80% with the largest drop ~20 cm below the KTB
523
correlative with increased phyllosilicates and MS values. A drop in CaCO3 to nearly zero
524
percent is indicated in the KTB clay in all three sections.
525
526
7.3 Fragmentation index
527
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Tests of planktic as well as benthic foraminifera may undergo fragmentation due
528
to a multitude of taphonomic processes. Acidic ambient waters react with planktic
529
foraminiferal test carbonate, which leads to test dissolution and enhanced fragmentation.
530
The number of fragments have been used as a quantitative estimate of low pH in
531
foraminiferal assemblages (Thunell, 1976; Berger et al., 1982). The fragmentation index
532
may be calculated based on the following equation (Williams et al., 1985; Malmgren,
533
1987):
534
535
Fragment % = (Fragments/8)/[(Fragments/8) + whole tests]
536
537
538
Based on the assumption that each (non-crystallized) test breaks into an average of 8
539
fragments, the equation requires the total number of counted fragments to be divided by 8
540
to estimate the original number of whole tests. This is because the number of fragmented
541
tests is a better approximation of dissolution effects than the total number of fragments
542
counted (Le and Shackleton, 1992). As most of the planktic foraminifera at Bidart, Elles
543
and Gamsbach are recrystallized and/or infilled with secondary calcite, they are relatively
544
more resistant to fragmentation than pristine tests. We adjust for recrystallization by
545
reducing the number of fragments per test to 6 instead of 8 to avoid underestimation of
546
fragmented tests. Similarly, we consider 2 fragments per test for benthic foraminifera for
547
Bidart as they are far more resistant to fragmentation.
548
The stacked area graphs of Fig. 10 show fragmentation indices for Elles, Bidart
549
and Gamsbach. At Elles, fragmented tests increase from 23% to 46% at the beginning of
550
the low MS interval and increased fragmentation and imperfect tests are observed in the
551
top 4 m below the KTB (Fig. 10A). At Bidart, the percentage of imperfect tests for the
552
lower part of CF1 is considerably higher than at Elles and that of fragmented tests is
553
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lower (Fig. 10 A-B). In the uppermost ~60 cm of zone CF1, the low-MS interval is
554
accompanied by a significant (p<0.0001) increase in the combined abundance of
555
imperfect and fragmented from ~25% to 70% (~0.3 m below KTB, Fig. 10B). A drop in
556
this percentage ~2.5 m below the KTB boundary is followed by a rapid increase to 90%.
557
It must be noted that the fragmented tests (not the imperfect tests) dominate at the KTB
558
boundary hiatus. The fragmentation index at Gamsbach also shows an abrupt increase in
559
the combined abundance of fragmented and imperfect tests from 40% to ~90% at the
560
onset of the low MS interval ~50 cm below the KTB boundary. Within this interval,
561
fragmentation continues to be high (~90%) up to the KTB boundary (Fig. 10C). A brief
562
episode of decreased fragmentation is observed ~20 cm below the KTB boundary similar
563
to the event recorded at Bidart (supplementary material Tables 2-4). Figure 11 compares
564
the fragmentation indices of planktic and calcareous benthic foraminifera at Bidart. The
565
proportion of fragments of benthic foraminifera remains 2-3 % for most of the analyzed
566
CF1 interval (Plate 3: M, N, Q-S). This is consistent with the general robustness of
567
benthic morphologies and likely indicates a limited/uniform influence of sample
568
processing techniques and post-depositional breakage on the assemblage. However, an
569
increase in the fragments to ~5% concurrent with the Deccan benchmark event may
570
imply enhanced post-depositional bottom water transport during the climate-cooling
571
event (Figs. 10, 11). The imperfect benthic tests of CF1 largely show mechanical damage
572
unlike the planktic counterparts that show chemically leached surfaces and holes (Plate 2:
573
A-L). However, at the KTB and the lowermost Danian, sediments contain benthic
574
foraminifera that show intense leaching as well as mechanical damage strongly indicating
575
a dominance of post-depositional dissolution and bottom water transport affecting the
576
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assemblage. The planktic FI is not useful to isolate water column dissolution effects in
577
these samples (Fig 11).
578
579
7.4 Preferential preservation of robust morphologies
580
Dissolution preferentially decreases the relative abundance of thin-walled test
581
morphologies and therefore increases the relative abundance of robust dissolution-
582
resistant tests (e.g. Globotruncana, Globotruncanita, Pseudotextularia and P.
583
brazoensis), which may explain the increased calcite at this interval. This bias is evident
584
just below the KTB mass extinction in all three profiles analyzed. In the middle shelf
585
environment of Elles, globotruncanids and pseudotextularids are rare in zone CF1 (~1%),
586
increase to 2% in the low-MS interval and peaks at 10% and 6% in the 1 m preceding the
587
mass extinction (Fig. 10A). In the middle bathyal sections of Bidart and Gamsach, the
588
abundance of globotruncanids average 20-30% throughout zone CF1. At Bidart
589
globotruncanids abruptly reach 70% correlative with increased fragmentation and
590
decreased percent CaCO3 beginning about 30-cm below the KTB (Fig. 10B). At
591
Gamsbach, globotruncanids and another robust species (Planoglobulina brazoensis show
592
anomalously high abundance throughout zone CF1 with peak abundance of P. brazoensis
593
correlative with the high fragmentation index ~15 cm below the KTB (Fig. 10C).
594
595
8. DISCUSSION
596
597
8.1 Paleoclimate
598
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Whole-rock stable oxygen and carbon isotopes approximate mixed layer (mostly
599
calcareous nannoplankton) values in the deep-water sediments at Bidart and Gamsbach.
600
Diagenesis and recrystallization of tests may have overprinted 18O signals but their
601
effects on the 13C trend are limited. This claim is supported by the low correlation
602
coefficients of 13C vs. 18O (Stüben et al., 2003; R2=0.53 for Bidart and R2=0.27 for
603
Gamsbach) and the low correlation coefficient of 13C v/s %CaCO3 (R2=0.40 for Bidart
604
and R2=0.39 for Gamsbach).
605
The lower ~3 m of zone CF1 at Bidart and lower ~1.2 m of Gamsbach record
606
faunal responses comparable with those observed in the upper part of the late
607
Maastrichtian global warming at Elles and DSDP Site 525A. Globally, this warm event
608
began in zone CF2 as a likely consequence of the onset of the main phase-2 of Deccan
609
volcanism (Punekar et al., 2014a). This is consistent with the more negative 18O values
610
for these intervals indicating higher temperatures (-2‰ for Bidart and -1.5 to -2‰ for
611
Gamsbach, Fig. 8). At Bidart, the late Maastrichtian warm event is associated with
612
changes in the relative abundance of heterohelicids (particularly H. planata, H.
613
navarroensis and H. globulosa) and increased P. hariaensis abundance in the >150 m
614
fraction.
615
The end of the late Maastrichtian warming at Elles is marked by abrupt cooling
616
concurrent with unprecedented low MS values and increased test fragmentation ~4 m
617
below the KTB. This could be an expression of increased volcanic SO2 emission and
618
acidification (Fig. 10 this study; Fig. 5 of Stüben et al., 2003). At Bidart the onset of this
619
same cooling event is recognized by a drop in MS associated with increased dissolution
620
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28
and disappearance of Globigerinelloides yaucoensis, P. costulata, G. subcarinatus and R.
621
rugosa and at Gamsbach the disappearance of G. subcarinatus (Figs. 8, 10).
622
623
8.2 Paleoproductivity
624
Low 13C values with multiple negative excursions are observed at Elles and
625
Bidart (~0.5‰ and ~0.7‰ respectively) near the end of the late Maastrichtian CF1 warm
626
event Fig. 7). Similar negative excursions (~1 ‰) are recorded at deep marine Site 525A
627
(~0.3 ‰) as well as in shallow marine environments of Texas and India (Meghalaya) (Li
628
and Keller, 1998a; Gertsch et al., 2011; Abramovich et al., 2011). The existing dataset
629
shows a greater magnitude of 13C negative shift in the shallow sites (e.g., Meghalaya
630
(India), Mullinax-1 (Texas) and slightly deeper Elles (Tunisia). The smaller 13C shift at
631
deeper Site 525A may be due to an incomplete record resulting from erosion of early
632
Danian and topmost Maastrichtian sediments (Li and Keller, 1998a).
633
A rise in sea level near the end of the Maastrichtian and across the KTB transition
634
accompanied by increased precipitation and continental weathering/erosion (Haq et al.,
635
1988; Li et al., 1999) may have been responsible for the increased delivery of organic
636
carbon with very low 13C values into shallow marine environments. Low primary
637
productivity could have been the other important contributor to the low 13C values. Low
638
nannofossil productivity is recorded in Elles, Bidart, DSDP Site 525A, DSDP Site 577A
639
and DSDP Site 216 during the late Maastrichtian warm event in CF1 (Gorostidi and
640
Lamolda, 1995; Gardin, 2002; Tantawy et al., 2009; Thibault and Gardin, 2007, 2010).
641
Heterotrophic planktic foraminifera may have in turn suffered, resulting in a decrease in
642
carbonate export and the eventual 13C value of bulk carbonate. The lower carbonate
643
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29
(40%) between 1.2-3.0 m below the KTB at Bidart, correlative with the low 13C interval
644
lends support to this interpretation.
645
These global effects of increased precipitation and enhanced continental
646
weathering/erosion can be attributed to climate warming caused by ongoing large scale
647
Deccan volcanism. Additionally, the outgassing of higher-than-background quantities of
648
volcanic CO2 (13C about -5‰) would also significantly contribute to lowering the 13C
649
of the global oceans dissolved inorganic carbon (DIC), although the sediment record of
650
this signal would lag by ~1000 years (Zeebe, 2012).
651
652
8.3 Planktic Foraminifera
653
The high abundance of planktic and near absence of benthic foraminifera at Bidart
654
is consistent with the high planktic:benthic ratio (>90% planktics) reported by Coccioni
655
and Marsili (2007). The middle bathyal paleobathymetry and open marine setting appears
656
to be the reason for the unusual globotruncanid abundance at Bidart and Gamsbach (Fig.
657
9). This is supported by the high abundance of globotruncanids at Site 525A (~35-40%,
658
>150 m) where deposition occurred at ~1000 m depth (Shackleton and Boersma, 1985;
659
Abramovich and Keller, 2003). In relatively shallow (<150 m) continental shelf
660
environments, such as Elles, the diversity and abundance of globotruncanids is much
661
lower (Fig. 9).
662
663
8.4 Pre-KTB Ocean Acidification
664
White et al. (1994) showed that under present-day conditions (pH rain=5.6),
665
magnetite grains have very long time residence (>107 years) on land, but can be rapidly
666
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30
dissolved under more acidic conditions. In marine sediments, iron oxide dissolution by
667
ocean acidification has previously been documented for the Triassic-Jurassic mass
668
extinction and the coeval Central Atlantic Magmatic Province (Abrajevitch et al., 2013),
669
and more recently in the case of the KTB mass extinction at Bidart and Gubbio (Font et
670
al., 2014). The top ~50 cm interval of low-magnetic susceptibility (MS) that immediately
671
precedes the KTB at Bidart was attributed to the main phase-2 of Deccan volcanism
672
(Font et al., 2011, 2014; Font and Abrajevitch, 2014). The reductive iron oxide
673
dissolution may have occurred on land and/or in seawater. The first scenario was tested
674
by Font et al. (2014) who used a numerical weathering model to test for the consequences
675
of acidic rains on a continental regolith. Results revealed nearly complete magnetite
676
dissolution after ~31kyr (with a pH of 3.3.). However, the dissolution of magnetotactic
677
bacteria, which generally thrive the oxic-anoxic boundary in deep-sea marine sediments,
678
evokes ocean acidification as well and requires validation (Font and Abrajevitch, 2014;
679
Abrajevitch et al., in review).
680
Factors affecting the nature and concentration of detrital magnetic minerals and
681
therefore the MS of sediments include the nature and proximity of continental sediment
682
sources carbonate productivity, sea-level changes and/or post-depositional alteration
683
mechanisms (oxidation due to weathering/ diagenetic reduction of oxides). The influence
684
of sea level change on bulk MS is based on the relative contribution of carbonate
685
(diamagnetic, low MS) versus detrital input (paramagnetic clays and ferromagnetic iron
686
oxides, high MS) and thus can be estimated by correlating MS data with phyllosilicates
687
where a direct positive correlation implies a strong dependence of both parameters.
688
At Elles, the pre-KTB low-MS interval is not evident probably because
689
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31
paramagnetic minerals (clays) dominate the MS signal, supported by very low Ca:detritus
690
ratios (Fig. 10). At Bidart, the correlation between percent phyllosilicates and MS in zone
691
CF1 is poor (r=0.088), indicating an overall weaker influence due to sea level changes or
692
turbidity currents. For Gamsbach, this correlation is more complicated due to the
693
presence of frequent turbidite beds that are rich in dia/paramagnetic-silicates (Fig. 10;
694
samples Gb 5, 10, 12, 15, 27-28). However, the MS profile of Gambsach does show low
695
MS values for the ~40 cm interval below the KTB, similar to the MS profile of the Bidart
696
section (Fig. 10). The MS data of the present study thus suggest that the Gambsach
697
section is a good analog of the Bidart section. A prolonged period of acid rain on the
698
continents resulting in dissolution of magnetic detrital minerals can therefore be the
699
principal mechanism that caused the low MS intervals antecedent to the KTB because sea
700
level changes are a secondary influence on the MS profiles of Bidart and Gamsbach,
701
Surface ocean acidification in the low MS intervals of Bidart and Gamsbach is
702
indicated by increased dissolution and fragmentation of planktic foraminiferal tests (Fig.
703
10). This increased fragmentation interval correlates with the abrupt cooling event at
704
Elles (~4 m to ~0.5 m below the KTB), Bidart (~0.5 m interval below the KTB) and
705
Gamsbach (~0.4 m interval below the KTB) despite their different paleogeography,
706
paleobathymetry, depositional conditions and faunal assemblages, suggesting a common
707
cause. An increase in the proportion of dissolved tests (in addition to physically
708
fragmented ones) in Bidart and Gamsbach confirm the contribution of chemical leaching
709
as cause for imperfect carbonate tests with holes. This implies that water column
710
acidification is a likely cause for the observed increase in FI.
711
Benthic foraminifera are well preserved and even pristine looking in the same
712
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32
samples alongside leached and fragmented planktic foraminiferal tests at Bidart and
713
Gamsbach (Plates 2, 3). This may indicate acidification restricted to the upper water
714
column, or may reflect the inherently more robust mechanically resistant benthic tests.
715
Alegret et al. (2004; fig. 4) noted an increase in the proportion of agglutinated
716
foraminifera relative to calcareous benthic foraminifera in the top 10-cm of the
717
Maastrichtian at Bidart. We confirm these to be arenaceous (do not dissolve with 1:1
718
HCl), which may be interpreted as a consequence of ocean acidification or dissolution.
719
Only in the KTB red clay layer are benthic species corroded suggesting that low pH
720
acidic waters reached through the water column into deeper waters precisely at the KT
721
boundary event. However, the benthic species were little affected by the KTB mass
722
extinction or ocean acidification as their survival is globally documented (Widmark and
723
Malmgren, 1992; Alegret et al., 2001, 2003. Alegret and Thomas, 2004). For the most
724
part preceding the KTB, ocean acidification was restricted to the upper water column
725
with surface waters in equilibrium with very high atmospheric pCO2 and low CO32-
726
concentrations. Dissolution of test calcite during sinking through the water column would
727
make tests more fragile in post-depositional transport. The degree of
728
dissolution/fragmentation appears to be largely affected by local paleobathymetry and
729
species composition of the assemblage.
730
At Elles, the percentage of fragments is high throughout zone CF1 owing to
731
greater bottom water currents at shallower depths and also dominance of thin-walled
732
fragile heterohelicids in the assemblage (Fig. 10). In contrast, at Bidart the overall test
733
fragmentation is quite low due to a high proportion of structurally more robust
734
globotruncanids and quieter deposition at a greater depth. However, the proportion of
735
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33
leached out tests with holes due to dissolution increased (Fig. 10; red). At Gamsbach, the
736
degree of fragmentation is high throughout zone CF1 despite a high abundance of
737
globotruncanids. This can be attributed to the frequent turbiditic activity at this site that
738
may have increased post-depositional transport and fragmentation. The variable lithology
739
(deposition of quartz rich beds/lenses) may have facilitated pore-water dissolution,
740
recrystallization-cementation leading to lithified sediments that are difficult to
741
disaggregate and free individual tests.
742
743
8.5 Ocean acidification: The missing link to Deccan Volcanism?
744
The main phase-2 of Deccan volcanism occurred over ~750kyr entirely within
745
chron C29r, straddling the KTB (Schoene et al., 2014). However, all volcanism did not
746
occur at uniform intensity within this interval, as inferred from the multiple eruptive
747
events of geologically short duration separated by red/green boles indicating periods of
748
quiescence (Subbarao et al., 2000; Jay and Widdowson, 2008; Chenet et al., 2007, 2008).
749
Four of the Deccan phase-2 longest lava-flows across the Indian sub-continent (likely
750
signifying peak volcanic activity) erupted within a duration of ~250 kyr (zone CF2-CF1)
751
as seen in the Krishna-Godavari basin, India. The overlying Danian zone P1a sediments
752
constrain the age of the KTB mass extinction as coincident with the final mega-flow of
753
the peak phase-2 eruptions. (Keller et al., 2011a, 2012). The resultant cumulative loading
754
of 12,000–28,000 Gigatons (Gt) of CO2 into the end Cretaceous atmosphere within tens
755
of thousand years could increase the pCO2 on timescales that are recorded in the
756
sediments. This excess CO2 equilibrates with surface ocean water thus altering the
757
carbonate chemistry. The carbonic acid (H2CO3) formed dissociates to bicarbonate anion
758
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34
(HCO-3) and H+ ions, reducing the pH of surface waters. These H+ ions combine with
759
CO32- anions forming more HCO3- and decreasing the bioavailability of CO32- to
760
calcifying organisms to build their tests.
761
The episodic release of hundreds to thousands of teragrams of volcanogenic SO2
762
per year for each Deccan eruption would form sulfate aerosols upon reaction with
763
atmospheric water vapor and precipitate out as toxic acid rain locally/regionally
764
centennial timescales, shorter than the millennial timescales for removal of CO2 (Self et
765
al., 2008; Chenet et al., 2009; Mussard et al., 2014; Callegaro et al., 2014). This could
766
have been directly toxic/lethal for continental flora and fauna of affected areas. On land,
767
acid rain would exacerbate continental weathering. Sulfur dioxide would also lower the
768
surface ocean pH further, significantly contributing to the calcification crisis and high-
769
stress conditions for calcifying organisms on shorter timescales.
770
Ocean acidification has been identified as an important mechanism associated
771
with faunal turnovers and mass extinction events through geological history (e.g. ocean
772
anoxic events (OAEs) of the Paleozoic and the Paleocene-Eocene thermal maximum
773
(PETM)) that have affected marine calcifiers e.g. coccolithophores, planktic and benthic
774
foraminifera (review in Hönisch et al., 2012). Physiological manifestations of high-stress
775
due to acidification recorded as dwarfism, deformed tests, and R- strategist–dominated
776
assemblages in the Late Maastrichtian have already been linked with phase-2 Deccan
777
volcanism (Erba et al., 2010; review in Punekar et al., 2014a). Moy et al. (2009) reported
778
a 30%–35% lower calcification in modern Globigerina bulloides from the Southern
779
Ocean as compared to Holocene specimens. The anthropogenic CO2 emissions have
780
resulted in acidification of the Southern Ocean in the past ~300 yr (drop in pH by 0.1
781
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35
units, expected drop of 0.7 units in the next ~300 yr; Orr et al., 2005; Zeebe et al., 2008).
782
Carbonate tests of planktic organisms can experience dissolution in the water column as
783
demonstrated by the in vitro pteropod shell dissolution within 48 hours of exposure to
784
low pH waters (Fabry et al., 2008; Doney et al., 2009). The cumulative effect of thinner
785
walled tests undergoing water-column dissolution can render a test increasingly fragile
786
and vulnerable to fragmentation, consistent with our taphonomic evidence for ocean
787
acidification.
788
There are multiple lines of evidence in support of global surface ocean
789
acidification associated with the main phase-2 Deccan volcanism: (1) intensely corroded
790
carbonate tests and rapid extinctions of Maastrichtian planktic foraminifera in the
791
intertrappean sediments of the lava mega-flows in the Krishna-Godavari Basin of India
792
(Keller et al., 2011a, 2012). (2) Strong carbonate dissolution and high-stress
793
environments indicated by intense Guembelitria blooms (>95%) in CF1 of Meghalaya
794
(NE India) (Gertsch et al., 2011). And (3) evidence for iron oxide dissolution by
795
acidification inferred from low MS as well as for surface ocean acidification near the end
796
of zone CF1 preceding the KTB at distal sites such as Bidart (France) and Gamsbach
797
(Austria) as documented in this study.
798
799
9. Conclusions
800
Good temporal correlation between the age of the main phase of Deccan volcanic
801
eruptions in India and the age and episodic nature of climate fluctuations worldwide has
802
strengthened the case for large scale volcanism as a significant contributor to the Late
803
Maastrichtian biotic stress that culminated in the KTB mass extinction. However,
804
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36
inherent limitations due to incompleteness of the stratigraphic record and the lack of a
805
convincing kill-mechanism have inspired strong skepticism for this hypothesis.
806
A multi-proxy study of the Late Maastrichtian zone CF1 in Bidart (France) and
807
Gamsbach (Austria) reveals events in the final ~160 ky of the Late Maastrichtian that are
808
critical to understanding the role of Deccan volcanism in global high stress environments
809
and leads to the following conclusions.
810
The Late Maastrichtian warm event in the lower part of zone CF1 at Bidart
811
(France) and Gamsbach (Austria) is recognized by faunal responses similar to
812
those observed at Elles (Tunisia) and DSDP Site 525A.
813
A period of low 13C values and decreased percent CaCO3 content during the
814
global warming may record a combination of increased continental 12C influx
815
through increased runoff, suppressed primary and calcifier productivity and
816
equilibration of surface ocean waters with increased isotopically lighter
817
volcanogenic CO2.
818
An increase in carbonate dissolution and foraminiferal test fragmentation
819
suggests surface ocean acidification in 60-cm immediately preceding the KTB
820
at Bidart (France) and Gamsbach (Austria). This event globally correlates
821
with the low MS interval defined as the Deccan benchmark interval in Bidart
822
by Font et al., (2011, 2014).
823
The widespread ocean acidification interval is coincident with the rapid
824
cooling. At Elles, evidence for another rapid warming following this interval
825
coincides with the KTB mass extinction. The acidification may be the result
826
of equilibration with huge amounts of CO2 injected rapidly into the
827
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37
atmosphere at rates overwhelming the response time of feedback mechanisms.
828
829
Acknowledgements:
830
This research was supported by Princeton University’s Scott and Tuttle Funds, the U.S.
831
National Science Foundation (grants NSF EAR-0207407, EAR-0447171 and EAR-
832
1026271) and FCT (ref. PTDC/CTE-GIX/117298/2010). We thank the three anonymous
833
reviewers and the Guest Editor Prof. Wolfram M. Kürschner for their insightful
834
comments and suggestions.
835
836
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Stüben, D., Kramar, U., Berner, Z. A., Meudt, M., Keller, G., Abramovich, S., Adatte, T.,
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Hambach, U. and Stinnesbeck, W. 2003. Late Maastrichtian paleoclimatic and
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paleoceanographic changes inferred from Sr/Ca ratio and stable isotopes.
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10.1016/S0031-0182(03)00499-1
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Subbarao, K.V., Bodas, M.S., Khadri, S.F.R., Beane, J.E., Penrose Deccan 2000, Field
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excursion guide to the western Deccan Basalt Province. Penrose Field Guides, B.
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Summesberger, H., Wagreich, M. and Bryda, G. 2009. Upper Maastrichtian cephalopods
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and the correlation to calcareous nannoplankton and planktic foraminifera zones
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Museums Wien 111A, 159–182.
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Tantawy, A. A., 2003. Calcareous nannofossil biostratigraphy and paleoecology of the
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Tantawy, A. A., Keller, G. and Pardo, A., 2009. Late Maastrichtian volcanism in the
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LIST OF FIGURES
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Figure 1. Palaegeographic map of 66 Ma showing the study sections Bidart (France) and
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Gamsbach (Austria) and the reference section Elles (Tunisia, GSSP) relative to the
1376
location of the Reunion hotspot (focal point of Deccan volcanism). Modified after ©2000
1377
C R Scotese PALEOMAP Project.
1378
1379
Figure 2. (A) Lithological log of the upper Maastrichtian-basal Danian interval studied at
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Bidart, the red layer marks the KTB (B) Google Earth image showing the present day
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location of Bidart (C) Field photograph of the Bidart section showing the sampled
1382
interval and the position of KTB (in red).
1383
1384
Figure 3. (A) Lithological log of the uppermost Maastrichtian interval of Gamsbach, the
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red layer marks the KTB (B) Google Earth image showing the present day location of
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Gamsbach (C) Field photograph of the Gamsbach section showing the position of KTB
1387
(in red).
1388
1389
Figure 4. The completeness of Bidart and Gamsbach sections relative to Elles (Tunisia)
1390
based on planktic foraminiferal biozonation scheme of Keller et al. (1995; 2002). The
1391
biozone ages can be extrapolated using a KTB age of 65.5 Ma (Gradstein et al., 2004) or
1392
66.04 Ma (Renne et al., 2013). Hiatuses are observed at the KTB and at the P1a(1)/P1a(2)
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transition at Bidart. A major hiatus is identified at the KTB at Gamsbach due to missing
1394
zones P0, P1a(1) and early P1a(2).
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Figure 5. (A) Key foraminifera and geochemical attributes of the KTB boundary and
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lower Danian at Bidart (B) Abundance of late Maastrichtian planktic foraminifera of the
1398
63-150 μm size fraction and the KTB mass extinction. The δ13C record shows the
1399
characteristic ~2‰ negative shift at the KTB.
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Figure 6. (A) Faunal and geochemical changes at the KTB boundary and in the lower
1402
Danian of Gamsbach. (B) Late Maastrichtian planktic foraminifera of the >150 μm size
1403
fraction and the KTB mass extinction. The δ13C record shows ~1.3‰ negative shift at the
1404
KTB.
1405
1406
Figure 7. A comparison of the relative abundances of some key species (63-150 μm) in
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the upper Maastrichtian zone CF1 assemblage of (A) Elles (Tunisia) and (B) Bidart
1408
(France). Note that Heterohelix dentata, H. globulosa and Pseudoguembelina costulata
1409
dominate the assemblage in Elles, in contrast to Bidart where they are rare. Whole-rock
1410
13C and 18O are shown.
1411
1412
Figure 8. A comparison of the relative abundances of some key species (>150 μm) in the
1413
Late Maastrichtian zone CF1 assemblage of (A) Elles (Tunisia), (B) Bidart (France) and
1414
(C) Gamsbach (Austria). The deep-water assemblages of Bidart and Gamsbach are very
1415
similar to each other and different from the neritic assemblages of Elles. Planktic 13C
1416
and 18O for Elles are obtained from Rugoglobigerina rugosa and benthic values are
1417
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from Cibicidoides pseudoacuta. Whole-rock isotope data are shown for the Bidart and
1418
Gamsbach sections.
1419
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Figure 9. Relative abundances of depth-ranked groups of planktic foraminifera species in
1421
(A) Elles (Tunisia), (B) Bidart (France) and (C) Gamsbach (Austria). Thermocline
1422
dwelling globotruncanids (blue) are more abundant in the >150 μm fraction of Bidart and
1423
Gamsbach (open marine settings) as compared to Elles (neritic setting). Poor preservation
1424
of foraminifera in the 63-150 μm fraction of the Gamsbach section precluded quantitative
1425
analysis.
1426
1427
Figure 10. Multi-proxy data shows a dissolution interval immediately preceding the
1428
KTB. A low magnetic susceptibility (MS) interval in the upper part of zone CF1 of Elles,
1429
Bidart and Gamsbach (yellow) marks a regional chemical benchmark of Deccan
1430
volcanism (after Font et al, 2011; 2014). Increased planktic foraminiferal test
1431
fragmentation in the low MS interval supports water column carbonate dissolution.
1432
1433
Figure 11. Magnetic susceptibility (MS) data for Bidart (Font et al., 2011) along with the
1434
fragmentation index (FI) data for planktic and benthic foraminifera. The geochemical
1435
Deccan benchmark interval coincides with a pronounced water column dissolution event
1436
recorded by the planktic foraminifera. The benthic FI for the same interval indicate only a
1437
minor contribution of post-depositional breakage.
1438
1439
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Plate 1. Characteristic taxa of the upper Maastrichtian zone CF1 assemblage at Bidart,
1441
France, scale bar= 100 m.
1442
A. Globotruncanita stuarti (de Lapparent), spiral view
1443
B. Abathomphalus mayaroensis (Brönnimann), spiral view
1444
C. Heterohelix rajagopalani Govindan
1445
D. Pseudoguembelina hariaensis Nederbragt
1446
E. Heterohelix globulosa (Ehrenberg)
1447
F. Pseudotextularia elegans (Rzehak)
1448
G. Racemiguembelina fructicosa (Egger)
1449
H. Planoglobulina brazoensis (Martin)
1450
I. Guembelitria cretacea (Cushman)
1451
J. Spiroplecta americana (Ehrenberg)
1452
K. Hartella harti Georgescu & Abramovich
1453
L. Heterohelix navarroensis (Loeblich)
1454
M. Plummerita aff. hantkeninoides (Brönnimann)
1455
N. Rugoglobigerina macrocephala (Brönnimann)
1456
O. Globigerinelloides volutus (White)
1457
P. Globigerinelloides subcarinatus (Brönnimann)
1458
1459
Plate 2. Planktic foraminifera indicating varying degrees of preservation in the upper
1460
Maastrichtian zone CF1 assemblage at Bidart, France, scale bar= 100 m.
1461
(A-L): “Imperfect” tests with minor breakages and/or holes and signs test surface
1462
dissolution
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(M-S): “Fragments” defined by less than two-thirds or the original test preserved.
1464
1465
Plate 3. Benthic foraminifera indicating varying degrees of preservation in the upper
1466
Maastrichtian zone CF1 and the KTB assemblage at Bidart, France, scale bar= 100 m.
1467
(A-G): “Perfect” tests with no signs of chemical or mechanical damage
1468
(H-L, O, P): “Imperfect” tests with minor breakages and/or holes and signs test surface
1469
dissolution. Note that the proportion of the tests with holes is maximum at the KTB and
1470
in the early Danian sediments.
1471
(M, N, Q-S): “Fragments” defined by less than two-thirds or the original test preserved.
1472
1473
1474
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Figure 1
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Figure 2
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Figure 3
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Figure 4
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Figure 5A
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Figure 5B
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Figure 6A
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Figure 6B
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Figure 7
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Figure 8
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Figure 9
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Figure 10
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Figure 11
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Plate 1
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Plate 2
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Plate 3
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1525
Highlights
1526
Plummerita hantkeninoides zone CF1 present in Bidart (France), Gamsbach
1527
(Austria)
1528
Low MS Deccan benchmark coincides with a carbonate dissolution event at both
1529
sites
1530
Dissolution event immediately precedes the KTB mass extinction
1531
Likely ocean acidification event may have resulted in carbonate crisis
1532
Deccan volcanic main phase-2 may be the likely cause
1533
