A critical evaluation of carbon isotope stratigraphy ...

Earth-Science Reviews 126 (2013) 116–146

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A critical evaluation of carbon isotope stratigraphy and biostratigraphic implications for Late Cretaceous global correlation Ines Wendler ⁎ Department of Geological Sciences, University of Bremen, P.O. Box 330 440, D-28334 Bremen, Germany

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Article history: Received 24 January 2013 Accepted 6 August 2013 Available online 19 August 2013 Keywords: Carbon isotope Global correlation Biostratigraphy Magnetostratigraphy Late Cretaceous Planktic foraminifera

a b s t r a c t Climate variability is driven by a complex interplay of global-scale processes and our understanding of them depends on sufficient temporal resolution of the geologic records and their precise inter-regional correlation, which in most cases cannot be obtained with biostratigraphic methods alone. Chemostratigraphic correlation based on bulk sediment carbon isotopes is increasingly used to facilitate high-resolution correlation over large distances, but complications arise from a multitude of possible influences from local differences in biological, diagenetic and physico-chemical factors on individual δ13C records that can mask the global signal. To better assess the global versus local contribution in a δ13C record it is necessary to compare numerous isotopic records on a global scale. As a contribution to this objective, this paper reviews bulk sediment δ13Ccarb records from the Late Cretaceous in order to identify differences and similarities in secular δ13C trends that help establish a global reference δ13C record for this period. The study presents a global-scale comparison of twenty δ13C records from sections representing various palaeo-latitudes in both hemispheres and different oceanic settings from the Boreal, Tethys, Western Interior, Indian Ocean and Pacific Ocean, and with various diagenetic overprinting. The isotopic patterns are correlated based on independent dating with biostratigraphic and paleomagnetic data and reveal good agreement of the major isotope events despite offsets in absolute δ13C values and variation in amplitude between the sites. These differences reflect the varying local influences e.g. from depositional settings, bottom water age and diagenetic history, whereas the concordant patterns in δ13C shifts might represent δ13C fluctuations in the global seawater dissolved inorganic carbon. The latter is modulated by variations in organic matter burial relative to re-mineralization, in the global-scale formation of authigenic carbonate, and in partitioning of carbon between organic carbon and carbonate sinks. These variations are mainly controlled by changes in climate and eustasy. Additionally, some globally synchronous shifts in the bulk δ13Ccarb records could result from parallel variation in the contribution of authigenic carbonate to the sediment. Formation of these cements through biologically mediated early diagenetic processes is related to availability of oxygen and organic material and, thus, can be globally synchronized by fluctuations in eustasy, atmospheric and oceanic oxygen levels or in large-scale oceanic circulation. Because the influence of early diagenetic cements on the bulk δ13Ccarb signal can, but need not be synchronized, chemostratigraphy should not be used as a stand-alone method for trans-continental correlation, and especially minor isotopic shifts have to be interpreted with utmost care. Nevertheless, the observed consistency of the δ13C correlations confirms global scale applicability of bulk sediment δ13C chemostratigraphy for the Late Cretaceous, including sediments that underwent lithification and burial diagenesis such as the sediments from the Himalayan and Alpine sections. Limitations arise from increased uncertainties (1) in sediments with very low carbonate content, (2) from larger δ13C variability in sediments from very shallow marine environments, (3) from unrecognized hiatuses or strong changes in sedimentation rates, and (4) in sections with short stratigraphic coverage or with few biostratigraphic marker horizons. The combination of chemostratigraphy with biostratigraphy and magnetostratigraphy substantially increases the precision and temporal resolution of inter-regional correlations and helps overcome problems that arise from differences in biostratigraphic schemes, facies or provincialism of key fossils. By using an iterative approach to stepwise increase precision of the correlations, isochroneity of first and last occurrences of marker species versus chemostratigraphy is tested, which helps to improve biostratigraphic zonations, to assess zonal boundary ages and to identify useful criteria for defining Late Cretaceous stage boundaries, many of which are still not formally defined. The presented correlations indicate a consistent position for most planktic foraminifer zonal boundaries relative to corresponding isotope shifts during the mid-Cretaceous sea-level high, whereas

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diachroneity appears to be more pronounced during the Late Campanian and Maastrichtian global sea-level fall. A similar pattern is observed for trans-continental consistency in the δ13C shifts. Graphic correlation of isotopic shifts, magnetostratigraphic and biostratigraphic events among the compared sections is used to detect hiatuses or relative changes sediment accumulation rates and visualizes consistency or offsets of individual biostratigraphic markers relative to chemo- and magnetostratigraphy. Finally, an attempt of a global average δ13C stack is presented for the Turonian through Maastrichtian. © 2013 Elsevier B.V. All rights reserved.

Contents 1.

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.1. Chemostratigraphic application of δ13C . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2. The δ13C record: bulk versus component . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.1. Organic carbon . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.2. Single-species fossils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.3. Radiaxial and radial-fibrous carbonate cements . . . . . . . . . . . . . . . . . . 1.2.4. Matrix micrite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.2.5. Bulk sediment carbonate . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.3. Reading the bulk carbonate δ13C record . . . . . . . . . . . . . . . . . . . . . . . . . . 1.3.1. Factors influencing bulk sediment δ13Ccarb . . . . . . . . . . . . . . . . . . . . 1.3.2. The δ13C of dissolved inorganic carbon (DIC) . . . . . . . . . . . . . . . . . . . 1.3.3. Diagenetic alteration . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.3.4. Biological factors: faunal composition and vital effects . . . . . . . . . . . . . . . 1.4. Applicability of δ13C for global correlation and causes of global δ13C excursions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1.5. Cretaceous δ13C data and objectives 2. Material and methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1. Compared sections and data sources . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2. Age models and methods of correlation . . . . . . . . . . . . . . . . . . . . . . . . . . 3. Results and discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.1. Biostratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2. Changes in sediment accumulation rates and hiatuses . . . . . . . . . . . . . . . . . . . 3.3. Regional correlation for Tibet . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.4. Global δ13C comparison: Turonian . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.5. Global δ13C comparison: Coniacian through Santonian . . . . . . . . . . . . . . . . . . . 3.6. Global δ13C comparison: Campanian through Maastrichtian . . . . . . . . . . . . . . . . . 3.7. Late Cretaceous global δ13C correlation . . . . . . . . . . . . . . . . . . . . . . . . . . 3.8. Relation of δ13C to magnetostratigraphy and biostratigraphy . . . . . . . . . . . . . . . . 3.9. Testing isochroneity of biostratigraphic index taxa: planktic foraminifera . . . . . . . . 3.10. Testing isochroneity of biostratigraphic index taxa: nannofossils . . . . . . . . . . . . . . 3.11. Testing isochroneity of biostratigraphic index taxa: macrofossils . . . . . . . . . . . . . . 3.12. Towards a global Late Cretaceous δ13C stack . . . . . . . . . . . . . . . . . . . . . . . . 3.13. Patterns in sediment accumulation rates . . . . . . . . . . . . . . . . . . . . . . . . . 4. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Appendix A. List and explanation for abbreviations used in this paper . . . . . . . . . . . . . . . . Appendix B. List and explanation for abbreviations of microfossil species names and biostratigraphic zones Appendix C. List and explanation for abbreviations of macrofossil species names and biostratigraphic zones References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1. Introduction 1.1. Chemostratigraphic application of δ13C Climate variability in Earth history is driven by a complex interplay of numerous factors such as astronomic, atmospheric, oceanic and tectonic processes. Because many of these processes act on a global scale, our understanding of these interactions strongly depends on sufficient temporal resolution of the geological records and our ability to correlate them globally. While biostratigraphy represents an essential tool for dating sedimentary sequences, its use for detailed correlation on inter-regional or global scales is often limited due to provincialism of key species, diachronous first and last occurrences, lack of marker species related to preservational effects or the use of different biostratigraphic schemes. Furthermore, the biostratigraphic zones are often too long to allow for the temporal resolution required to reconstruct

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variations of the relevant processes. The combination of biostratigraphy and carbon isotope chemostratigraphy has proven helpful to increase the temporal resolution and precision of correlations (Scholle and Arthur, 1980; Berger and Vincent, 1986; Shackleton, 1986; Weissert, 1989; Jenkyns et al., 1994; Jarvis et al., 2006 and many others). It further allows correlation between sites that are dated with different fossil groups or that reveal diachronous occurrence of species. 1.2. The δ13C record: bulk versus component Depending on the material that is available and the questions to be addressed, researchers have used carbon isotope data from organic carbon, bulk sediment carbonate or from various carbonate components such as the calcareous tests of benthic or planktic organisms, radialfibrous carbonate cements or isolated inter-component matrix micrite. All of these approaches have their advantages and limitations.

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1.2.1. Organic carbon Carbon isotope records may be obtained from organic matter if the sediments contain sufficient organic material. Some δ13C records from terrestrial organic matter were shown to correlate with δ13C of marine organic matter (e.g. Heimhofer et al., 2003), allowing for land–sea correlation. In some cases, the δ13Corg records show similar trends as δ13Ccarb records obtained from the same material (e.g. Tsikos et al., 2004), but often there are differences in the magnitude or exact timing of isotopic shifts. Similarity or differences among δ13Ccarb and δ13Corg records appear to be related to early diagenetic processes that involve the formation of authigenic carbonate (Schrag et al., 2013). Because organic matter is much more reactive than carbonate and a large part of the exported organic material is remineralized in the upper part of the sediment column, the processes influencing the δ13Corg are complex and much less understood than for carbonate (Werne and Hollander, 2004). 1.2.2. Single-species fossils The study of isolated single-species fossils allows reconstructions of surface to deep ocean gradients, but taxonomic knowledge is required, and the interpretation is complicated by taxon-specific, biologically controlled disequilibrium effects (e.g. Birch et al., 2013; Wendler et al., 2013) and the possibility of reflecting seasonal signals (e.g. King and Howard, 2005; Jonkers et al., 2013). Additionally, infaunal benthic species record the δ13C signature of the pore water rather than that of the bottom water (Belanger et al., 1981; Rathburn et al., 1996; McCorkle et al., 1997; Mackensen et al., 2000; Schmiedl et al., 2004; Basak et al., 2009). Because in most pelagic sediments benthic taxa are relatively rare, this method is very time-consuming, which limits stratigraphic resolution with a given time or financial budget. Furthermore, the limited stratigraphic range of most species makes single-species δc13C records over long geologic periods impossible. Preservation of calcareous fossils is good to excellent in many Mesozoic and Cenozoic sediments, but older sequences are usually too lithified for separation of single species or lack macroscopic skeletal components (Precambrian carbonates). A common practice for Paleozoic sediments is analysis of δ13C from marine radiaxial and radial-fibrous carbonate cements or from matrix micrite. 1.2.3. Radiaxial and radial-fibrous carbonate cements The advantages of studying cements are that (1) they are an in situ precipitate and (2) they are abiotic so that their δ13C is not affected by vital effects. However, the reliability of radiaxial and radial-fibrous cements as recorder of ambient seawater δ13CDIC cannot be tested because these cement types are unknown from modern marine carbonate settings (Immenhauser et al., 2008). As these cements most likely represent an in situ alteration product it cannot be excluded that they have experienced a significant geochemical resetting during their diagenetic transformation (Wilson and Dickson, 1996). 1.2.4. Matrix micrite Micrite that is found in mudstone or as matrix between larger components is derived from a microcrystalline carbonate ooze consisting of small detrital carbonate particles that can be autochthonous or transported, and/or of automicrite that represents an in situ organomineralic carbonate precipitate. The latter is common in Paleozoic to modern neritic carbonates, whereas pelagic micritic carbonates from the Mesozoic and Cenozoic are largely composed of calcareous nannofossils. The main problem with analyzing bulk matrix micrites from neritic environments is syndepositional diagenesis driven by organic matter degradation (Patterson and Walter, 1994b; Walter et al., 2007). The large surface to volume ratio of these small carbonate particles and the high contribution of metastable carbonate phases in shallow marine sediments make them especially prone to early marine diagenetic alteration. Nevertheless, δ13C data from matrix micrite or from marine radaxial cements may still represent

approximations of the first-order trends in secular seawater geochemistry as documented by the global reproducibility of some of these δ13C records that cannot be explained by diagenetic processes (Immenhauser et al., 2008). 1.2.5. Bulk sediment carbonate Generation of δ13Ccarb data from bulk sediments does not require taxonomic knowledge and does not depend on well preserved fossils that can be extracted from the sediments. Furthermore, the method is comparably easy and fast, which is a big advantage for producing long δ13C records at high resolution. This explains the wide application of bulk δ13C studies to compare and correlate stratigraphic sections (e.g. Scholle and Arthur, 1980; Ripperdan et al., 1992; Saltzman et al., 2000; Herrle et al., 2004; Payne et al., 2004). Comparisons of bulk sediment δ13C data with those obtained from single species from the same successions usually show consistent short and longterm trends with less scatter in the bulk sediment data (possibly due to the above mentioned early diagenetic equilibration of micritic components). A systematic offset to more negative values due to contribution of early diagenetic authigenic calcite (Jeans et al., 2012; Schrag et al., 2013) and burial diagenesis is typically observed in the bulk sediment δ18O and δ13C curves as compared to single species measurements, but the δ13C curves often show the same trends (e.g. Schönfeld et al., 1991; Niebuhr and Joachimski, 2002; Voigt and Gale, 2002). However, the δ13C values measured from bulk carbonate represent the average of the individual calcareous components, including fossils and cements, and as such can be influenced by numerous factors, similar to the other methods described in this section. Therefore, in order to understand and correctly interpret bulk δ13C data, the potential contribution from each of these processes has to be critically assessed for each case. 1.3. Reading the bulk carbonate δ13C record 1.3.1. Factors influencing bulk sediment δ13Ccarb The carbon isotopic signal measured from bulk sediments is the net result of a highly complex interplay of sedimentological, physicochemical and biological processes that have affected the rock record (Fig. 1). The major influencing components are (1) the δ13CDIC of the ambient water, (2) the type of carbonate grains and taxonomic composition of calcareous shells with their specific habitat and vital effects, and (3) diagenetic alterations (syn-depositional and burial). 1.3.2. The δ13C of dissolved inorganic carbon (DIC) Variations in δ13CDIC in the oceans over geologic time scales mainly revolve around changes in the distribution of carbon among the Earth's surface carbon reservoirs and in the size and rate of fluxes of organic carbon relative to carbonate stored in the lithosphere. While precipitation of carbonate involves little carbon isotope fractionation, and the δ13Ccarb is relatively insensitive to changes in temperatures, the formation of new organic matter through photosynthesis occurs under strong preferential uptake of 12C relative to 13C. This negative fractionation causes low δ13C values of organic compounds relative to atmospheric CO2 with values around − 26‰ (− 9 to − 33‰) for land plants and around −22‰ (−10 to −32‰) for marine phytoplankton, while δ13CDIC values of modern oceanic surface waters are ~2.0‰ (Sarmiento and Gruber, 2006), although these values may have been different in the Cretaceous (Dean et al., 1986). Because photosynthesis in the ocean is restricted to the photic zone the preferential uptake of 12 C during photosynthesis leads to enrichment of 13C in the dissolved carbon of surface waters. If the sites of organic matter production and remineralization are separated, such as in the open ocean and in some stratified shallow seas, a distinct vertical gradient in δ13CDIC exists with higher values in the surface waters and lower values in bottom waters where 12C is released together with nutrients as organic matter is degraded.

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Fig. 1. Processes influencing bulk sediment δ13C values. Scheme visualizing the complex interplay of biological, sedimentological and physico-chemical factors that influence the δ13C values measured on bulk sediment carbonate. Shallow epeiric seas experience much more variability in most of these factors than the open ocean. Local differences in one or more of these processes can lead to regional variations in absolute bulk δ13C values and in the magnitude of bulk δ13C shifts, while global fluctuations in oceanic δ13CDIC can result in bulk δ13C shifts that can be recognized and correlated on global scale. Many of these factors are influenced and synchronized by changes in eustatic sea-level and in the globally integrated weathering flux, or by major perturbations of the global carbon cycle (e.g. by large igneous provinces, OAEs, asteroid impacts).

Strong horizontal gradients in δ13CDIC are characteristic for shallow epeiric seas and carbonate platforms (Patterson and Walter, 1994a; Immenhauser et al., 2008; Gischler et al., 2009). This horizontal variability is caused by the comparably small reservoir size in these settings and the more or less restricted exchange with the deep ocean. Therefore, compared to the deep ocean, shallow epeiric seas experience much stronger coastal influences such as influx of freshwater and dissolved chemical solutes (especially DIC from terrestrial organic matter), evaporation, as well as influences from fluctuations in sea-level (potentially leading to transient exposure), nutrient upwelling, primary productivity, temperature and bottom water ventilation (Immenhauser et al., 2008). For example, even though high rates of freshwater influx and evaporation may be balanced and result in normal marine salinities in the basin, these high flux rates clearly complicate the δ13CDIC budget (Patterson and Walter, 1994a). During much of the Cretaceous global sea-level was relatively high, topography was generally low (especially around the Tethys and the Atlantic Ocean) and large shelf areas were inundated (Skelton et al., 2003). Many of the Cretaceous sections that are used as paleoceanographic archives were deposited on these shelf areas (e.g. sections in the Tethys Himalaya), on carbonate platforms (e.g. Levant platform) or in epicontinental seas (e.g. Western Interior or Boreal Chalk sections). The strong spatial and temporal variability that is typical for shallow epeiric seas with b100 m water depth does not apply to shelf areas and epicontinental seas with several hundreds of meters water depth (Immenhauser et al., 2008). However, it should be kept in mind that during periods of lower sea-level some of these areas may have experienced influences comparable to those of the shallow epeiric seas. Spatial δ13CDIC differences in the open ocean are much less severe. They include a gradient in δ13CDIC in bottom waters as a function of residence time. This aging effect leads to lower δ13CDIC values in older as compared to younger bottom waters due to the continued release of 12C from remineralized organic matter at the seafloor. There is also a latitudinal gradient in the surface waters of modern oceans with significantly lower δ13CDIC values at high latitudes than at temperate latitudes (Goericke and Fry, 1994). This trend is opposite to what would be expected from the temperature distribution (colder waters

generally have higher δ13CDIC values than warmer waters) and appears to be related to differences in deviation from CO2 equilibrium between the atmosphere and the mixed layer and to seasonality in phytoplankton productivity (Goericke and Fry, 1994). Secular variations in the global oceanic δ13CDIC occur as a result of changing proportions of organic to inorganic (carbonate) carbon that is buried in the sediments and is thus removed from the surface reservoirs (ocean–atmosphere–biosphere). A global net deposition of organic matter causes an increase in δ13CDIC in the whole ocean, while global net oxidation of organic matter decreases the δ13CDIC in the whole ocean. It was recently proposed that early diagenetic formation of authigenic carbonates represents a major sink in the global carbon cycle with the potential to influence oceanic δ13CDIC (Schrag et al., 2013). The changes in global ocean δ13CDIC are reflected e.g. in the calcitic shells of microfossils and can cause synchronous shifts in the bulk carbonate δ13C that allow for high-resolution global correlation. From the multitude of factors that variably influence the δ13CDIC at different locations (Fig. 1) it is clear that the absolute δ13C values as well as the magnitude of these shifts can be expected to vary among different regions. Difficulties with correlation of δ13C records arise when local influences are overriding the global signal or when diagenetic overprinting has erased the original trends. 1.3.3. Diagenetic alteration Syndepositional dissolution and reprecipitation of carbonate are significant processes that affect all calcareous marine sediments more or less strongly (Patterson and Walter, 1994b; Minoletti et al., 2005; Sexton and Wilson, 2009; Jeans et al., 2012; Schrag et al., 2013). There is not necessarily a relation to undersaturated bottom waters because syndepositional dissolution of CaCO3 can be driven by organic matter degradation in the sediments. These early diagenetic processes are particularly important in shallow epeiric seas because of the higher abundance of metastable carbonates (high-Mg calcite, aragonite) in these sediments (Patterson and Walter, 1994b). While in some cases, the δ13C signal from shallowwater sections can still be correlated to pelagic sections (Jenkyns, 1995; Huck et al., 2011) even if they experienced repeated subaerial

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exposure and significant impact of diagenesis (Grötsch et al., 1998), other studies have demonstrated significant differences in δ13C shifts between shallow and deep-water sites (e.g. Colombié et al., 2010). Shallow water sequences can further be affected by meteoric syndepositional diagenesis, leading to pronounced negative shifts in δ13C especially during icehouse periods with large glacio-eustatic sealevel changes (Immenhauser et al., 2008). Early diagenetic alteration of the δ13C signal can also be problematic in oceanic porous sediments with relatively high TOC contents that are deposited under oxygenated bottom waters (Wendler et al., 2002). The decrease in pore water pH caused by organic matter oxidation is buffered by carbonate dissolution leading to chemically evolved pore waters that adopt a mixture of the isotopic signal of the most abundant carbonate grains and that of the degraded organic matter. Formation of syn-depositional cements is favored in pore waters with elevated alkalinity resulting from microbial reduction of sulfate and iron under reduced oxygen availability. These authigenic carbonates can cause considerable shifts in the bulk sediment δ13Ccarb towards lower values (Jeans et al., 2012; Schrag et al., 2013). It appears that cements formed under sub-oxic or oxic conditions affect the bulk δ13Ccarb signal much less than anoxic cements that can substantially lower the bulk isotopic values (Jeans et al., 2012). With reprecipitation of carbonate the compositions of all carbonate grains will shift towards equilibrium with respect to pore water (Patterson and Walter, 1994b), thereby reducing much of the δ13Ccarb variability within the sediment. Although the carbon isotopic signature then does not solely reflect the δ13CDIC of seawater, this process has some advantage because it reduces the intra-sample scatter and uncertainty for bulk carbon isotope measurements, leading to better reproducibility of the results. However, large regional differences in the contribution of authigenic carbonate to the sediment can produce isotopic signatures that are not comparable among different sites and, thus, should not be correlated. Late diagenetic processes include cementation of pore space, recrystallization of shells in fossils and pressure solution. In rock-dominated, closed to semi-closed diagenetic systems (without or with just minor influence by external meteoric fluids) the primary trends in the δ13C record are often preserved and can be correlated to distant areas despite strong lithification and cementation. Such examples for Upper Cretaceous sections were described from the Tethys Himalaya (Wendler et al., 2009; Wendler et al., 2011b) or from the limestone of the Italian sections (Jenkyns et al., 1994). It should be noted though, that pressure-related re-distribution of carbonate primarily affects smaller particles that, in Meso- and Cenozoic pelagic carbonates, mainly consist of coccoliths. As the latter typically have lower δ13C values than most other co-occurring fossils a distinct isotopic pattern is often observed with higher δ13C values in the (coccolith-deprived) marls than in the limestone beds (Frank et al., 1999; Jeans et al., 2012). While pressure dissolution usually does not obliterate the main isotopic trends in the sequence, strong meteoric influence can obscure the original isotopic trends, making long-distance correlation impossible. These processes typically lead to decreased δ13C values from meteoric waters containing DIC from oxidized organic matter, and to a concurrent decrease in δ18O. Correlation between oxygen and carbon isotopes is, therefore, commonly observed in material that was strongly altered by burial diagenesis. However, covariance between oxygen and carbon isotopes cannot be used for unequivocal identification of meteoric late diagenesis because this covariance is also observed in modern calcareous skeletons and in pristine fossils where it reflects simultaneous depletion in δ13C and δ18O due to (1) kinetic fractionation effects that are related to growth rate (McConnaughey, 1989; Wendler et al., 2013) or (2) vital effects that are related to pH (Spero et al., 1997; Bijma et al., 1999; Zeebe, 1999; Adkins et al., 2003; Zeebe, 2007). 1.3.4. Biological factors: faunal composition and vital effects Vital effects refer to biologically controlled deviations from isotopic equilibrium between the secreted fossil shell material and the ambient

seawater that are related to the influence of physiological fractionation processes on biomineralization, as was demonstrated in numerous laboratory and field experiments (Erez, 1978; Duplessy et al., 1981; Grossman, 1984; Dudley et al., 1986; Spero et al., 1991; Dudley and Nelson, 1994; Spero and Lea, 1996; Bijma et al., 1998; Cooke and Rohling, 1999; Bemis et al., 2000; Erez, 2003; Bauch et al., 2004; Chang et al., 2004; Bentov and Erez, 2006; McCorkle et al., 2008; Bentov et al., 2009; Bernhard et al., 2010; Dueñas-Bohórquez et al., 2011; Kisakürek et al., 2011; Ziveri et al., 2012; Birch et al., 2013). For the Cretaceous, species-specific offsets have been demonstrated for co-occurring planktic foraminifera (e.g. Corfield et al., 1990; D'Hondt and Arthur, 1995; Houston et al., 1999; Abramovich et al., 2003) as well as for benthic foraminifera (Friedrich et al., 2006; Wendler et al., 2013). Because the amount of this isotopic disequilibrium depends upon the calcifying organisms (McConnaughey, 1989; Wefer and Berger, 1991; Norris, 1998), large changes in the faunal composition can cause shifts in the bulk δ13Ccarb record. Additionally, changes in the proportions of planktic to benthic species as well as of epifaunal to infaunal species may contribute to changes in bulk sediment δ13Ccarb as these organisms form their shells in different habitats along the vertical gradient in δ13CDIC between the surface and deep ocean and within the upper centimeters of the sediment (Belanger et al., 1981; Rathburn et al., 1996; McCorkle et al., 1997; Mackensen et al., 2000; Schmiedl et al., 2004; Mackensen, 2008). This potential influence on bulk δ13Ccarb is minimized in most Mesozoic and Cenozoic pelagic sediments where the contribution from planktic organisms (especially nannofossils) is by far overwhelming that from benthic organisms. However, large differences in vital effects have been found for different coccolith taxa (Dudley et al., 1986; Dudley and Nelson, 1994; Ziveri et al., 2003; Birch et al., 2013). A major taxonomic turnover in the nannoplankton could, therefore, contribute significantly to shifts in the bulk δ13Ccarb record, as e.g. described for the end of the Cretaceous (Jiang et al., 2010; Alegret et al., 2012) or for the Paleocene/Eocene Thermal Maximum (Bralower, 2002). Similarly, changes in the relative contribution of nannofossils and of other carbonate microparticles (mostly of diagenetic origin) can cause shifts in the bulk δ13Ccarb signal (Minoletti et al., 2005). Trophic recourses determine the composition of an ecosystem, and the relative abundance of calcifying and non-calcifying organisms as well as the relative contribution of photozoan and heterozoan carbonate production have an important influence on the δ13C signal (Föllmi et al., 2006). In high-productive environments where dissolved oxygen and sedimentary carbonate contents are low, additional effects from carbonate-ion undersaturation and incorporation of 13C depleted methanotrophic biomass can significantly influence the δ13C in calcareous microfossil tests (Mackensen, 2008). Other biological factors that may contribute to shifts in the δ13Corg include compositional variations of terrestrial and marine organic carbon in the sediments and the biological reworking of organic material through zooplankton gut communities (Hayes, 1993).

1.4. Applicability of δ13C for global correlation and causes of global δ13C excursions From the large number of potential local influences (biological, diagenetic, spatial gradients in DIC) on the δ13C signal of bulk sediments it becomes clear that these records have to be interpreted and correlated with great caution. From the complexity of processes that can be involved (Fig. 1) it almost seems impossible that any two sections from different regions would show comparable patterns in their δ13C record. However, in reality we observe great similarity of secular trends in δ13C records and their wide application for global chemostratigraphic correlation. This is probably related to the much larger scale of the global carbon cycle relative to local budgets, and

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provides confidence that in many cases the primary global trends can be recognized despite the influence of local factors. Because many of the studied sequences contain lithified sediments, the question on reliability of δ13C shifts in diagenetically affected sediments is especially important. In a study of diagenetic influences on the bulk δ13C signal in the English Chalk Jeans et al. (2012) demonstrated that the δ13C of the lithified red chalk is less altered than that of the soft white chalk, indicating that early diagenetic redox conditions are decisive for alteration of the δ13C record rather than the degree of lithification and cementation of pore space. It can be expected that the isotopic signal in most sediments has been affected by diagenesis, which has important implication for the reconstruction of flux rates but does not hamper chemostratigraphic applicability as long as primary shifts are not obliterated. Successful application of bulk sediment δ13C data for global correlation in a wide range of Mesozoic (e.g. Grötsch et al., 1998; Hesselbo et al., 2002), Paleozoic (e.g. Underwood et al., 1997; Saltzman et al., 2000; Zhu et al., 2006) and Neoproterozoic sequences (e.g. Corsetti and Kaufman, 1994; Hill and Walter, 2000) indicate that first order trends in δ13C are often preserved despite diagenetic overprint. In a review paper on Phanerozoic and Neoproterozoic inorganic carbon isotopes Holser (1997) concluded that “Despite widespread effects of late diagenesis on the carbon isotope record, an important fraction of isotope events can be verified on a global scale”. The global reproducibility of secular trends in bulk δ13C values indicates that many of the factors shown in Fig. 1 are either restricted to specific environmental settings (e.g. to the very shallow marine), or are of subordinate importance compared to the global atmospheric and oceanic carbon fluxes, or they are synchronized by some process of global magnitude. Such synchronizing processes include: (1) eustatic sea-level fluctuations, (2) changes in the globally integrated weathering flux, and (3) major perturbations of the global carbon cycle, e.g. by sudden emission of large amounts of CO2 from large igneous provinces or of methane from clathrates, by global oceanic anoxic events (often associated with high CO2 emissions) or by catastrophic events such as an asteroid impact. Because burial of organic carbon and the formation of authigenic carbonate are controlled by early diagenetic redox conditions (Jeans et al., 2012) the biologically influenced availability of oxygen plays a key role in modulating the global carbon cycle and, thus, the bulk δ13Ccarb record (Katz et al., 2005). Although redox conditions can be expected to vary regionally, it has been proposed that net sequestration of organic matter and formation of authigenic cements may be globally synchronized by a number of factors, including eustasy (Katz et al., 2005; Grotzinger et al., 2011; Schrag et al., 2013). A relation between global sea-level and δ13C has been observed and discussed by numerous authors (e.g. Jenkyns et al., 1994; Mitchell et al., 1996; Voigt and Hilbrecht, 1997; Grant et al., 1999; Gröcke et al., 1999; Jarvis et al., 2001; Jarvis et al., 2002, 2006). Many of the processes summarized in Fig. 1 are influenced by relative sea-level changes, reflecting the susceptibility of local carbon cycles to sea-level driven perturbations. Because sea-level has a major control on the exchange between the open ocean and restricted shallow epeiric or continental seas, the lateral gradients (and, thus, some regional δ13C differences) can be expected to decrease with rising sea-level. Eustatic changes have the potential to modulate the global carbon cycle, e.g. via changing the area of shelf seas where most of the organic carbon is produced and deposited, by changing the amount of organic matter that is remineralized with erosion of previously deposited organicrich sediments as areas are exposed or seafloor topography steepens during lower sea-level, or by influencing the relative weathering fluxes from silicate or carbonate rocks. These mechanisms result in a positive correlation of sea-level and δ13C. However, some periods are characterized by negative correlation of δ13C with sea-level, such as during the Hirnantian and early Silurian δ13C excursions when vast areas of previously deposited carbonates were eroded, thereby introducing large amounts of isotopically enriched DIC to the oceanic reservoir (Immenhauser et al., 2008). This demonstrates that the relationship

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between sea-level and δ13C is not simple and straight forward, as can be expected from an “equation with many variables” (Fig. 1). In summary, an observed δ13C trend represents changes in the global primary seawater DIC values (or globally synchronized early diagenesis) only if it is found in multiple sections on a global scale, in a wide range of depositional settings and with differing degrees of diagenetic alteration (Saltzman and Thomas, 2012). This emphasizes the need to compare a large number of carbon isotope records from various paeoenvironmental settings based on independent dating methods such as a precise biostratigraphic framework or magnetostratigraphy. 1.5. Cretaceous δ13C data and objectives The potential of bulk sediment δ13C data for chemostratigraphic correlation in the Late Cretaceous has been demonstrated especially for Boreal chalk successions as more data became available during the past decades (Scholle and Arthur, 1980; Arthur et al., 1987; Gale et al., 1993; Mitchell et al., 1996; Voigt and Hilbrecht, 1997; Jarvis et al., 2002, 2006; Voigt et al., 2010). Good inter-regional correlation was observed between the English Chalk and sections from Italy, France, Germany and Tunisia, representing various depositional settings (e.g. Jenkyns et al., 1994; Voigt and Hilbrecht, 1997; Voigt, 2000; Jarvis et al., 2002; Voigt et al., 2007, 2008; Galeotti et al., 2009; Robaszynski et al., 2010; Wiese, 2010). A number of detailed global-scale comparisons of δ13C records have been published for the Early Cretaceous (e.g. Herrle et al., 2004; Tsikos et al., 2004; Bornemann et al., 2005) and the Cenozoic (e.g. Zachos et al., 2001). However, few studies on Upper Cretaceous sequences present correlations over larger distances, e.g. between Europe and South America (Crespo de Cabrera et al., 1999) or between Europe and Tibet (Li et al., 2006; Wendler et al., 2009, 2011b). By compiling available Late Cretaceous δ13C data and using an iterative approach to compare them with biostratigraphic and magnetostratigraphic data, this paper aims: (1) to test the reproducibility of individual Late Cretaceous δ13C shifts on a global scale, (2) to provide a scheme for Late Cretaceous global correlation by linking macro- and microfossil biozonations through δ13C records, and (3) to test for isochroneity of first and last occurrences of Late Cretaceous index taxa and to assess their reliability for biostratigraphy. For this purpose, sections from the northern and southern hemisphere as well as from various palaeo-latitudes and ocean basins are compared (Fig. 2, Table 1). At the same time, these sequences represent a wide range of oceanic settings from epicontinental basins to the slope and open ocean, as well as different diagenetic histories. 2. Material and methods 2.1. Compared sections and data sources Most sections presented in the literature have a relatively short stratigraphic range and it is difficult to find long and continuous, highresolution δ13C records. For the English Chalk δ13C reference curve this problem has been circumvented by combining a number of sections, thereby extending the stratigraphic range and eliminating intervals of local condensation or omission (Jarvis et al., 2006). The recent study of a Turonian through Maastrichtian sedimentary sequence from the Tethys Himalaya (Guru section, South Tibet; Wendler et al., 2011b) provides an excellent opportunity for inter-hemispheric comparison of δ13C records. With its paleogeographic position on the northern passive continental margin of the Indian plate at about 30° S during the mid-Cretaceous, the Guru section represents an area in the southern Tethys Ocean for which very few isotope data are available to date (Li et al., 2006; Wendler et al., 2009). Because of its diagenetic history that accompanied the formation of the Himalayas, the material from Guru provides a good test for the robustness of bulk

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Fig. 2. Paleogeographic map. Paleogeographic map of the Early Maastrichtian (70 Ma; modified from Voigt et al., 2012) showing position of the twenty sections compared in this study. Some localities represent several sites (S-Tibet: 3, Gubbio: 3, Boreal: 7).

sediment δ13C stratigraphy. The δ13C data from Wendler et al. (2011b) for Guru are complemented here with 117 new data points for the Upper Campanian and Maastrichtian interval. Detailed information on lithology, microfacies, biostratigraphy, fossil preservation and additional geochemical data (δ18O, % carbonate, TOC and sulfur) from Guru are given in Wendler et al. (2011b). In addition to the data from Guru, the following data sources from nineteen other sections are used for δ13C comparisons: Jarvis et al. (2006) for the English Chalk reference curve and time scale; Voigt and Hilbrecht (1997) for Salzgitter-Salder (Germany); Richardt and Wilmsen (2012) for Anröchte (Germany); Voigt et al. (2007) for Halle-Oerlinghausen (Germany); Voigt et al. (2008) for Wunstorf

(Germany); Voigt et al. (2010, 2012) for Lägerdorf-KronsmoorHemmoor (Germany); Thibault et al. (2012a) and Sheldon (2008) for Stevns-1 (Denmark); Voigt et al. (2012) for Norfolk coast (UK), Tercis (France) and the Campanian through Maastrichtian data from Bottaccione (Italy) and Contessa (Italy); Tsikos et al. (2004) for Gubbio S2 core (Italy); Jenkyns et al. (1994) for Turonian through Lower Campanian high-resolution data from Bottaccione; Corfield et al. (1991) for low-resolution δ13C curve from Bottaccione; Stoll and Schrag (2000) for Turonian through Coniacian data from Contessa; Wagreich et al. (2012) for Postalm (Austria); Jarvis et al. (2002) for El Kef (Tunisia); Pratt et al. (1993) and Locklair et al. (2011) for Berthoud State 3 core (Western Interior, USA); Wendler

Table 1 Summary of paleogeography, depositional setting, stratigraphic range and references for carbon isotope data for the compared sections. Paleogeography

Region

Section

Depositional setting

Stratigraphic range

Carbon isotope data sources

Indian Ocean

Exmouth Plateau

ODP Site 762C

Pelagic

Campanian–Maastrichtian

Central Pacific Ocean

Mid-Pacific mountains Shatsky Rise S-Tibet

DSDP Site 463 ODP Site 1210B Guru Tingri Gongzha El Kef Bottaccione

Pelagic Pelagic Neritic–pelagic Hemipelagic–pelagic Hemipelagic–pelagic Hemipelagic–pelagic Pelagic

Contessa

Pelagic

Gubbio S2 core Postalm Tercis les Bains Stevns-1

Pelagic Hemipelagic–pelagic Inner–outer shelf Epicontinental

Campanian–Maastrichtian Campanian–Maastrichtian Turonian–Maastrichtian Turonian–Santonian Turonian–Santonian Campanian–Maastrichtian Turonian–Maastrichtian Campanian–Maastrichtian Turonian–Campanian Turonian–Coniac Campanian–Maastrichtian Turonian Campanian Campanian–Maastrichtian Campanian–Maastrichtian

English Chalk Reference Norfolk coast Lägerdorf-Kronsmoor-Hemmoor Salzgitter-Salder Anröchte Halle-Oerlinghausen Wunstorf core Berthoud State 3 core

Epicontinental Epicontinental Epicontinental Epicontinental Epicontinental Epicontinental Epicontinental Epicontinental

Turonian–Maastrichtian Campanian–Maastrichtian Coniacian–Maastrichtian Turonian–Coniacian Turonian Turonian Turonian Coniacian–Campanian

Stoll and Schrag (2001), Thibault et al. (2012b) Li and Keller (1999) Jung et al. (2012) Wendler et al. (2011b), this study Wendler et al. (2009) Li et al. (2006) Jarvis et al. (2002) Corfield et al. (1991) Voigt et al. (2012) Jenkyns et al. (1994) Stoll and Schrag (2000) Voigt et al. (2012) Tsikos et al. (2004) Wagreich et al. (2012) Voigt et al. (2012) Thibault et al. (2012a), Sheldon (2008) Jarvis et al. (2006) Voigt et al. (2012) Voigt et al. (2010, 2012) Voigt and Hilbrecht (1997) Richardt and Wilmsen (2012) Voigt et al. (2007) Voigt et al. (2008) Pratt et al. (1993), Locklair et al. (2011)

S-Tethys

NW-Tethys

Boreal Sea

Tunisia Gubbio, Italy

Alps, Austria Aquitaine Basin, France Denmark UK Germany

Western Interior Seaway

Denver Basin, Colorado, USA

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et al. (2009) for Tingri (Tibet); Li et al. (2006) for Gongzha (Tibet) with regional lithostratigraphic correlations from Wendler et al. (2009); Li and Keller (1999) for Site 463 (equatorial Pacific); Jung et al. (2012) for Site 1210B (equatorial Pacific); combined data from Stoll and Schrag (2001) and Thibault et al. (2012b) for Site 762C (Indian Ocean) with additional biostratigraphic data from Howe et al. (2003). Because the detailed δ13C curves for the Berthoud State 3 core (Niobrara Formation, Coniacian through lowermost Campanian) and the Maastrichtian part from Guru (Limestone 3) are rather “noisy”, the 3-point moving average is shown instead. The detailed data points can be seen in Fig. 6 for Guru and in Locklair et al. (2011) for Berthoud State 3 core. A summary on paleogeography, depositional setting and stratigraphic coverage of the compared sections is given in Table 1 and Fig. 2. For the Bottaccione section δ13C data (Corfield et al., 1991; Jenkyns et al., 1994) and biostratigraphic and magnetostratigraphic data (Alvarez et al., 1977; Arthur and Fischer, 1977; Premoli Silva, 1977; Chan et al., 1985; Premoli Silva and Sliter, 1995; Gardin et al., 2001, 2012) were published with two different depth scales. Correlation of the different data sets was achieved through consultation with Isabella Premoli Silva who has intensely worked on the section. Both depth scales are given in Figs. 4 to 6 together with the biostratigraphy and magnetostratigraphy, taking into account 25 m of section that was not covered in the low-resolution δ13C curve of Corfield et al. (1991; dashed line in Fig. 7; Jenkyns et al., 1994; Premoli Silva, written communication; Gardin et al., 2001). The depth scale for the Upper Campanian and Maastrichtian of the combined Bottaccione/Contessa record is from Contessa as in Chauris et al. (1998) and Gardin et al. (2012) and with cumulative thickness from Bottaccione in the upper section part as given in the online data from Voigt et al. (2012). Note that in their graphs Voigt et al. (2012) use a new depth scale for Contessa that differs by 35 m from that used by Gardin et al. (2012) and in earlier studies, in order to adjust the depth to the nearby section Bottaccione. Magnetostratigraphy for Lägerdorf-Kronsmoor-Hemmoor and for Tercis is from Lewy and Odin (2001), and for Site 762C from Galbrun (1992) with updates from Husson et al. (2011, 2012) and Thibault et al. (2012b). 2.2. Age models and methods of correlation All carbon isotope records compared in Figs. 3 to 6 are plotted versus depth except the English Chalk reference curve that is shown relative to time (Jarvis et al., 2006). The original age model published by Jarvis et al. (2006; based on GTS 2004) is given in Figs. 4 to 6. These figures illustrate detailed comparisons of the δ13C records covering the respective intervals together with biostratigraphic information and depth scales for each section for the Turonian (Fig. 4), the Coniacian through Santonian (Fig. 5) and the Campanian through Maastrichtian (Fig. 6). Fig. 3 provides higher resolution correlations for the Cenomanian/Turonian oceanic anoxic event 2 (OAE 2) and for the Early to Middle Turonian post-OAE 2 interval, in which the age model for the English Chalk has been adapted using ammonite zone ages from GTS 2012 (Ogg et al., 2012). In all graphs the δ13C curves are shown to the same horizontal scale except in Fig. 3. For isotope correlations the terminology of δ13C events from the English Chalk (Jarvis et al., 2006) is used, with the addition of few events from Jarvis et al. (2002), Wendler et al. (2011b) and Voigt et al. (2012). Fig. 3 also shows the terminology used for Turonian isotope shifts in the German sections (Voigt et al., 2007, 2008; Richardt and Wilmsen, 2012). Abbreviations used in this paper are summarized in Appendices A to C. Based on the correlations illustrated in Figs. 3 to 6, a compilation of all aligned isotope records with positions of planktic foraminifer zonal boundaries is given in Fig. 7, while an attempt to synthesize these results is presented in Fig. 9. In these two summarizing correlation panels both age models are shown for comparison. Fig. 7 gives the original GTS 2004 ages (Jarvis et al., 2006) for the English Chalk and GTS

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2012 ages for the combined Bottaccione/Contessa record, together with ages for planktic foraminifer zonal boundaries from the different age modes. The position of ammonite zones is correlated from the English Chalk to Bottacione/Contessa based on δ13C correlation as shown in Fig. 7. For the GTS 2012 age model, ages from Ogg et al. (2012) for magnetic polarity chrons and for ammonite zones are used as tie-points for the interval from the OAE 2 to the top of Chron C33n (ages of tie-points are given in Table 2), and the age model from Voigt et al. (2012) is applied between the top of Chron C33n and the Cretaceous/Paleogene boundary. Note that their age model uses 74.1 Ma for the top of Chron C33n instead of 74.3 Ma as in Ogg et al. (2012). Some adjustments in the vertical scales were applied in accordance with biostratigraphy in order to compensate for major variations in sediment accumulation rates and to allow for easier visual comparison between the curves (arrows in Figs. 3 to 6 indicate breaks in depth scales). Precision of the correlations is stepwise increased by combining the different stratigraphic methods. This should not be confused with circular reasoning as it rather represents an iterative approach. A first low-resolution correlation of the sections is based on biostratigraphy and, if available, on magnetostratigraphic data. Within this bio- and magnetostratigraphic framework the δ13C curves are compared on the basis of distinct peaks and sometimes characteristic shapes of isotope excursions, points of inflection and troughs. The main isotope events are then used as tie-points (Table 2), allowing for more detailed correlations in between them. The detailed correlation of similar isotopic patterns can in turn be used to test whether or not the position of a biostratigraphic zonal boundary relative to the isotope shifts is the same among the different sections Table 2 Description and ages of tie-points used for δ13C correlations and GTS 2012 age model for Bottaccione/Contessa (Figs. 7 and 9) and for calculation of accumulation rates (Fig. 10). * = isotope events of Voigt et al. (2012). For abbreviations see Appendices A and C. Tie-points for δ13C correlations δ13C tie-point

Age (Ma)

Description

1 2

93.91 92.43

3 4 5 6

91.16 90.22 89.91 89.50

7

88.21

8

85.92

9 10 11 12

83.65 79.07 75.71 72.80

13

72.20

14

70.59

15

69.13

16 17

68.38 66.20

Lower Turonian δ13C maximum of the Holywell Event Base Middle Turonian δ13C maximum of the Round Down Event Top Middle Turonian δ13C maximum of the Pewsey Event Middle Upper Turonian δ13C minimum of the Bridgewick Event Higher Upper Turonian δ13C maximum of the Hitchwood Event Turonian/Coniacian boundary δ13C minimum of the Navigation Event Middle Coniacian δ13C maximum of the White Fall Event: base of δ13C plateau Lower Santonian δ13C minimum of the Haven Brow Event: after δ13C plateau Santonian/Campanian boundary δ13C maximum (SCBE) Base Upper Campanian δ13C maximum (BUCE) Middle Upper Campanian δ13C minimum (LCE) Uppermost Campanian δ13C maximum before drop of CMBE (= CMBE1*) Campanian/Maastrichtian boundary: second δ13C minimum of CMBE δ13C minimum before δ13C rise towards the middle Maastrichtian maximum First δ13C maximum in the middle Maastrichtian (= MME1*) δ13C minimum after the MME, near top of Chron C31n Uppermost Maastrichtian δ13C minimum (= KPgE3*)

Tie-points for age model based on GTS 2012 (Figs. 3, 7, 9 and 10) 94.57 Base Metoicoceras geslinianum Zone 94.15 Base Neocardioceras juddii Zone 93.90 Base Watinoceras devonense Zone 93.35 Base Mammites nodosoides Zone 92.90 Base Collignoniceras woollgari Zone 90.86 Base Subprionocyclus neptuni Zone 83.64 Top Chron C34n 79.90 Top Chron C33r 74.10 Top Chron C33n

4.04 12.26 18.54 1.42 4.47

Söhlde

1.55

Wunstorf

0.88

Anröchte

2.03

Halle-Oerling.

3.35 6.16 4.20 4.86

Stevns-1

3.64 2.77 1.97 2.21 3.18 1.89 3.63 4.00 2.75

Läg.-Kron.-H.

2.36 2.60 2.34 6.45 4.63 1.59 2.08 2.64 2.15 2.45 1.89 3.00

English Chalk

1.03 0.72 0.92 1.06 1.87 2.48

Site 762C

1.25 1.27 0.90 1.30 1.85 1.33 0.92

Site 1210B

6.25 4.12

El Kef

0.71

Postalm

1.43 1.20 1.53 4.04

Tercis

0.54 1.81 0.53 1.03 2.56 1.53 1.00 1.23 1.31 0.60 0.69 0.77 0.93 1.03 1.67 0.66

Cont./Botta.

6.12 2.09 2.20

Tingri

2.84 3.31 2.13 7.26 5.98 2.90 0.25 0.29 0.37 1.91 2.80 0.90 2.88 2.81 11.33

0.41 0.39 0.43 5.26 5.49 5.93 1.65 1.38

Gongzha Guru Top

2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17

Base

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16

1.48 1.27 0.94 0.31 0.41 1.29 2.29 2.27 4.58 3.36 2.91 0.60 1.61 1.46 0.75 2.18

Duration (Myr) Tie-point interval

or if it is somewhat diachronous. Sediment accumulation rates (AR, uncorrected for compaction) were calculated for the intervals between the main δ13C tie-points (Table 3 and Fig. 10), based on the correlations and GTS 2012 age model from Fig. 7. A more detailed investigation of relative changes in AR, detection of hiatuses and assessment of the reliability of biostratigraphic marker species are possible with the method of simple graphic correlation (e.g. Paul and Lamolda, 2009). Such plots are given for some of the main sections compared in this paper (Figs. 11 and 12). Although intervals of positive and negative δ13C shifts rather than individual peaks are correlated here, the method of graphic correlations requires determination of specific depth values for each shift. These positions are indicated with “+” in Figs. 4 to 6, but it is not stated that these peaks exactly correlate. Accuracy in reflecting the true position of the peak value of an isotopic shift strongly depends on the sample spacing relative to AR in each section, as pointed out in Paul and Lamolda (2009). The peaks chosen for graphic correlation are rather used to test the proposed correlation of the δ13C intervals, to show their consistency with magneto- and biostratigraphy and to visualize AR-changes that are assumed with the δ13C correlations. The correlations presented here integrate the correlations proposed by earlier works and for the most part follow these suggestions. These are correlations between: the English Chalk and Tunisia (Jarvis et al., 2002), the English Chalk and the Italian sections (Jarvis et al., 2006), the English Chalk and the Austrian Postalm section (Wagreich et al., 2012); the English and German Chalk sections (Voigt, 2000; Jarvis et al., 2006; Voigt et al., 2007, 2008, 2010; Richardt and Wilmsen, 2012; Voigt et al., 2012), the German Chalk and the Italian sections (Voigt et al., 2007), the Tibetan sections and the English Chalk (Wendler et al., 2011b), the Indian Ocean record and the Italian and Boreal Chalk sections (Thibault et al., 2012b), and the Campanian– Maastrichtian correlations between the English and German Chalk, Stevns-1, Tercis, Gubbio and Site 1210B (Voigt et al., 2012). In few cases the current interpretation differs from correlations proposed by earlier works: the correlation of the Lower-Middle Turonian (Holywell Event to Round Down Event) between the English Chalk and Gubbio as proposed by Jarvis et al. (2006) and the lowermost Turonian between Eastburn (English Chalk) and Gubbio as proposed by Tsikos et al. (2004). In both cases, the arguments and alternative interpretation for correlating the Italian sections by Voigt et al. (2007) are followed, who suggest (1) a deeper level for the Lulworth Event and (2) that the Bonarelli level at Gubbio covers the entire interval of the OAE 2, including the C-peak, so that the elevated values just above the Bonarelli horizon are correlated with the Holywell Event rather than with the C-peak of the OAE 2 excursion. Correlation between the English Chalk and Bottaccione proposed here further differs from those in Jarvis et al. (2006) for the Middle Coniacian (above the j2-peak) to the Santonian/ Campanian Boundary (SCBE, Figs. 5 and 7), suggesting a lower position for the SCBE for Bottaccione at the top of chron C34n and of the Dicarinella asymetrica Zone. This interpretation is more consistent with earlier studies on the Bottaccione section (Premoli Silva and Sliter, 1995; Gardin et al., 2001) and with results in Gale et al. (2008), as discussed in more detail in the last paragraph of Section 3.8. Most authors cited in the previous paragraph have included correlation of isotopic shifts that are b0.5‰, although it might be questioned if these small shifts are meaningful with respect to global correlation. Especially in sections from shallow epeiric seas with water depths of less than ~100 m it is not advisable to correlate isotopic shifts of such small magnitude because of the generally high lateral and temporal δ13CDIC variability in these settings (Immenhauser et al., 2008). None of the sections compared here represents such shallow marine environments, with the only exception of the Maastrichtian part of Guru that is not correlated in detail but the trend in the 3-point average is compared. Furthermore, most δ13C shifts in the sections with high carbonate content (e.g. the Boreal Chalk successions and the Pacific and Indian Ocean sites) have a relatively small magnitude compared to sections with lower

Salzgitter

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Table 3 Sediment accumulation rates (AR) for intervals between numbered δ13C tie-points as given in Table 2 and Fig. 7. Plots of AR for some of these sections are shown in Fig. 10.

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carbonate content (e.g. Tibetan sections), but nevertheless these small δ13C shifts display comparable patterns that can be correlated on global scale (e.g. Paul and Lamolda, 2009). It seems, therefore, inappropriate to generally define a lower boundary for a meaningful correlation of δ13C shifts, and this boundary should rather be critically evaluated for each case. It should be noted, though, that even small amounts of organic carbon may considerably reduce δ13C values in the bulk carbonate during early diagenesis under reducing conditions, and it has been argued that all of the defined isotope events (except the OAE 2) in the English Chalk could be caused by early diagenesis (Jeans et al., 2012). This view is certainly on the other extreme of interpreting bulk δ13Ccarb data and does not explain the similarities in isotopic patterns that are observed on global scale. However, the study of Jeans et al. (2012) should caution the use of minor δ13C shifts for global correlation. The main isotope excursions that were used here as chemostratigraphic tie-points (Table 2) are ~0.5‰ or larger in most of the correlated sections. In addition, some of the minor peaks are correlated if similar patterns were observed in multiple sections on global scale, because this similarity cannot be explained by intra-sample variability of δ13C values. 3. Results and discussion 3.1. Biostratigraphy Planktic foraminifera in Guru are well preserved in most parts and planktic foraminiferal biozones are comparable to the standard WTethyan biozones with few differences as discussed in Wendler et al. (2011b). The additional Maastrichtian part from Guru presented here consists of shallow marine limestone that is rich in larger benthic foraminifera. These foraminifera include Orbitoides media, Orbitoides tissoti and Orbitoides apiculata, indicating an Early to mid-Maastrichtian age of the sediments. Because Siderolites calcitrapoides first occurred in Tibet in the second half of the Late Maastrichtian (Willems and Zhang, 1993), its absence from the studied material points to a midMaastrichtian age of the top of the measured section in Guru. Biostratigraphic data for the other records were taken from the sources given in Section 2. Thirteen of the twenty sections are dated with planktic foraminifera and nannofossils (Figs. 3 to 6). Biostratigraphy in the seven epicontinental sections from England, Northern Germany and the Western Interior is based on macrofossils with some additional nanno- and microfossil data (Figs. 3 to 6). 3.2. Changes in sediment accumulation rates and hiatuses The sediment accumulation rate (AR) calculated for Guru (Tibet) varies strongly over the section as indicated by bio- and chemostratigraphy as well as by sedimentological evidences (Wendler et al., 2011b). A relatively high rate is calculated for the Turonian and lower part of the Coniacian with a mean of ~3.8 cm/kyr (~7–8 cm/kyr in the Upper Turonian part), and it drops to about 0.3 cm/kyr at the base of the 29 m thick Limestone 1 Member (Upper Coniacian through Lower Campanian). Sharp lithological contacts at the base and top of Limestone 1 and small channels at its base, point to periods of omission or erosion. Low AR during that period are also indicated by comparably small thickness of foraminifer biozones (though all zones are present), e.g. ~10 m for the D. asymetrica Zone as compared to N60 m in Tingri and 40 m in Bottaccione. The transition to the Coniacian plateau of high δ13C values occurs in Guru at the base D. asymetrica Zone, while in Tingri, Bottaccione and Contessa it is found slightly below the first occurrence (FO) of D. asymetrica. Together with the more abrupt change in slope of the δ13C curve in Guru, this supports the assumption of a small hiatus at the base of Limestone 1, possibly spanning the period between the White Fall Event and the j2-peak (Fig. 7). Correlations in Figs. 5 and 6 also suggest that the positive δ13C shifts of the Santonian/Campanian Boundary Event (SCBE) and the Base Upper Campanian Event (BUCE, in the Middle Campanian of the GTS 2012) are not fully represented in

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Guru. The section parts above and below the Limestone 1 in Guru appear to be complete, and the Late Campanian AR is estimated at 2.8 cm/kyr (Wendler et al., 2011b). To account for the condensed sedimentation of the Limestone 1, the vertical scale for this interval has been expanded to make visual comparison with other sections easier (Figs. 5 and 6). Through this adaptation it becomes clear that the present sample spacing in Guru is insufficient for resolving small-scale δ13C variations for the Middle Coniacian to Early Campanian, while it provides very good resolution for the Turonian through Early Coniacian and for the Late Campanian through Early Maastrichtian (Fig. 7). Similar to Guru, AR at Tingri and Gongzha (both located at the northern slope of the Zhepure mountain range in South Tibet) were high in the Late Turonian, but remained high into the Middle Coniacian with ~5–6 cm/kyr. The Upper Cenomanian through Lower Turonian part in Gongzha is relatively condensed with AR of ~0.4 cm/kyr. For both sections, a decrease in AR is indicated for the Middle Coniacian, but in contrast to Guru values remained at moderate levels of ~2 cm/kyr throughout the Late Coniacian and Santonian, and without indication of any hiatus. Tectonic faults complicate the uppermost Santonian and Campanian at Tingri and Gongzha, and the Campanian is largely missing in this area (Wendler et al., 2011b). Seafloor paleo-topography and local bottom currents in the epicontinental Chalk Sea resulted in considerable variation of AR among the different Chalk sections. Common hardgrounds in these sections indicate periods of condensed sedimentation or omission, including a number of hiatuses that are variably present among the sections and represent different length of non-deposition. Based on numerous biostratigraphic and lithostratigraphic marker beds the different Chalk successions can be correlated in great detail (e.g. Gale, 1996; Gale et al., 2005; Jarvis et al., 2006), which allowed producing a stacked δ13C reference curve for the English Chalk that compensates for local periods of condensed sedimentation or non-deposition (Jarvis et al., 2006). Using GTS 2012 ages (Table 2), AR at Trunch section were also high in the Late Turonian (~6–7 cm/kyr) and then dropped to ~2 cm/kyr in the Coniacian through Campanian, which is a typical value of Late Cretaceous Chalks (Scholle et al., 1983; Jarvis et al., 2002). A similar value was also reported for the Campanian and Maastrichtian of the Rørdal core from the Danish Basin, whereas the mean AR calculated for the Stevns-1 core (Danish Basin) is comparably high with 5.7 cm/kyr, and apparently represents continuous sedimentation (Thibault et al., 2012a). The average AR at the N-German Chalk section Lägerdorf-KronsmoorHemmoor was 2.4–2.5 cm/ka from the Santonian through Early Maastrichtian (Schulz et al., 1984; Ehrmann, 1986; Voigt et al., 2010). However, some variability in AR during this period becomes apparent when applying macrofossil and chemostratigraphic correlations between Lägerdorf-Kronsmoor-Hemmoor and the English Chalk (Voigt et al., 2010), which results in values of ~3–4 cm/kyr for the Santonian, ~ 2 cm/kyr for the Campanian and ~ 3–4 cm/kyr for the Maastrichtian. Late Cenomanian through Middle Turonian average AR at Wunstorf and Halle-Oerlinghausen (both N-Germany) were around 2 cm/kyr with higher values around 3 cm/kyr during OAE 2 at Wunstorf (Voigt et al., 2007, 2008). The calculated Early through Middle Turonian AR at Anröchte (N-German Münsterland Basin) are ~1 cm/kyr. In the N-German Lower Saxony Basin at Salzgitter-Salder and Söhlde, the large thickness of Upper Turonian sediments (N120 m at SalzgitterSalder as compared to b 20 m at Contessa, Fig. 4) reflects a very high AR that increased from ~1.4 cm/kyr in the early Middle Turonian (Söhlde) and ~4–5 cm/kyr in the late Middle Turonian to ~18 cm/kyr in the latest Turonian at Salzgitter-Salder. The Italian sections in the Gubbio area (Contessa, Bottaccione and Gubbio S2) are considered to represent continuous pelagic sedimentation that can locally be interrupted by younger tectonic faulting (Arthur and Fischer, 1977). At Bottaccione, a ~10 m hiatus in the top of Chron C31n and another gap in the Radotruncana calcarata Zone have been suggested based on biostratigraphic and chemostratigraphic comparisons (Gardin et al., 2012; Voigt et al., 2012). The AR in these

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Italian sections remained relatively low (~ 0.5–1.8 cm/kyr) during the Turonian through Maastrichtian, but appear slightly elevated (~2–3 cm/kyr) near the Turonian/Coniacian boundary. Based on the GTS 2012 ages for magnetic polarity Chrons, short periods in the Maastrichtian (Chrons C31n and C32n2n) with AR of ~2–2.5 cm/kyr were calculated for Contessa and Bottaccione (Gardin et al., 2012). The Tercis section represents the GSSP for the base Maastrichtian, and the section has been intensely studied. The lithostratigraphic and biostratigraphic results are summarized in a dedicated volume (Odin, 2001). In their chapter on planktic foraminifera of the section, Odin et al. (2001) report that they did not find evidence for any important breaks in sedimentation at Tercis, and suggest an average AR of 2.5 cm/kyr, following Odin and Amorosi (2001). The δ13C correlations for Tercis as proposed by Voigt et al. (2012) suggest AR of ~1.5 cm/kyr in the Campanian and an increase to ~4 cm/kyr in the Maastrichtian. An alternative interpretation for correlating the Maastrichtian part of the section (see Section 3.6) would invoke a less severe increase in AR (version * in Fig. 6; version ** assumes no change in AR). For the Tunisian section El Kef, an average Campanian AR of ~5 cm/kyr can be calculated based on the correlations with the English Chalk as suggested by Jarvis et al. (2002). Similar to Tercis, an alternative interpretation is possible for correlation of the upper section part at El Kef (*** in Figs. 6 and 7, see also Section 3.10), which would imply reduced AR or a hiatus in the post-LCE interval of the section. The Campanian sediments of the Austrian Postalm section represent pelagic deposition of marl and limestone at low AR (Wagreich et al., 2012). While these authors report an AR of ~2 cm/kyr for the Postalm section based on the their assumed duration of the R. calcarata Zone, their δ13C correlation with the English Chalk suggests a lower AR of ~0.7 cm/kyr for the ~24 m thick sequence between the BUCE and the LCE that represents 3–4 Myr (Table 3, Fig. 7). This discrepancy reflects a number of uncertainties that arise from correlating this interval between sections based on microfossil and macrofossil biostratigraphy, and from problems with the determination of the top R. calcarata Zone and its consequence for assessing the duration of this zone as discussed in Section 3.9. The average AR for the Niobrara Formation in the Western Interior (Coniacian through lowermost Campanian) was estimated to 1.4 cm/kyr, based on dated bentonites (Obradovich, 1993; Locklair and Sageman, 2008). This value agrees well with the average AR of ~ 1.3 cm/kyr for the ~ 90 m thick Niobrara Formation at Berthoud State 3 core with respect to the ~ 7 Myr represented by the record, based on δ13C correlations and the GTS 2012 (Fig. 7). Relatively continuous pelagic sedimentation is indicated for the Campanian and Maastrichtian at Site 1210B in the equatorial Pacific with AR of ~ 1.3 cm/kyr (Jung et al., 2012). The Indian Ocean Site 762C at the northwestern margin of Australia (Exmouth Plateau) has recently been astronomically calibrated, allowing for a detailed reconstruction of AR at this site (Thibault et al., 2012b). According to these calibrations, the Campanian AR at Site 762C varied between ~0.7 and 1.8 cm/kyr, then dropped to about 0.6 cm/kyr in the Early Maastrichtian and increased again to ~1–2.5 cm/kyr in the Late Maastrichtian. These results also indicate a hiatus of about 500 kyr during Chron C31n, which is taken into account in the correlations presented here. 3.3. Regional correlation for Tibet Regional comparison of the Tibetan δ13C records from Guru, Tingri and Gongzha shows similar trends among all three sections. A δ13C minimum characterizes the base Helvetotruncana helvetica Zone at Guru and Gongzha, followed by a broad interval of higher values (correlated to the Round Down Event) before they decrease again in the upper part of the zone (Fig. 4). The positive shift in the top H. helvetica Zone at Guru and Gongzha is interpreted to represent the late Middle Turonian Pewsey Event, and reduced AR or a small hiatus is suggested below this interval at Gongzha. The Pewsey Event is followed by an interval of lower δ13C values, above which a pronounced positive shift is found in

the Upper Turonian of N2‰ at Guru and ~0.7‰ at Gongzha, thought to represent the Late Turonian Hitchwood Event (Wendler et al., 2011b). The base Dicarinella concavata Zone is defined by the FO of the nominate taxon. The difference in the stratigraphic range of this zone between Guru, Gongzha and the Italian sections can be explained by the rarity of the species in the Middle Turonian while it becomes abundant in the latest Turonian (Robaszynski and Caron, 1995). Rare specimens of D. concavata were found in Guru at 101 m, 141 m and 161 m (washed samples) and the species becomes abundant above 177 m. Biostratigraphy in the limestone of the two Italian sections is based on thinsections and might not have recorded the earliest rare occurrence of D. concavata. Following the Hitchwood Event, the δ13C values decrease towards the Turonian/Coniacian boundary in all three Tibetan sections (Figs. 4 and 5). Above this minimum, a strong positive shift occurs in all three sections in the Coniacian to reach a first maximum below the FO of D. asymetrica. This maximum is correlated to the White Fall Event and forms the beginning of a plateau of high values in the Upper Coniacian (upper D. concavata Zone and lower D. asymetrica Zone, Fig. 5). A sudden drop of ~0.8‰ is found in the middle D. asymetrica Zone at Tingri and Gongzha and forms the end of this isotope plateau. Also at Guru, the δ13C values decrease by more than 0.5‰ at the respective level, although this interval cannot be correlated with certainty because of the low AR and sampling resolution in this section part. The exact stratigraphic position of the section top in Tingri has been uncertain because D. asymetrica is still present in the top. However, comparison of planktic foraminifer assemblages between Tingri and Guru indicates that it has a lower stratigraphic level than the top D. asymetrica Zone (Wendler et al., 2011b). This means that the SCBE is not present in the Tingri record as was tentatively suggested in Wendler et al. (2009). The new δ13C and biostratigraphic data from Guru therefore allow for more refined correlation of the isotope curves from Tingri and Gongzha. 3.4. Global δ13C comparison: Turonian Figs. 3 and 4 present δ13C data and biostratigraphy for Turonian sediments from the Boreal (English Chalk and the German sections Salzgitter-Salder, Halle-Oerlinghausen, Anröchte and Wunstorf), the NW-Tethys (Italian sections Contessa, Bottaccione and Gubbio S2) and the S-Tethys (Tibetan sections Guru and Gongzha). Comparison of the isotope curves shows great similarity in secular δ13C variations from these sections, but also reveals differences in absolute values and amplitude of the isotopic shifts. The amplitude of the OAE 2 is especially large in the English Chalk, resulting in a stronger decreasing trend above the OAE 2 from the Lower Turonian through the lower part of the Upper Turonian. This decreasing trend is less pronounced in the German sections (Fig. 3), and the general trend in Guru and in the Italian sections shows δ13C values remain relatively high before they start to decrease in the middle H. helvetica Zone (Middle Turonian, Fig. 4). Apart from the OAE 2 excursion, two major intervals with elevated δ13C values are evident in all of the compared sections that cover the respective levels: in the lower part of the Middle Turonian (lower Collignoniceras woollgari Zone and lower/middle H. helvetica Zone) and in the Upper Turonian, corresponding to the Round Down and Hitchwood Events in the English Chalk, respectively. The Upper Turonian excursion looks relatively symmetrical in the English Chalk, while it shows a steeper increase and a more gradual decrease in the other five sections (Fig. 4) and in the Spanish Liencres section (Wiese, 1999; Jarvis et al., 2006). Except at Trunch, which was used for this part of the English Chalk Reference curve, most of the other English Chalk sections also have a lower thickness between the Bridgewick and Hitchwood Event than between the Hitchwood and Navigation Event (Jarvis et al., 2006). Between the Round Down and Hitchwood Event, a decreasing trend is obvious from all records in Fig. 4. Superimposed on this trend is a smaller positive shift in the upper Middle Turonian, at the level of the Pewsey Event in the English Chalk. Between the Pewsey Event and the Hitchwood Event a relatively broad interval of low δ13C values is evident in the lower

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Fig. 3. Cenomanian–Turonian OAE 2 and post-OAE 2 interval: δ13C correlation and biostratigraphy. Biostratigraphy and correlation of Upper Cenomanian to Middle Turonian δ13C records from the Boreal Realm (English Chalk, UK; Halle-Oerlinghausen, Anröchte and Wunstorf, Germany), the southern Tethys (Guru and Gongzha, Tibet) and the north-western Tethys (Gubbio S2 and Contessa, Italy). Data sources are given in Section 2.1 and Table 1. Data for the Bonaralli level in the Gubbio S2 core are from organic carbon (bold gray curve; from Tsikos et al., 2004), all other curves are from bulk carbonate. Dark and light gray shadings mark correlation of positive and negative isotope shifts, respectively. English Chalk curve is shown relative to age, adapted from Jarvis et al. (2006) and using GTS 2012 ages for ammonite zones as tie points (marked by triangles). All other curves are plotted relative to thickness (indicated after section name). Some adjustments in vertical scale (arrows mark breaks in scale) were performed to compensate for changes in SR during and after the OAE 2, allowing for better visual comparison. For more details on biostratigraphy and section meters for Gongzha, Guru and Contessa see Fig. 4, and for Gubbio S2, Anröchte and Wunstorf see Tsikos et al. (2004), Richardt and Wilmsen (2012) and Voigt et al. (2008), respectively. Carbon isotope events for English Chalk from Jarvis et al. (2006), and for Wunstorf from Voigt et al. (2007). For abbreviations see Appendices A to C.

part of the Upper Turonian from all six records. This part is remarkably similar between the English Chalk, Guru, Salzgitter-Salder and Contessa, showing three minor peaks punctuating the broad δ13C low. Similarly, the minor double peaks of the Holywell Event and the c1/c2 shifts in the English Chalk might also be recognizable in the Lower Turonian (upper Whiteinella archaeocretacea Zone) at Guru, Gongzha and Contessa.

3.5. Global δ13C comparison: Coniacian through Santonian Fig. 5 compares δ13C and biostratigraphic data from the English and German Chalk (Boreal), two Italian sections (NW Tethys) and three Tibetan sections (S-Tethys). Following the Upper Turonian Hitchwood Event, a pronounced δ13C minimum characterizes the base of the Coniacian in all sections, corresponding to the Navigation Event in the English Chalk. This event is also present in the sections SalzgitterSalder and Słupia Nadbrzeżna that were proposed as GSSP for the Coniacian (Walaszczyk et al., 2010). Slightly higher δ13C values, followed by another δ13C minimum (East Cliff Event) are found in the Lower Coniacian of the English Calk and in the lower/middle D. concavata Zone in Italy and Tibet. This interval appears expanded in the Italian and Tibetan sections relative to the Chalk.

A characteristic feature in the compared records is a δ13C increase in the Middle Coniacian, following the East Cliff Event and reaching a first peak at the White Fall Event. This shift occurs in the upper part of the D. concavata Zone in the Tibetan and Italian sections. The δ13C values remain elevated through the Middle and Upper Coniacian and into the Lower Santonian in all seven sections. The two intervals of highest δ13C values in the upper half of this isotopic plateau, above the base of the D. asymetrica Zone in Tibet and Italy, might correspond to the j2/Kingsdown and k1/k2 peaks in the English Chalk. Lower and more variable δ13C values are found through the Middle and Upper Santonian, bracketed by two minima that are well comparable between the sections (Fig. 5). A positive shift in the uppermost Santonian (upper D. asymetrica Zone) marks the beginning of the double peaked Santonian/Campanian Boundary Event (SCBE). 3.6. Global δ13C comparison: Campanian through Maastrichtian The Campanian and Maastrichtian climatic decline of the Late Cretaceous greenhouse has been intensely studied, and a large number of stable isotope records have become available with good global coverage (D'Hondt and Lindinger, 1994; Jenkyns et al., 1995; Stenvall, 1997; Stoll and Schrag, 2001; Jarvis et al., 2002; Paul and Lamolda, 2007; Melinte-

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Fig. 4. Turonian δ13C correlation and biostratigraphy. Biostratigraphy and correlation of Turonian bulk sediment δ13C records from the Boreal Realm (English Chalk, UK; Salzgitter-Salder, Germany), the southern Tethys (Guru and Gongzha, Tibet) and the north-western Tethys (Contessa and Bottaccione, Italy). Data sources are given in Section 2.1 and Table 1. Age model (GTS 2004) and carbon isotope events for English Chalk from Jarvis et al. (2006); for GTS 2012 age model see Figs. 3 and 7. Dark and light gray shadings mark correlation of positive and negative isotope shifts, respectively. Note great inter-hemispheric similarity in δ13C patterns especially for the Middle and Late Turonian and regional differences for Early Turonian longterm trends. See Section 3.9 for discussion on the base D. concavata Zone. For Bottaccione both scales are given as published [A]: for bio- and magnetostratigraphy and [B]: for isotopes. Arrows mark breaks in scale to compensate for large changes in AR. + = positions for graphic correlation in Fig. 11A–C. For abbreviations see Appendices A to C.

Dobrinescu and Bojar, 2009; Voigt et al., 2010; Wendler et al., 2011b; Jung et al., 2012; Thibault et al., 2012a; Thibault et al., 2012b; Voigt et al., 2012). In Fig. 6, δ13C data from eleven sections are compared from the Boreal (English, German and Danish Chalk), the NW-Tethys (sections from Tunisia, Austria, France and Italy), the S-Tethys (Tibetan Guru section), the Indian Ocean (Site 762C) and the Pacific Ocean (Site 1210B). In the Lower Campanian, these δ13C records show little secular variation with generally elevated values and a stable to slightly decreasing trend (Fig. 6). A small positive δ13C shift at the base of the Upper Campanian (BUCE) is followed by a somewhat stronger decreasing trend in the lower half of the Upper Campanian. This trend is relatively steady in the English and German Chalk sections but shows a pronounced minimum below the R. calcarata Zone in Guru that can be attributed to remineralization of organic matter as evidenced by reworked foraminifera in some samples of this interval (Wendler et al., 2011b). A positive δ13C shift of ~1‰ near the base R. calcarata Zone (Base Calcarata Eevent, BCE) in Guru is correlated to peaks at a similar level at Postalm, Bottaccione, El Kef, Site 1210B and Tercis. This δ13C maximum can also be correlated to the boreal Chalk sections based on correlations between the English Chalk and El Kef (Jarvis et al., 2002; Fig. 6), and to Site 305 at Shatsky Rise (equatorial Pacific, not shown in Fig. 6) where a similar positive shift is found near the base R. calcarata Zone (Voigt et al., 2010). The most pronounced feature in the compared Campanian δ13C records is a broad negative shift above the middle of the Upper Campanian

(base Upper Campanian in the GTS 2012), called the Late Campanian Event (LCE, Fig. 6). In Northern Germany it occurs in the Nostoceras polyplocum Zone. The position of the planktic foraminifer biozonation relative to the LCE varies among the sections and seems to depend on how well the top of the R. calcarata Zone can be identified (see Section 3.9). In many sections the LCE is a regular, V-shaped minimum with a magnitude of 0.5–1‰. Although it is very well developed in the S-Tethys, the LCE is less pronounced in the compared sections from the Pacific and the Indian Ocean. Nevertheless, in most cases the LCE represents a very good tie-point for global correlation. Correlations of minor peaks above and below the LCE are tentative. In all sections compared, the LCE is followed by a broad interval of elevated δ13C values in the higher part of the Upper Campanian. The Campanian/Maastrichtian boundary is characterized by a strong δ13C decrease (CMBE) towards a broad δ13C low in the Lower Maastrichtian, with several minor sub-peaks (CMBE 1 to 5 of Voigt et al. (2012)) that appear to vary in magnitude among the sections (Figs. 6, 7). This broad isotopic minimum is followed by a pronounced stepwise increase of δ13C values towards the Mid-Maastrichtian Event (MME) that consists of two maxima interrupted by a short minimum. After reaching a third maximum, the δ13C values subsequently decrease again towards the K/Pg. The detailed correlation at the K/Pg is unclear as this boundary appears to occur at a positive δ13C shift at Bottaccione and at a negative shift at Stevns-1. As mentioned in Section 3.2, two alternative interpretations

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Fig. 5. Coniacian through Santonian δ13C correlation, magneto- and biostratigraphy. Biostratigraphy, magnetostratigraphy and correlation of Coniacian through Santonian bulk sediment δ13C records from the Boreal Realm (English Chalk, UK; Lägerdorf-Kronsmoor-Hemmoor, Germany), the north-western Tethys (Contessa and Bottaccione, Italy) and the southern Tethys (Guru, Tingri and Gongzha, Tibet). Data sources are given in Section 2.1 and Table 1. Age model (GTS 2004) and carbon isotope events for English Chalk from Jarvis et al. (2006); for GTS 2012 age model see Fig. 7. Dark and light gray shadings mark correlation of positive and negative isotope shifts, respectively. Arrows mark breaks in scale to compensate for large changes in AR. + = positions for graphic correlation in Fig. 11D–F. For abbreviations see Appendices A to C.

for the Maastrichtian part of the Tercis section imply a less severe increase in AR and are indicated by arrows with asterisk that mark the respective positions of the section top. The version marked with * appears plausible regarding the similarity to δ13C trends at the respective interval at Site 762C (Fig. 7). 3.7. Late Cretaceous global δ13C correlation A summary of the Upper Cretaceous δ13C records discussed above is presented in Fig. 7, with two additional records from the Pacific Ocean (Site 463, thickness ~150 m) and the Western Interior (Berthoud State 3 core, thickness ~90 m). The curves have been aligned according to the correlations shown in Figs. 3 to 6 and using the major isotopic shifts as tie-points (Table 2). The stage boundaries as given in Locklair et al. (2011) were used to align the Western Interior record. On the left and right hand side in Fig. 7, the GTS 2004 time scale and respective GTS 2012 ages are shown for the English Chalk and for the combined Bottaccione/Contessa record, respectively, allowing for comparison of the two age models. Note that tie-point number 6 lies at the δ13C minimum of the Navigation event at the Turonian/Coniacian boundary in the English Chalk. The respective δ13C minimum at Bottaccione lies at ~89.5 Ma using the GTS 2012 ages for the top of Chron C34n and for the base Subprionocyclus neptuni Zone, but the GTS 2012 age for the Turonian/Coniacian boundary is 89.77 Ma. This discrepancy reflects the age difference of the base Coniacian between the GTS 2004 and 2012. Isotopic shifts that are common among all or most of the compared sections are marked with dark and light gray shading in Fig. 7. Because the records are plotted to the same horizontal scale, the

graph not only visualizes similarities in most of the secular δ13C trends but also shows the regional differences in amplitude of the δ13C shifts and in absolute δ13C values. Carbon isotopes in Guru show long-term variation between 0 and 3‰. This magnitude of about 3‰ is larger than the 1–2‰ shifts in the two other Tibetan sections (Tingri and Gongzha), and about three times larger than the Upper Cretaceous δ13C shifts in the English, German and Danish Chalk (except OAE 2 and CMBE), and in the sections from the NW-Tethys and the Western Interior (Fig. 7). The lowest amplitudes are observed for the Pacific and Indian Ocean sites. Most of the absolute δ13C values in all three Tibetan sections vary between 0.5 and 2.5‰ with a mean of ~1.5–2‰. These values are comparable to the data from El Kef and from the Western Interior site, but are lower than in the more carbonate-rich sediments from the Boreal Chalk sections and the NW-Tethyan sections, where Upper Cretaceous δ13C values mainly vary between 2 and 3‰ (except the LCE and CMBE). Generally high values between 2.5 and 3‰ were obtained from the Pacific and Indian Ocean sites. Despite the differences in amplitude and offsets in absolute δ13C values there is good agreement of long-term δ13C trends in all the sections compared in Fig. 7, based on their biostratigraphic and magnetostratigraphic framework. The major δ13C excursions are clearly recognizable in all sections covering the respective intervals. As explained in Section 1.3, this is strong evidence for many of the δ13C shifts to have been caused by fluctuations in the global oceanic DIC, supposedly related to changes in climate, oceanic circulation and sea-level and their influence on the partitioning of carbon between organic and carbonate carbon sinks (e.g. Voigt and Hilbrecht, 1997; Jarvis et al.,

130 I. Wendler / Earth-Science Reviews 126 (2013) 116–146 Fig. 6. Campanian through Maastrichtian δ13C correlation, magneto- and biostratigraphy. Biostratigraphy, magnetostratigraphy and correlation of Campanian through Maastrichtian bulk sediment δ13C records from the Boreal Realm (English Chalk, UK; Lägerdorf-Kronsmoor-Hemmoor, Germany; Stevns-1, Denmark), the north-western Tethys (El Kef, Tunisia; Postalm, Austria; Bottaccione, Italy; Tercis, France), the southern Tethys (Guru, Tibet), the Western Interior (Berthoud 3 State core), the Indian Ocean (Site 762C) and the equatorial Pacific (Site 1210B). The Berthoud 3 State core record and the Maastrichtian part from Guru show the 3-point moving average. Data points in Guru: black from this study, gray from Wendler et al. 2011. Data sources are given in Section 2.1 and Table 1. Dark and light gray shadings mark correlation of positive and negative isotope shifts, respectively. A strong increase of AR in the Maastrichtian at Tercis is assumed to follow correlations of Voigt et al. (2012), but two alternative interpretations for correlating the Maastrictian part are indicated by arrows with asterisk (arrows point to position of section top). Similarly, the post-LCE part at El Kef could alternatively be correlated to the Lower Maastrichtian (***). The enhanced δ13C minimum near 230 m in Guru represents local influence of organic matter reworking during relative sea-level lowstand. Uncertainties in the R. calcarata zonal boundaries are explained in Section 3.9. Depth scales for Lower Campanian of Bottaccione are shown on the left as published for isotopes and on the right as published for bio- and magnetostratigraphy. References: 1 = Jarvis et al. (2006); 2 = Voigt et al. (2012); 3 = Thibault et al. (2012a). For abbreviations see Appendices A to C.

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2002, 2006). A recent study demonstrated detailed correlation of Upper Cretaceous δ13C shifts between two German sections (Salzgitter and Burgberg) whereby absolute values differed by ~ 1‰ (Wiese, 2010). Differences in absolute values and amplitude most likely reflect local influences on the water chemistry by factors such as nutrient levels and primary productivity, bottom water age, fluvial supply of terrestrial organic matter, shelf topography and its influence on accumulation rates and sediment re-distribution (Fig. 1), but they may also express the varying contribution of early diagenetic cements. Given the large distances between the compared sites it is remarkable how similar in overall shape the δ13C curves are in some intervals. However, Fig. 7 also reveals that consistency in δ13C shifts among the sections varies across the Upper Cretaceous. There is striking similarity among the curves for the Turonian, apparently allowing for correlation of minor isotope shifts such as the Pewsey, Caburn and Southerham Events. Relatively detailed correlation also seems possible for the Coniacian and Santonian, though the number of high-resolution records is limited for this interval. Note that in the Lower/Middle Coniacian of the Berthoud State 3 core, the high-resolution data show a steeper increase towards a first peak at the level of the White Fall Event (Locklair et al., 2011) that is not apparent from the 3-point moving average shown here (the high-resolution curve is rather “noisy” in some intervals). In contrast to the Turonian through Santonian, more regional variability seems to be expressed in the Campanian and Maastrichtian records, making δ13C correlation difficult apart from the major isotope events (such as the LCE or CMBE) that can be recognized in all sections. Additionally, chemostratigraphic correlation is severely hampered for periods with relatively stable δ13C values such as in the Early Campanian. This situation might improve as more detailed δ13C records from this period become available. For Site 463, bulk δ13C data were not available but it is interesting to note that the Δδ13C data very nicely parallels the trends in the δ13C curve from Guru. These Δδ13C values represent δ13C differences between planktic and benthic foraminifera and were interpreted to reflect variations in surface water productivity and deep ocean circulation (Li and Keller, 1999). Superposition of global changes in oceanic δ13CDIC and local signals can also cause regional differences in long-term trends despite good agreement of distinct isotope shifts (e.g. the Round Down and Hitchwood Event or LCE) within the given bio- and magnetostratigraphic frame. For example, in the English Chalk and in Gongzha, a long-term decrease is superimposed on the Lower through middle Upper Turonian carbon isotope events, whereas in Guru, in both Italian sections and in the German sections the overlaying trend after the Holywell Event is rather stable (or even increasing at Wunstorf and Halle-Oerlinghausen) and only starts to decrease after the Round Down Event (Fig. 7). Another example is the Middle Campanian, where locally decreased δ13C values below the R. calcarata Zone in Guru cause an increasing trend in the upper part of the Middle Campanian as opposed to stable or decreasing trends in the equivalent interval of most other sections, while recognition of the LCE is unequivocal due to its stratigraphic position and characteristic shape. The pronounced δ13C minimum just below the R. calcarata Zone in Guru likely reflects reworking during a relative sea-level lowstand and subsequent transgression (Wendler et al., 2011b), possibly related to local tectonics. It could, however, correspond to a global sea-level lowstand around 77 Ma (Miller et al., 2005). A lowstand just below the R. calcarata Zone and transgression within this zone has been suggested for the New Jersey coastal plain based on sequence stratigraphy and planktic foraminifer distribution (Georgescu, 2006). The hiatuses and strong changes in AR at the base and top of the Limestone 1 in Guru are likely related to the northward drift of the Indian plate from the temperate to the arid climate belt in the Middle Coniacian and into the tropics during the Late Campanian (Wendler et al., 2011b). These strong environmental changes caused large fluctuations in carbonate production and clastic influx that are expressed in the strongly varying lithology across the section and can explain the

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comparably large δ13C shifts in Guru. The patterns of lithology and δ13C across the Guru section and among the sections compared in Fig. 7 indicate a relation between the magnitude of δ13C shifts and the content of carbonate and TOC in the sediments. Sequences with high carbonate content display generally low amplitudes in δ13C shifts (except OAE 2) whereas larger amplitudes of δ13C shifts are observed in sediments with lower carbonate and higher TOC content. At the same time, the δ13C data from sediments with lower carbonate content appear to be more “noisy” such as in El Kef or in the Turonian part of Guru, for which eutrophic conditions were reconstructed (Wendler et al., 2011b) and where enhanced fluctuations in organic matter production and degradation can be expected. In contrast, the highest absolute δ13C values with lowest fluctuations are observed from the equatorial Pacific and Indian Ocean sites that represent carbonate-rich (96–100 wt.% at Site 1210B; Jung et al., 2012) sediments from comparably low-trophic, upper bathyal to upper abyssal settings (Thierstein, 1979; Zepeda, 1998; Kaiho, 1999). One possibility to explain smaller δ13C amplitudes in deeper setting is a depth-dependent truncation of the δ13C shifts by carbonate dissolution, as was suggested for the Paleocene– Eocene δ13C excursion (McCarren et al., 2008). However, for most of the less extreme δ13C shifts an influence by early diagenetic effects appears to offer a more plausible explanation. Because degradation of organic carbon is the main driver for early diagenetic processes (see Section 1.3.3), sediments with high TOC and low carbonate contents would have a lot of “fuel” but little capacity to “buffer” these reactions, and a higher contribution of authigenic carbonate (typically with low δ13C values) can be expected if early diagenetic conditions were anoxic (Jeans et al., 2012). Such sediments are typically found in nutrient-rich settings close to the continental shelf rather than at deeper sites in the open ocean. It has been suggested that δ13C values should not be considered reliable if the ratio of carbonate carbon to organic carbon is below 7:1 (Scholle and Arthur, 1980; Saltzman and Thomas, 2012). The lowest carbonate contents in Guru are found in the Lower and Middle Turonian part, but most of these values vary between 30 and 50% and with the exception of very few samples the ratio of carbonate carbon to organic carbon is well above 7:1. This explains why the δ13C curve in this part is rather “noisy” but isotope events can nevertheless be recognized. An example for such a low ratio can be seen in TDP Site 31 that was drilled in Tanzania and recovered very clay-rich Turonian sediments with carbonate contents usually b20% and averaging at ~13%, and with moderately high TOC values averaging at ~1% (as opposed to ~0.3% in the Turonian at Guru) but with some values over 2% (Jiménez Berrocoso et al., 2012). Despite the large abundance of exceptionally well preserved microfossils in the material from Site TDP 31 (Wendler et al., 2011a; Wendler and Bown, 2013; Wendler et al., 2013; MacLeod et al., in press; Wendler et al., in press), the bulk sediment δ13C values are comparably low (mostly between 1‰ and −5‰) and show large fluctuations that make correlative interpretations very difficult. Apart from the early diagenetic effects, there is also a primary influence of carbonate versus organic carbon production on ocean water δ13CDIC. While periods of increased production (and preservation) of organic matter are characterized by positive shifts in δ13C (e.g. during the OAEs), an increased burial flux of carbonate tends to decrease the δ13C signal, which has the potential to affect the carbon isotope records on regional and on global scale (Locklair et al., 2011). Correlative patterns between δ13C, TOC and carbonate content have been repeatedly observed in the sedimentary record. Generally, in sediments from higher tropic environments with low carbonate and high organic matter fluxes there is a positive correlation between TOC and δ13C and their negative correlation to the carbonate content, while co-variation between carbonate and δ13C are observed during periods with high carbonate burial flux (Wendler et al., 2009; Wendler et al., 2011b). This pattern can be explained by the opposing effects of early diagenesis and export production of TOC and carbonate on the δ13C signal. For example, elevated organic matter production increases δ13Ccarb through C12 depletion of the surface waters, but elevated TOC contents in the

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sediments lower δ13Ccarb through early diagenetic C12 release to pore waters. The interplay of these two processes can lead to varying net results in the final bulk δ13C signal, depending on which process is the dominating one. With this respect, not only the export production of organic and inorganic carbon have to be considered, but the preservation potential for organic carbon plays a key role and depends on a number of factors that influence the oxygen exposure time in the sediments (Fig. 1). Because 13C is enriched in carbonates (see Section 1.3.2), comparably higher δ13C values are typically obtained for carbonate-rich sediments relative to carbonate-poor sediments. However, the prolonged production of carbonate leads to a long-term δ13C decrease in the DIC, which can have a global influence on the δ13C records, e.g. during OAE 3 (Locklair et al., 2011).

3.8. Relation of δ13C to magnetostratigraphy and biostratigraphy In addition to biostratigraphy, magnetostratigraphy can provide an important independent control for global δ13C correlations, especially for the Late Cretaceous period after the Cretaceous Normal-Polarity Super Chron C34. Detailed microfossil biostratigraphy together with good magnetostratigraphic data are available for Bottaccione/Contessa and for the Campanian and Maastrichtian from Site 762C, and in part for Tercis and Lägerdorf-Kronsmoor-Hemmoor. There is good agreement of magneto- and chemostratigraphy in the Upper Campanian and Maastrichtian (dashed lines in Fig. 6) with the only exception at the top of Chron C32r at Lägerdorf-Kronsmoor-Hemmoor. However, the interval of Chron C32 has been re-interpreted in Lewy and Odin (2001), and the top of Chron C32r might actually be slightly deeper (in the Belemnitalla langei Zone) according to the previous interpretation by Hansen et al. (1992), which would be more consistent with the δ13C correlations (Fig. 6). A δ13C maximum (“Exmouth Plateau Event” of Thibault et al., 2012b) occurs in the lower part of Chron C30n at Bottaccione and Site 762C. The Mid-Maastrichtian Event spans most of Chron C31n, and the broad Lower Maastrichtian δ13C minimum is consistently found in Chron C32n1r through lower C31r with a subsequent δ13C rise in the upper part of Chron C31r. The Campanian/Maastrichtian boundary as defined at Tercis correlates to the uppermost part of Chron C32n2n and lies at the base of the broad Lower Maastrichtian δ13C minimum. This level is found within the planktic foraminifer Gansseria gansseri Zone at Bottaccione and Site 1210B, and above the FO of Belemnella lanceolata in the English Chalk and in Lägerdorf-Kronsmoor-Hemmoor, consistent with correlations of Voigt et al. (2012) but different from the assumed correlations of Odin and Lamaurelle (2001). At Contessa and Site 762C, two minor δ13C minima are observed just below and above Chron C32r that might also be recognizable at similar levels in the other sections in Fig. 6. The distinct minimum of the LCE consistently occurs in the upper part of Chron C33n. Some inconsistency is found for the top of Chron C33r, which occurs at a δ13C peak at Bottaccione and at a δ13C minimum at Lägerdorf-Kronsmoor-Hemmoor. With respect to nannofossils and sediment thickness, the top of Chron C33r appears to be at a lower level at Bottaccione than at LägerdorfKronsmoor-Hemmoor (Fig. 6). However, correlation of this interval

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between these two sections is complicated by the use of different biostratigraphic schemes, the absence of a continuous isotope record at Bottaccione and the generally low δ13C variability and comparably poor data coverage during Chron C33r (Fig. 7). The top of Chron C34n is found at 222 m (scale A in Fig. 5) in Bottaccione, very close to the top of the D. asymetrica Zone with the last occurrence of the species at 221.4 m (Petrizzo et al., 2011). The SCBE in Bottaccione is composed of a double spiked broad interval of elevated carbon isotopes. A similar structure of the event is also seen at Lägerdorf-Kronsmoor-Hemmoor, whereby the nannofossil Broinsonia parca parca appears within the second maximum in both sections (Fig. 5). The last species of the planktonic foraminifer Sigalia in Guru occurs above the SCBE in the Middle Campanian (Fig. 7), consistent with observations from northern Spain (Lamolda et al., 2007). The criteria for defining the Santonian/Campanian boundary have not been fixed to date, but it has been suggested to use the extinction of the crinoid Marsupites testudinarius (Gale et al., 1995, 2008). This would place the boundary at the first peak of the SCBE in the English Chalk (Jarvis et al., 2006) and at the second peak of the SCBE in the Waxahachie Dam Spillway section (Texas, Gale et al., 2008) and in the Fizeşti section (Romania, Melinte-Dobrinescu and Bojar, 2009). The two latter works show that the LO of M. testudinarius is below the LO of D. asymetrica, as was also reported by Lamolda et al. (2007) from northern Spain. In accordance with these results, the SCBE at Bottaccione is placed at a lower level (top of D. asymetrica Zone and top of Chron C34n, consistent with Premoli Silva and Sliter (1995) and other works on the section) than in Jarvis et al. (2006), where it was interpreted to lie in the top of the Globotruncanita elevata Zone. Note that in Fig. 14 of Jarvis et al. (2006), the bio- and magnetostratigraphic zones are misplaced by ~6 m as compared to Fig. 10 in Jenkyns et al. (1994). The new interpretation of the SCBE in Bottaccione results in a discrepancy in magnetostratigraphy between Bottaccione and the English Chalk as shown in Jarvis et al. (2006), where the top of Chron C34n lies well below the SCBE, in the lower Uintacrinus socialis Zone (Gale et al., 1995; Montgomery et al., 1998). The magnetostratigraphy published for Bottaccione (Alvarez et al., 1977) was later confirmed in several sections (e.g. Chan et al., 1985). In contrast, the sediments of the English Chalk with their high carbonate content were long regarded as too weakly magnetic for reliable measurement, and magnetostratigraphic interpretation for the English Chalk is inconsistent among workers. Hampton et al. (2007) note that the exact position of the top of Chron C34n is uncertain, considered to be either in the Santonian (Montgomery et al., 1998) or in the Lower Campanian, close to the FO of B. parca parca (Barchi, 1995), basically bracketing the position of this reversal in Bottaccione.

3.9. Testing isochroneity of biostratigraphic index taxa: planktic foraminifera Correlation of δ13C shifts has to be based on biostratigraphic and/or magnetostratigraphic data, but temporal resolution of global correlation can be considerably enhanced through an iterative combination of these methods as described in Section 2. The fact that many δ13C excursions

Fig. 7. Summary of global δ13C correlations, magneto- and biostratigraphy for the Late Cretaceous (post-OAE2) and planktic foraminifer zonal boundaries. Global correlation of δ13C data from twenty Upper Cretaceous records from the Boreal Realm, Tethys, Western Interior, Indian Ocean and Pacific Ocean. Data sources are given in Section 2.1 and Table 1. The δ13Ccarb data are from bulk sediments except for Site 463 (~150 m) that shows the difference between δ13C of planktic and benthic foraminifera. The Berthoud State 3 record (~90 m, Western Interior, re-drawn from Locklair et al., 2011) and the Maastrichtian part from Guru represent the 3-point moving average. All curves are plotted to the same horizontal scale and are aligned according to the correlations in Figs. 3–6; stage boundaries were used for alignment of the Berthoud State 3 record. Dark and light gray shadings mark positive and negative isotope shifts, respectively. Curves are shown relative to the time scale for the English Chalk from Jarvis et al. (2006; based on GTS 2004; left). Respective GTS 2012 ages are shown for the combined Bottaccione/Contessa record (right) with varying scales between age tie-points (black diamonds). The age model from Voigt et al. (2012; based on GTS 2012 ages for magnetic polarity Chrons) is used above Chron C33n; additional tie-points are GTS 2012 ages for the base and top of Chrons C33n and C33r, and GTS 2012 ages for ammonite zones (correlated from English Chalk) during Chron C34n (Table 2). Note that the Turonian is shown at expanded scale for more detail. Arrows with asterisks at Tercis and El Kef indicate alternative interpretations for correlating the upper section part (Maastrichtian and post-LCE, respectively; arrows point to position of section top); similarity of version * with the respective part of Site 762C is shown in gray next to the latter. Colored lines indicate positions of planktic foraminifer zonal boundaries in the sections (short, solid lines: varying position relative to δ13C curve; dashed lines: similar position relative to δ13C curve) or where they project to the macrofossil-based records (dotted). Note isochroneity of most zonal boundaries during the Turonian through Santonian long-term high in global sea-level and diachroneity during falling sea-level in the Late Campanian and Maastrichtian. References: 1 = Gradstein et al. (1994); 2 = Ogg et al. (2004) and Huber et al. (2008); 3 = Petrizzo (2000); 4 = Petrizzo et al. (2011); 5 = Voigt et al. (2012) and Ogg et al. (2012). Abbreviations: Salzg. = Salzgitter-Salder, Lägerdorf-Kron.-Hem. = Lägerdorf-Kronsmoor-Hemmoor, Ha.-O. = Halle-Oerlinghausen; see also Appendices A to C.

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can be correlated globally (Fig. 7) offers the opportunity (1) to test for isochroneity of biostratigraphic datum levels, (2) to better assess the age of zonal boundaries, and (3) to compare and link different biostratigraphic schemes such as macro- and microfossil zonations. Planktic foraminifer biostratigraphic data are available for most of the compared sections, and the exact positions of these biozones relative to the isotope shifts disclose whether or not the occurrence of marker species is isochronous. The colored dashed lines in Fig. 7 indicate the position of planktic foraminifer zonal boundaries in the respective sections and where they project to sections that are based on nannoor macrofossils (dotted). For comparison, arrows with the respective color (next to time scales) indicate the ages for these zones as given in the GTS 2012 and in older time scales and other sources (Gradstein et al., 1994; Petrizzo, 2000; Ogg et al., 2004; Huber et al., 2008). The graph shows that in the Turonian through Middle Campanian, most of these tops and bases appear to be fairly isochronous (up to and including the base R. calcarata Zone) and likely represent reliable markers for correlation. Exceptions are the base H. helvetica and base D. concavata Zone. In contrast, all foraminiferal marker species in the Upper Campanian and Maastrichtian seem to be more or less diachronous. The problem of diachronous first and last occurrences of biostratigraphic index taxa during this period has been noted in many studies (e.g. Huber, 1990; Schönfeld and Burnett, 1991; Petrizzo, 2003; Huber et al., 2008; Thibault et al., 2010; Petrizzo et al., 2011; Voigt et al., 2012). The reasons for species diachroneity can include differences in biogeography as well as taxonomic and/or preservational effects such as different species concepts among workers, rare occurrence of marker species, poor fossil preservation and reworking. It is interesting to note the parallels in the general patterns of biostratigraphy and δ13C correlations that become apparent from Fig. 7: largely isochronous planktic foraminifera datum levels in the Turonian through Middle Campanian and their diachroneity in the Upper Campanian through Maastrichtian, and great similarity of δ13C records in the lower half of the Upper Cretaceous (13 sections), especially in the Turonian, and more regional variability in the Upper Campanian and Maastrichtian (11 sections). A reasonable explanation for both trends can be seen in the long-term development of global sea-level and temperature. During the Cretaceous, global sea-level was highest in the Cenomanian through Middle Campanian and decreased in the Late Campanian and Maastrichtian (e.g. Haq et al., 1987; Leckie, 2009). Similarly, the warmest temperatures with lowest latitudinal gradients were reconstructed for the mid-Cretaceous, followed by decreasing temperatures especially during the Campanian and Maastrichtian (Jenkyns et al., 2004; Cramer et al., 2011; MacLeod et al., 2011). During lower sea-level and cooler climate with larger latitudinal gradients, enhanced differentiation of facies and environments leads to larger regional differences in local factors that influence the δ13C and favor faunal provincialism. From this interpretation it follows that much of the observed diachroneity in planktic foraminifer key species in the Campanian and Maastrichtian reflects true differences in biogeographic distribution, while taxonomic and preservational/ abundance effects appear to be the main reasons for diachroneity of some of the mid-Cretaceous marker species. Indication thereof and a more detailed discussion on some of the planktic foraminifer index taxa are given in the following. The base of the H. helvetica Zone is found at similar positions relative to the δ13C curves (just after the c2 peak) in most of the compared sections, but occurs deeper at Wunstorf and Gubbio S2 (close to the Holywell Event, Figs. 3 and 7). At Gubbio, rare specimens of H. helvetica are found at about 2 m above the “Bonarelli” black shale horizon (Petrizzo written communication) at the end of the OAE 2 δ13C excursion (Tsikos et al., 2004). An earlier FO of H. helvetica was also reported from the English Chalk at Eastbourne within the OAE 2 between the Band C-peak (Pearce et al., 2009). A diachronous FO of H. helvetica has further been suggested by Desmares et al. (2007) and Robaszynski et al. (2010). This diachroneity appears to mainly involve differences

in taxonomic concepts between authors regarding the gradual evolution of H. helvetica from H. praehelvetica, the sporadic occurrence of the species at the base of the zone and its absence from higher latitude and nearshore settings (for a comprehensive review see Huber and Petrizzo, in press). At the GSSP site for the base Turonian at Pueblo (USA), where bentonites allow for absolute dating of the sequence, the identified position for the FO of H. helvetica varies among different workers: (1) at the Cenomanian/Turonian boundary in the middle of the OAE 2 δ13C excursion (Morel, 1998; Desmares et al., 2007), (2) in the upper W. devonense Zone at the end of OAE 2 (Eicher and Diner, 1985 in Caron et al., 2006), (3) in the lower part of the M. nodosoides Zone just after the OAE 2 (Caron et al., 2006). Huber and Petrizzo (in press) follow the concept of Caron et al. (2006) and give an age of 93.52 Ma for the FO of H. helvetica, based on the astrochronology and radioisotope age model of Meyers et al. (2012) and the GTS 2012 (Ogg et al., 2012). The top H. helvetica Zone consistently falls on the decline just after the Pewsey Event (Middle/Late Turonian) in all sections covering this interval (Fig. 7), which projects closer to the age given in the 1994 time scale (Gradstein et al., 1994) than in the GTS 2004 and 2012 that both suggests an older date. According to Wiese (2010), the LO of H. helvetica might be a little higher in the German Burgberg section, falling between the Caburn and the Bridgewick Event in the Late Turonian, but reworking could explain this exceptionally high LO because the extinction of the species is an abrupt event that appears to be reliable for global correlation (Huber and Petrizzo, in press). The δ13C correlations with the English Chalk suggest that the LO of H. helvetica is close to the base S. neptuni ammonite Zone that has an age of 90.86 Ma in the GTS 2012. Based on this age and the age of 93.52 Ma for the base of the H. helvetica Zone (Huber and Petrizzo, in press), the duration of the zone is ~2.6 Myr, which is considerably longer than the 0.53 Myr given in the GTS 2012 (Ogg et al., 2012) and also exceeds the minimum duration of 0.75 Ma given in Huber and Petrizzo (in press) by a factor of 3.5. Dicarinella concavata has its FO at the base Upper Turonian in Guru while in Italy it first appears more than one million years later, near the minimum of the Navigation Event and close to the Turonian/ Coniacian boundary (Fig. 7). These discrepancies likely reflect the rarity of the species in the Late Turonian while it is usually more abundant and consistently present in the Coniacian (e.g. Wendler et al., 2011b). The range of the species is therefore often indicated with a dashed line from the top Middle Turonian and with a solid line from the latest Turonian (e.g. Robaszynski and Caron, 1995; Ogg et al., 2008). Based on the current data set, the base and top of the D. asymetrica Zone appear to represent reliable biostratigraphic levels, and their projected ages are close to the respective ages in the GTS 2012 (Fig. 7). It should, however, be noted that relatively few of the compared sections cover the Santonian through Lower Campanian, and that this interval is condensed at Guru. Globotruncana ventricosa first occurs in the lower part of the D. asymetrica Zone on the Exmouth Plateau (Petrizzo, 2000), but has its FO in the Early Campanian in the western Tethys and in Guru, about 5–6 Myr later. In a recent paper Petrizzo et al. (2011) discuss the reliability of G. ventricosa as a zonal marker, pointing out discrepancies in taxonomic concepts as well as the diachronous appearance of the species and suggesting to replace the zone by the Contusotruncana plummerae Zone. At Guru, the FO of C. plummerae lies above the FO of R. calcarata, which means the C. plummerae Zone is not applicable at Guru based on the current data set. Further studies on the range of C. plummerae will have to consolidate the applicability and isochroneity of this zone. The R. calcarata Zone is a well-defined zone that is globally used for biostratigraphic correlation and is assumed to have a short duration of less than 0.5 Myr (Ogg et al., 2012). The zone is defined by the total range of the nominate taxon, which is considered a reliable and easily recognizable marker by most workers (e.g. Premoli Silva and Sliter, 1995; Huber et al., 2008). Comparison of the GTS 2012 ages for the top and base of the R. calcarata Zone with their position relative to the δ13C curves in Fig. 7 reveals that (1) the base of the zone is fairly

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isochronous, (2) the top of the zone is diachronous and (3) the duration appears to be longer than given in the GTS 2012. As discussed in recent papers by Wagreich et al. (2012) and Voigt et al. (2012), there is considerable uncertainty about the age and duration of the R. calcarata Zone. Reasons include the absence of direct geochronological dating for this zone, uncertainties in identification of the top of the zone in some sections and conflicting results with respect to the position of the zone top relative to the LCE. At Guru, the R. calcarata Zone has been placed entirely below the LCE with respect to the consistent occurrence of the species, while a single occurrence much higher in the section (just above the LCE) was considered as reworked (Wendler et al., 2011b). A similar situation is found at Tercis where the top of the zone was identified just below the LCE despite two single specimens that were found above the LCE (Ion and Odin, 2001; Voigt et al., 2012). Accordingly, the thickness of the R. calcarata Zone at Tercis would be either 22.5 m or 80 m, resulting in an estimated duration of 0.9 Myr and 3.2 Myr, respectively (Odin et al., 2001), based on the estimated average AR of 2.5 cm/kyr (Odin and Amorosi, 2001). Jarvis et al. (2002) show the R. calcarata Zone at El Kef below the LCE. However, their biostratigraphic ranges are based on a preliminary study of the planktic forams at El Kef and on lithostratigraphic correlations to Kalaat Senan, where more detailed biostratigraphy was available. The ranges of R. calcarata at El Kef are therefore not well constrained. This is also expressed in the different depth for the FO of the species at El Kef and the base of the zone between Figs. 2 and 3 in Jarvis et al. (2002), which is shown as uncertainty in Fig. 6 here. At Postalm, the R. calcarata Zone lies entirely below the LCE, whereas the top of the zone is slightly above the LCE at Contessa, and well above the LCE at Shatsky Rise, although identification of this isotopic event is not very clear at Site 1210B. At Bottaccione, the base and top of the zone are well constrained, but there seems to be a gap at the level of the LCE (Voigt et al., 2012). At Contessa, the LCE is well recorded, but there is uncertainty about the level of the FO of R. calcarata in the section. Voigt et al. (2012) state they started sampling the Contessa section at the base of the R. calcarata Zone, but they do not mark the FO of the species in their graphs. Similarly, no information on the FO of R. calcarata at Contessa is given in Gardin et al. (2012). If Voigt et al. (2012) assessed the level of the base of the zone at Contessa based on the thickness of the zone at Bottaccione and the LO of the species at Contessa, they must have underestimated the true thickness because part of the zone (at least 10 m) is missing at Bottaccione. The observed thickness of the R. calcarata Zone at Bottaccione together with the inferred missing part results in an original total thickness of the zone of at least 20 m. With an assumed duration for the zone of ~0.5 Myr, this would result in AR at Bottaccione of ~4 cm/kyr, which is four to ten times higher than AR calculated by Gardin et al. (2012) for the Campanian and Maastrichtian (mostly 0.4–1.2 cm/kyr). However, sedimentological evidence for such dramatic increase in AR was not reported. A more steady AR at Bottaccione implies a longer duration of the R. calcarata Zone of 2–3 Myr, which is closer to the duration of the zone based on the maximum range of the species at Tercis: 3.2 Myr (Odin et al., 2001) or 2–3.5 Myr as suggested from the δ13C correlations in Fig. 7. A longer range of the species of 0.9 Myr and more than 1 Myr was also suggested by Gardin et al. (2012) and Voigt et al. (2012), respectively. Extending the isotope record at Contessa down to the base of Chron C33n together with identification of the base R. calcarata Zone at this section would help to solve the problem. The best estimates probably come from astrochronological studies that indicate a duration of ~0.8 Myr for the R. calcarata Zone in Austria (Wagreich et al., 2012) and Tunisia (Hennebert et al., 2009; Robaszynski and Mzoughi, 2010). This duration is also favored in Odin (2010). However, with respect to the data from Contessa, it appears possible that the total range of R. calcarata ends above the LCE and that the few specimens found at this level in Guru and Tercis are not reworked. The LO of R. calcarata below the LCE in some sections could either reflect true diachroneity or rarity of the species towards the top of the zone (as suggested from the longer range at Guru and Tercis)

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and/or preservational effects. At any rate, the base of the zone obviously provides higher precision for biostratigraphic correlation than the top of the zone. Other planktic foraminifera that are commonly used for biostratigraphy in the Upper Campanian include Globotruncanella havanensis, Rugoglobigerina hexacamerata, Globotruncana aegyptiaca and G. gansseri, all showing a diachronous FO with respect to magneto- and chemostratigraphy (Figs. 6 and 7). The FO of G. havanensis lies just below the LCE at El Kef, Guru and Bottaccione, at the base of the LCE at Tercis, in the middle of the LCE at Contessa and above the LCE at Site 1210B, and it can occur below or above the LO of R. calcarata (Fig. 7). The GTS 2012 age for the FO of G. havanensis is 75.94 Ma, which is closer to the respective level in Fig. 7 (~76.6 Ma) than the age of 75.36 Ma suggested by Petrizzo et al. (2011). At Guru and El Kef, R. hexacamerata first occurs below the LCE but much later at Site 463 and Contessa. The FO of G. aegyptiaca is below the LCE at Guru, within the LCE at Site 463 and above the LCE at Site 1210B and Contessa. The FO of G. gansseri lies at the base of the LCE at El Kef and well above the LCE in Guru, Site 463, Site 1210B and Contessa. Inconsistent position of foraminiferal index species relative to the isotope curves are also observed for the Maastrichtian. The FO of Contusotruncana contusa is in the lowermost Maastrichtian at Guru and Tercis, slightly higher at Site 463, in the upper part of the Lower Maastrichtian at Contessa and in the Upper Maastrichtian in the Pacific and Indian Ocean Sites. Similar diachroneity in FO is found for Racemiguembelina fructicosa and Abathomphalus mayaroensis, with a reverse order of FO at Site 762C. The maximum difference in FO of the Campanian and Maastrichtian marker species among the compared sections (based on the GTS 2012 ages in Fig. 7) is: 5.6 Myr for R. hexacamerata, 1.5–2.3 Myr for G. havanensis (uncertainty due to gap at LCE in Bottaccione), 3.6 Myr for G. aegyptiaca, 3.3 Myr for G. gansseri, 4.4 Myr for C. contusa, 2.5 Ma for R. fructicosa and 1.7 Myr for A. mayaroensis. These results confirm previous evidence for diachroneity of these species (e.g. Odin, 2001; Odin and Lamaurelle, 2001; Petrizzo, 2003). The obvious diachroneity of planktic foraminiferal marker species in the Late Campanian and Maastrichtian (and partly in the Turonian) should stimulate future research to identify additional foraminiferal bio-events in the Late Cretaceous. For example, it was proposed to use the FO of Pseudoguembelina palpebra to identify the Campanian/Maastrichtian boundary at Zumaia (Pérez-Rodríguez et al., 2012). At Guru, the base of consistent occurrence of Pseudotextularia nuttalli coincides with a pronounced negative δ13C shift (Navigation Event) that correlates to the macrofossil-based Turonian/Coniacian boundary. If found at a similar level in other sections, this species could serve as an important biostratigraphic marker for the top of the Turonian.

3.10. Testing isochroneity of biostratigraphic index taxa: nannofossils Some first and last occurrences of nannofossil key taxa and nannofossil zonal boundaries are indicated relative to the δ13C curves in Figs. 3 to 6. While some of these datum levels appear to be fairly isochronous such as the FO of B. parca parca or the LO of Uniplanarius trifidus, there is considerable diachroneity for many species with respect to magneto- and chemostratigraphy. Because of their small size, these fossils are especially prone to preservational effects and reworking, and rarity of a species and differences in taxonomic concepts among workers might add to the observed discrepancies in nannofossil datum levels. However, similar to the planktic foraminifera, true diachroneity is certainly involved for some species especially in the Campanian and Maastrichtian, as was e.g. demonstrated for Micula murus (Thibault et al., 2010). Problems with correlating nannofossil datum levels in the Upper Campanian and Maastrichtian are also acknowledged in a recent paper by PérezRodríguez et al. (2012). In the following, some of the nannofossil index taxa are discussed in more detail.

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The FO of U. trifidus varies slightly among the sections with respect to chemostratigraphy and the planktic foraminiferal zonation: it is at the base of the R. calcarata Zone at El Kef, Postalm and Bottaccione, below the R. calcarata Zone at Site 1210B and in the middle of that zone in Tercis (Fig. 6). The LO of Eiffellithus eximius lies at a similar level (slightly varying within Chron C32r) in five of the compared sections, but it is considerably deeper at Bottaccione, Site 1210B and El Kef. As mentioned before, biostratigraphic levels of key taxa at El Kef are largely based on lithostratigraphic correlation with Kalaat Senan (Jarvis et al., 2002) and their true position relative to the isotope record might actually be different. This could also explain the much lower position of the LO of B. parca constricta, U. trifidum and Tranolithus orionatus as compared to the other sections in Fig. 6. Correlation of the post-LCE record from El Kef as shown in Fig. 6 follows correlations with the English Chalk as proposed by Jarvis et al. (2002). Alternatively, the postLCE part at El Kef could be interpreted as Lower Maastrichtian, which would shift the LO of these three nannofossils to a similar level as in the other sections and would still be consistent with the planktic foraminifera (the Campanian/Maastrichtian boundary lies within the G. gansseri Zone), but would imply reduced AR or a hiatus in this part of the section (*** in Figs. 6 and 7). In this scenario, the interval of the CMBE would not be characterized by a distinct drop of δ13C values as in most sections, but would be similar to Sites 762C and 1210B. Chemostratigraphic correlation between Site 762C and Bottaccione/ Contessa is well constrained by magnetostratigraphy (provided correct interpretation of the magnetostratigraphic data) and shows that the LO of B. parca constricta and T. orionatus is much later (top of Chron C31r) at Site 762C than at all other sections in Fig. 6 (base of Chron C31r at Bottaccione). Based on this observation, it seems plausible to correlate the section top at Tercis to a higher level, different from the interpretation of Voigt et al. (2012) that is shown in Fig. 6 and that assumes a strong increase in Maastrichtian AR (see Section 3.2). Correlation of the section top with the level marked with ** (Fig. 6) assumes constant AR and shifts the LO of B. parca constricta to the same level as at Site 762C, while version * assumes a moderate increase in AR and results in very good comparability of δ13C trends between the Maastrichtian part from Tercis and the respective record at Site 762C (gray curve marked with * next to the record of Site 762C in Fig. 7). The LO of B. parca constricta and T. orionatus are thought to be reliable markers for biostratigraphic correlation (Voigt et al., 2012), and the LO of B. parca constricta was suggested to approximate the Campanian/Maastrichtian boundary (Pérez-Rodríguez et al., 2012). Both datum levels are close to each other in most of the compared sections but vary in their relative order (Fig. 6). The observed diachroneity in the LO of these two species is mainly caused by their higher level at Site 762C, and amounts to 1.3 Myr for the LO of T. orionatus (Figs. 6 and 7). Diachroneity in LO of B. parca constricta among the compared sections is 2.6 Myr, with the species' LO varying between the base of the Maastrichtian at Site 1210B (if not considering El Kef) and the top of Chron C31r at Site 762C. The FO of L. quadratus varies from the base of Chron C31n at Bottaccione to the base of Chron 30n at Site 762C, accounting for a diachroneity of 1.4 Myr. The observed range in the FO of M. murus is 1.6 Myr, but this is based on three sections only. Micula prinsii first occurs at the base of Chron C29r at Bottaccione and Site 762C, but slightly below that level at Site 1210B if following the correlations proposed by Voigt et al. (2012). Fig. 8 shows an alternative interpretation of the magnetostratigraphy at Site 762C (column B) and the resulting correlation with Bottaccione/Contessa. At first sight, this version looks compelling because (1) it would shift the LO of B. parca constricta and T. orionatus and the FO of Lithraphidites quadratus to a similar level as in Contessa and in the other presented sections, and it would also decrease the offset for the FO of C. contusa and R. fructicosa relative to Contessa and Site 1210B, (2) it does not require postulation of a hiatus, (3) no change in depth scale is required to align the two sections with respect to magneto- and chemostratigraphy, implying fairly constant AR or similar changes in both sections, depending on the age model, (4) isotopic

Fig. 8. Alternative interpretation of magnetostratigraphy at Site 762C. Chemostratigraphic correlation and selected nannofossil datum levels between Bottaccione/Contessa and Site 762C based on alternative interpretation (B) of magnetostratigraphy at Site 762C, showing good agreement for δ13C, nannofossils and a large part of the paleomagnetic data, but conflicting at positions marked with “!” and with cyclostratigraphic model of Thibault et al. (2012b, A). See Section 3.10 for further explanation and Appendix B for abbreviations.

shifts are very well comparable between both sections. However, there are two major problems with this interpretation: (1) it conflicts with magnetostratigraphic data at two positions (“!” in Fig. 8), postulating an additional short interval with normal polarity in Chron C31r, and the short reversal of Chron C30r would not be reflected in the data set, and (2) the duration of the Chrons based on the cyclostratigraphic model (Thibault et al., 2012b) would not be consistent with the GTS 2012. That the LO of B. parca constricta and T. orionatus is indeed higher at Site 762C than in the other sections is also supported by the fact that the FO of A. mayaroensis is below the LO of B. parca constricta at Site 762C as opposed to the other sections (Figs. 6 and 7). This example shows that chemostratigraphic correlation is far from being trivial, and that diachroneity of the biostratigraphic datum levels discussed above obviously has to be accepted. To identify the LO of a nannofossil species can be very time consuming if the species is rare in the top of its range. The LO of a species might shift upwards in a section as more time is invested to search for it, as e.g. appears to have happened with the LO of B. parca constricta at Lägerdorf-KronsmoorHemmoor (Voigt et al., 2010, 2012). On the other hand, reworking of nannofossils can never be excluded and might account for an unusually

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high LO. To enhance the applicability of nannofossils for global correlation in the Upper Cretaceous, more sections are needed with good nannofossil biostratigraphy, with detailed δ13C data and, if possible, with a good magnetostratigraphic record. Furthermore, improvement of the nannofossil zonation can be reached through agreement of nannofossil workers on discrepancies in taxonomic concepts. Nevertheless, chances are that despite these efforts a fair number of key taxa (especially in the Campanian and Maastrichtian) are not isochronous. These taxa should not be used for global correlation at high-resolution but rather for a rough stratigraphic estimate or for regional correlation. 3.11. Testing isochroneity of biostratigraphic index taxa: macrofossils Macrofossil biostratigraphy is available for the German and English sections discussed in this paper. A more detailed zonation than for the English Chalk is used for Lägerdorf-Kronsmoor-Hemmoor and for the other German sections (Figs. 3 to 6; see also Richardt and Wilmsen (2012) for Anröchte and Voigt et al. (2007) for Halle-Oerlinghausen). Despite this difference, occurrences of some key species relative to the isotope curves are well comparable between the sections. Macrofossils typically have a much lower abundance in the sediments as compared to micro- and nannofossils, which limits their biostratigraphic precision (especially at drill sites) and can explain some of the discrepancies between macrofossil boundaries among the sections with respect to chemostratigraphy. The main δ13C increase towards the OAE 2 occurs in the Metoicoceras geslinianum Zone in the German and English Chalk, while an interval of lower δ13C values commonly occurs in the middle Neocardioceras juddii Zone (between peaks A2 and B, Fig. 3). However, the base of the N. juddii Zone lies slightly above the A2-peak in the English Chalk and just below this peak in Wunstorf. Two further δ13C peaks are found: in the top N. juddii Zone (B-peak) and in the base Watinoceras devonense Zone (C-peak), before δ13C values decrease again. The pronounced δ13C minimum of the Lulworth Event (= Tu 5) consistently occurs in the top of the Mammites nodosoides Zone, at the Lower/Middle Turonian boundary (Fig. 3). In the Upper Santonian of the English Chalk and of LägerdorfKronsmoor-Hemmoor, the crinoid Uintacrinus socialis Zone (s in Fig. 5) is found at the Buckle and Hawks Brow Event, followed by the crinoid Marsupites Zone (M in Fig. 5) across the Foreness Event and at δ13C increase of the SCBE. The echinoid Offaster pilula is consistently observed at the δ13C decline after the SCBE in the Early Campanian, though it occurs slightly earlier in the English Chalk than at LägerdorfKronsmoor-Hemmoor with respect to chemostratigraphy. In both sections, the belemnite Belemnitella mucronata (B m in Fig. 6) sets in just before the BUCE, and Belemnella lanceolata (lan in Fig. 6) occurs close to the beginning of the δ13C decline of the CMBE, again slightly earlier in the English Chalk than at Lägerdorf-Kronsmoor-Hemmoor. More problematic is the stratigraphic position of the ammonite N. polyplocum Zone (pol in Fig. 6) that occurs well below the LCE and bracketing with its range the base R. calcarata Zone at El Kef and Tercis, but occurs just above the LCE at Lägerdorf-KronsmoorHemmoor. Differences in taxonomic concepts might be involved here. The ranges of Belemnella obtusa and Belemnella sumensis are similar with respect to the isotope curves in the English Chalk and in Lägerdorf-Kronsmoor-Hemmoor (Fig. 6). Macrofossils were traditionally used to define the stage boundaries in the Cretaceous. However, macrofossils are rarely applicable for biostratigraphic correlation on global scale because of: (1) provincialism and regional differences in macrofossil schemes, and (2) their lower abundance as compared to micro- and nannofossils and the resulting limitations for stratigraphic precision and for their applicability in drill cores. The traditional thinking in macrofossil terms might be part of the reason why most of the Late Cretaceous stage boundaries are still not formally ratified to date. From Fig. 7 it is evident that most macrofossils-based stage boundaries do not coincide with planktic foraminifer zonal boundaries. This highlights the problems that arise for

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global correlation of micro- and macrofossil based sections. Fig. 7 also illustrates the importance of chemostratigraphy to relate the individual biostratigraphic datum levels, which can help select useful criteria for the definition of Late Cretaceous stage boundaries. For periods of good global representation of δ13C events, it appears helpful to include chemostratigraphic aspects into the stage definition. 3.12. Towards a global Late Cretaceous δ13C stack Average δ13C curves are provided in the GTS 2012 next to the magneto- and biostratigraphy (Gradstein et al., 2012; Saltzman and Thomas, 2012). While for some epochs (e.g. the Ordovician) these curves are detailed enough to be used for chemostratigraphic correlation, the Late Cretaceous δ13C curve brings out little more than the OAE 2 and slightly elevated values in the Late Maastrichtian (Ogg et al., 2012; Saltzman and Thomas, 2012). However, there are more Upper Cretaceous δ13C shifts that seem to be reflected globally and that appear useful for global chemostratigraphic correlation, as is evident from Fig. 7. Therefore, an attempt is made in Fig. 9 to develop an average δ13C stack, based on the sections compared here. Using the same tie-points for alignment of the records as in Fig. 7, the curves are plotted onto the same scale in order to show the distribution of absolute δ13C values for each period and to visualize similarities and differences in the regional δ13C trends (Fig. 9A). Although in such a plot some of the details are hard to recognize at places where the curves are crowded (these details can be seen in Fig. 7), the graph is useful to estimate the proportion of global versus local factors reflected in a δ13C record, and to assess the average δ13C values across the Upper Cretaceous. It should be noted that there is no such thing as the global δ13C value for any given period because there are always local influences that result in more or less strong regional offsets in δ13CDIC. However, a hypothetical average δ13C stack can be useful to visualize globally synchronous shifts in δ13CDIC that can be applied for global correlation. The black curve in Fig. 9B represents such an estimated average δ13C stack for the Late Cretaceous. Note that it is an estimation based on the current data set (Fig. 9A) with an attempt to reflect characteristic trends and δ13C events that are found in all or most of the presented records because these shifts can be assumed to reflect a global signal. The curve does not represent the calculated average; this would require a more detailed alignment of the curves based on numerous additional tie-points and would imply a higher precision of the correlations than seems reasonable without improving the knowledge on global isotopic trends. The gray envelop in Fig. 9B indicates the range of absolute values in the curves shown in Fig. 9A. Of course, this envelop partly reflects the sampling situation in the data set, and additional data will certainly modify the picture. Nevertheless, for intervals with similar coverage of data the outline of the envelope and its width may help evaluate the relative importance of global factors, in addition to comparison of the individual regional isotopic trends (note, e.g., narrow envelop at Hitchwood Event). Intervals with strong similarity of patterns in all sections indicate dominance of global factors on the δ13C records during that period (Fig. 9A). Examples are the Late Turonian Hitchwood Event, the broad Middle Coniacian through Early Santonian maximum and the LCE. Local influences can be recognized by features present in only one of the sections, such as the enhanced minimum around 77 Ma in Guru. As noted before, large consistency in isotopic shifts among the sections can be seen for the Turonian through Santonian and for the main trends in the Maastrichtian. Part of the variability in δ13C values among the sections appears to reflect regional differences in trophic resources: lower δ13C values are related to meso-/eutrophic conditions, e.g. at El Kef or in the Turonian trough Early Coniacian at Guru. Sections with carbonate-dominated sedimentation under meso-/oligotrophic conditions (the boreal Chalk sections and especially the Pacific and Indian Ocean Sites) show generally higher δ13C values and less fluctuation. This can be explained by the amount of re-mineralized organic matter being higher and more variable in a eutrophic regime as compared to

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Fig. 9. Upper Cretaceous δ13C variability and global average δ13C stack. A: Compilation of the discussed δ13C records plotted on the same scale and using the vertical alignment as in Fig. 7 relative to the age-calibrated English Chalk curve (Jarvis et al., 2006, GTS 2004). The data range and coverage for each interval are shown, and similarities and differences in δ13C trends between individual records help assess the relative contribution of global versus local factors reflected the δ13C records. B: Gray envelope illustrates variability in δ13C records as shown in Fig. 9A. The GTS 2012 scale is based on ages of polarity Chrons (all Chrons above C33n as in Voigt et al. (2012) and additional ages marked with black triangles) and of ammonite zonal boundaries within Chron C34n (see Figs. 3 and 7 for more detail). The bold black curve represents an estimated average δ13C stack based on the data in Fig. 9A, visualizing the Upper Cretaceous δ13C shifts that are common to all or most of the compared records. Abbreviations: Läg.-Kr.-Hem. = Lägerdorf-Kronsmoor-Hemmoor, Salzgitter-Sald. = Salzgitter-Salder, Halle-Oerling. = Halle-Oerlinghausen, Berthoud St. 3 = Berthoud State 3 core; see also Appendices A to C.

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an oligotrophic regime. Nutrient availability further controls the relative abundance of individual groups of calcifying organisms, which influences the bulk δ13C signal and can contribute to regional δ13C offsets. These early diagenetic and biological influences seem to be stronger than the effects of latitude or bottom water age on the δ13C signal of the bulk sediment, because both types of δ13C patterns (high δ13C/low variability and low δ13C/high variability) are found in various latitudes and ocean basins (Figs. 2 and 9). 3.13. Patterns in sediment accumulation rates Based on the δ13C tie-points and their age with respect to the GTS 2012 (Fig. 7), AR were calculated for the periods between these tiepoints for the sections with a sufficiently long stratigraphic range (Fig. 10, Table 3). The graph illustrates the differences in absolute AR values among the sections, but also shows a general pattern that appears to be common to all sections: elevated values are calculated for latest Turonian/Early Coniacian; generally low AR prevail through the Santonian and Campanian before the values variably increase again in the Maastrichtian. That AR appear to be simultaneously elevated on global scale might at least in part reflect an underestimation of the time represented by this portion in the age model. Although astronomical tuning and radiometric dating are available for the Turonian and Maastrichtian, and detailed magnetostratigraphy can be performed for the Maastrichtian, considerable uncertainties still exist for the correlation between different biostratigraphic schemes, e.g. between the Western Interior ammonite zones (where absolute ages are available for the Turonian) and the Boreal and Tethyan ammonite zones. On the other hand, oceanographic and climatic factors might have had a global-scale effect on AR. For example, high AR in the Turonian and Early Coniacian could be caused by increased weathering rates due to high CO2 levels and an accelerated hydrological cycle during the midCretaceous “Supergreenhouse”, although model simulations indicate non-uniform regional responses of the hydrological cycle to climate forcing (e.g. Flögel et al., 2008). The strong Maastrichtian increase in AR at Guru certainly reflects the onset of the formation of larger benthic foraminifera shallow-water limestone that represented an efficient local carbonate factory, but other factors need to be involved at the deeper sites. Changes in ocean water chemistry combined with the continuous fast evolution of calcifying organisms (Ridgwell and Zeebe, 2005) might have caused an increase in carbonate production rates at many sites in the Maastrichtian. Additionally, the less distinct isotopic

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patterns and diachronous biostratigraphic zones enhance uncertainties in correlation among the sections, which might contribute to the variably increased AR in the Maastrichtian. Graphic correlation of the isotopic shifts, polarity chron boundaries and bioevents (FO and LO) is presented for some of the compared sections for the Turonian (Fig. 11A–C), Coniacian through Santonian (Fig. 11D–F) and Campanian through Maastrichtian (Fig. 12), allowing for more detailed study of relative changes in AR. At steady AR in both sections, the values should plot along a linear line of correlation. This is the case for many intervals in Figs. 11 and 12, and R2 values for the trendlines are mostly above 0.99, supportive of the chemostratigraphic correlations in Figs. 3 to 6. However, there are conspicuous breaks in some of the trendlines that indicate AR-changes or in some cases a hiatus, provided the δ13C correlations are correct. The latter assumption is also supported by good agreement between chemo- and magnetostratigraphic correlations, where available (Fig. 12). Slight variations in relative AR could be responsible for some undulation of the isotope data points around the trendlines (e.g. Fig. 11A), which in some section pairs appears to be related to differences in AR among individual sections within a composite record, such as in the Turonian of the English Chalk (Fig. 11B, C). A break or change in the trendline slope can be caused by an AR-change in either one or in both sections, as long as the change is not equal in both sections (in which case it would not be visible in the plot). It is therefore useful to compare a record with several other records in order to identify the cause for the change in trendline slope. A decrease in AR across the Coniacian/Santonian boundary in Tingri and Bottaccione relative to the English Chalk (Fig. 11E, F) is supported by the results in Fig. 10 that show decreasing AR across the Late Coniacian for Tingri and Bottaccione and an increase for the English Chalk. This AR decrease in Tingri and Bottaccione is not apparent in the cross-plot between the two sections (Fig. 11D), probably because it is of similar magnitude. In Tingri, this drop in AR coincides with the top of a limestone unit and a marked decrease in carbonate content at ~175 m (Wendler et al., 2009). A strong break in sediment AR at Tercis across the Campanian/ Maastrichtian boundary is implied from Fig. 12A–E, based on the correlations proposed in Voigt et al. (2012), but alternative isotopic correlations of the Maastrichtian part of Tercis are possible that require less or no change in AR (* and ** in Figs. 6 and 7). Reduced AR are indicated between ~40 and 55 m at Tercis (Fig. 12A, B, D and E), coinciding with higher carbonate content in this interval (Voigt et al., 2012). Slighly higher AR at Contessa than at Bottaccione are suggested from Fig. 12F

Fig. 10. Upper Cretaceous sediment accumulation rates. Sediment accumulation rates for intervals between δ13C tie-points (numbered triangles; see also Tables 2, 3 and Fig. 7) for selected sites, showing different values among the sections but similarity in overall patterns.

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Fig. 11. Graphic correlation—Turonian through Santonian. Graphic correlation of carbon isotope shifts (circles) among some of the compared sections. A–C: Turonian; D–F: Coniacian through Santonian. The sample positions used for the cross-plots in A–C and D–F are marked with “+” in Figs. 4 and 5, respectively. Numbers indicate slope/R2 of trendlines; triangles mark bioevents (FO and LO); dotted lines shows stage boundary; Tr. = Trunch. Note that the English Chalk Reference curve is composed of several sections with slightly different sediment accumulation rates as expressed in slope variations in B and C. For legend see Fig. 12; For abbreviations see Appendices A and B.

and G, consistent with results in Gardin et al. (2012). The large jump in isotopic and magnetostratigraphic data in panel 12G, however, mainly reflects a hiatus at ~595 m at Site 762C that has been reported before based on cycle analysis and magnetostratigraphy (= 597 amsf in Thibault et al., 2012b) and is also evident from Fig. 12H. Note that this hiatus could not be resolved with biostratigraphic data alone. The difference in slope among the Campanian and Maastrichtian data in Fig. 12I–L appears to reflect different AR between Trunch and Norfolk that represent the English Chalk. The plots in Fig. 12K and O could indicate a minor hiatus at the top of the Limestone 2 unit in Guru (368 m) that has not been recognized before, but this interpretation is not confirmed in Fig. 12 M and N. Alternatively, the break in slope could be expleined by the switch in section from Trunch to Norfolk in the English Chalk record in Fig. 12K, and by a decrease in AR at Site 762C across the Campanian/Maastrichtian boundary, which would be in line with results in Thibault et al. (2012b). The increase in AR with the onset of shallow water carbonates at ~430 m in Guru is clearly reflected in Fig. 12M and N. A slight contribution to the slope change in Fig. 12N could come from the switch from Lägerdorf-Kronsmoor to Hemmoor at 300 m in the composite record for the German Chalk, but a possible difference in AR among these section parts must be minor as evidenced from Fig. 12F and H. The plots with biostratigraphic data in Figs. 11 and 12 clearly show that—although biostratigraphy is essential for providing a frame for chemostratigraphic interpretation—in detail, few bioevents are consistent with the chemo- and magnetostratigraphic correlations. As pointed out in Paul and Lamolda (2009): though not completely independent from biostratigraphy, these simple graphic correlation plots can still be very useful for assessing the reliability of individual biostratigraphic marker species. 4. Conclusions Carbon isotope data are compiled from twenty Upper Cretaceous sections that represent various palaeo-latitudes on both hemispheres as well as different oceanic settings from the Boreal, Tethys, Western Interior, Pacific and Indian Ocean, and different diagenetic histories.

The global-scale comparison of these records reveals great similarity in secular δ13C variations despite regional differences in amplitude and absolute δ13C values. While interpretation of palaeoproductivity and assessment of absolute values in past ocean water DIC from bulk δ13C data is complicated by a complex interplay of local sedimentary, environmental, biological and diagenetic factors, relative shifts in bulk sediment δ13C appear to represent a robust tool for chemostratigraphic correlation, and in many cases are applicable to diagenetically overprinted sequences. Because δ13C correlations are usually not unequivocal they require independent control with bio- and magnetostratigraphy. In combination with these methods, δ13C stratigraphy is a powerful correlation tool for high-resolution studies in the Late Cretaceous, independent of facies, latitude or ocean basin. Limitations may arise: (1) for shallow-water settings, (2) for stratigraphically short records or sections with insufficient biostratigraphic control, (3) for sequences with an unrecognized hiatus, (4) for intervals with insufficient temporal resolution due to inappropriate sample spacing with respect to AR, (5) for sections with low carbonate and high TOC content, and (6) for intervals that reflect a dominance of local influences over the global carbon budget. The compared Upper Cretaceous sections display the highest consistency in δ13C trends among the records as well as between biostratigraphic datum levels and the δ13C records for the mid-Cretaceous “Supergreenhouse”, when long-term sea-level was highest. Global sea-level and temperatures decreased in the Late Campanian and Maastrichtian, and, although the main δ13C trends from this period are reflected in most sections, this interval is characterized by considerable diachroneity of biostratigraphic index taxa and by higher regional variability in δ13C. This observation highlights the importance of magnetostratigraphic control in the Campanian and Maastrichtian. The presented global-scale compilation of δ13C records reveals both, similarities and regional differences in isotopic signatures that can help assess the influence of global versus local factors on bulk δ13C at a given time and place for the Late Cretaceous. Alignment of the δ13C curves is based on iterative combination of bio-, magneto- and chemostratigraphy, and by

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Fig. 12. Graphic correlation—Campanian through Maastrichtian. Graphic correlation of carbon isotope shifts among six of the compared sections for the Campanian through Maastrichtian, together with bioevents (FO and LO) and magnetostratigraphic boundaries. The sample positions used for isotope cross-plotting are marked with “+” in Fig. 6. Dotted lines indicate stage boundary. Breaks in the trendlines enable identification of relative changes in sediment AR (e.g. at 106 m at Tercis, panels A–E, or at Guru with the onset of shallow water limestone at 430 m, representing increased AR, panels M and N) or of a hiatus (e.g. at 196 m at Site 762C, panels G and H). Some of these breaks might reflect differences in sediment AR among different sections of a composite record, such as in the English Chalk (panels I–L) or in the Gubbio composite (Bottaccione and Contessa, panels F and G). For abbreviations see Appendices A to C.

using the main carbon isotope events as tie-points. Through comparison of absolute values in the presented records and estimating an average value, a hypothetical Upper Cretaceous δ13C stack is proposed with an attempt to reflect the δ13C shifts that are common to all or most of the records and that are considered to represent global δ13C fluctuations in Late Cretaceous ocean water DIC. The difference in absolute δ13C values and in magnitude of δ13C shifts among the sections appear to be related to the trophic regime and the resulting sedimentary content of carbonate and organic carbon that determine primary (bio-geochemical) and secondary (early diagenetic) influences on the δ13C signal. These processes seem to be more important than latitudinal effects or bottom water age.

The δ13C correlation scheme is further used to relate the different biostratigraphic zonations and to test for isochroneity of index taxa, which contributes to identify problems with biostratigraphic or taxonomic concepts. In combination with biostratigraphy and sedimentology, the comparison of δ13C records may also provide a means to recognizing local hiatuses that are sedimentologically not obvious, and to assess their duration. Most importantly, the highresolution δ13C data allow for more precise correlations, which is essential for a better understanding of the duration and of the temporal and spatial course of oceanographic and climatic processes in the strange world of the Late Cretaceous. Increasing availability of new isotope

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records and biostratigraphic data will have to verify and improve the correlations proposed here. Their future combination with additional absolute dating and astronomical tuning has the potential for more accurate dating of biozone boundaries and can help select suitable criteria for the definition of Late Cretaceous stage boundaries.

Acknowledgments The work of many colleagues is acknowledged that form the basis for this review in numerous publications on litho-, chemo-, magnetoand biostratigraphy of the sections compiled here. I especially thank I. Jarvis, S. Voigt, N. Thibault, M. Wagreich and X. Li for providing the δ13C data from their studies and for answering questions on stratigraphic details. I am grateful to B.T. Huber and M.R. Petrizzo for discussions on planktic foraminifera biostratigraphy, and to Isabella Premoli Silva for her patience in answering questions that helped combine the two depth scales for Bottaccione. Field work in Tibet was led by H. Willems and made possible by the support of L. Ding and his group from ITP-CAS, Beijing, and H. Luo from NIGPAS-CAS, Nanjing. I thank H. Willems and Q. Zhang for additional sampling at Guru and for providing larger benthic foraminifera data and discussing the Maastrichtian biostratigraphy for Guru. Review and critical comments on early versions of the manuscript by B.T. Huber, P. deMenocal and A. Immenhauser and are highly appreciated. Constructive reviews by M. Wagreich and an anonymous reviewer helped to improve the final version of the manuscript. During the study the author was funded by the German Science Foundation (DFG Wi 725/25), the University of Bremen (Germany) and the Smithsonian Institution (USA).

Appendix A. List and explanation for abbreviations used in this paper Abbreviation

Explanation

FO LO OAE TOC DIC AR GTS GSSP pl. foram. n. d. Carbon Isotope Events: L/U S.h. Cab. Navig. Light Pt. Kingsd. Hawks B. Hav. B. SCBE BUCE BCE LCE CMBE MME KPgE Stages: Ce Tu Low./Mid./Up. Co/Coniac. S/Sant. Ca M Pg

First occurrence Last occurrence Oceanic anoxic event Total organic carbon Dissolved inorganic carbon Accumulation rate(s) Geologic time scale Global Standard Section and Point Planktic foraminifera No data Lower/Upper Southerham Event Caburn Event Navigation Event Light Point Event Kingsdown Event Hawks Brow Event Haven Brow Event Santonian/Campanian Boundary Event Base Upper Campanian Event Base Calcarata Event Late Campanian Event Campanian/Maastrichtian Boundary Event Mid-Maastrichtian Event Cretaceous/Paleogene Event Cenomanian Turonian Lower/Middle/Upper Coniacian Santonian Campanian Maastrichtian Paleogene

Appendix B. List and explanation for abbreviations of microfossil species names and biostratigraphic zones Abbreviation

Explanation

Foraminifera: arch/W. archaeo. helv/H. helvetica flan Ms Argl sigali D. primitiva conc/concavata nutt forn elev/G. elevata asym/asymetrica arca stfo vent/G. ventricosa rugosa plum/C. plummerae calc/R. calcarata fal hava hexa stua ae/aegy gans/G. gansseri cont fruc maya/A. mayaroensis O. tissoti hant Nannofossils: Qg Ee B pp Rm Ca Ug Ut B pc Hb To Lq Mm Mp

Whiteinella archaeocretacea Hetvetoglobotruncana helvetica Hedbergella flandrini Marginotruncana schneegansi Archaeglobigerina Marginotruncana sigali Dicarinella primitiva Dicarinella concavata Pseudotextularia nuttalli Contusotruncana fornicata Globotruncanita elevata Dicarinella asymetrica Globotruncana arca Globotruncanita stuartiformis Globotruncana ventricosa Globotruncana rugosa Contusotruncana plummerae Radotruncana calcarata Globotruncana falsostuarti Globotruncanella havanensis Rugoglobigerina hexacamerata Globotruncanita stuarti Globotruncana aegyptiaca Gansseria gansseri Contusotruncana contusa Racemiguembelina fructicosa Abathomphalus mayaroensis Orbitoides tissoti Plummerita hantkeninoides Quadrum gartneri Eiffellithus eximius Broinsonia parca parca Rucinolithus magnus Ceratolithoides aculeus Uniplanarius gothicus Uniplanarius trifidus Broinsonia parca constricta Heteromarginatus bugensis Tranolithus orionatus Lithraphidites quadratus Micula murus Micula prinsii

Appendix C. List and explanation for abbreviations of macrofossil species names and biostratigraphic zones Abbreviation

Explanation

Macrofossils: ges/M. geslinianum ju/N. juddii dev/W. devonense cat/F. catinus nod/M. nodosoides woo/C. woollgari nep/S. neptuni Mc coranguin. Us M pil/O. pilula G. quadrata pol muc/B. mucronata lan o/obt sum

Metoicoceras geslinianum Neocardioceras juddii Watinoceras devonense Fagesia catinus Mammites nodosoides Collignoniceras woollgari Subprionocyclus neptuni Micraster cortestudinarium Micraster coranguinum Uintacrinus socialis Marsupites Offaster pilula Gonioteuthis quadrata Nostoceras polyplocum Belemnella mucronata Belemnella lanceolata Belemnella obtusa Belemnella sumensis

I. Wendler / Earth-Science Reviews 126 (2013) 116–146 (continued) Appendix C (continued) Abbreviation Macrofossil zones (LägerdorfKronsmoor-Hemmoor): Koen in p pu cw r/wf r/g s/gr ts/gr gq li/q pil p/s senon papi c/g g/s co/s bas/spin vulgaris pol langei gri/gra lan 1 obt sum/tridens cimb fas teg/jun arg/jun d ba/d

Explanation

Volviceramus koeneni Volviceramus involutus Gonioteuthis praewestfalica Sphenoceramus pachti/Cladoceramus undulatoplicatus Micraster coranguinum/Gonioteuthis westfalica Micraster rogalae/Gonioteuthis westfalica Micraster rogalae/Gonioteuthis westfalica/ Gonioteuthis granulata Uintacrinus socialis/Gonioteuthis granulata Marsupites testudinarius/Gonioteuthis granulata Gonioteuthis granulata/Gonioteuthis quadrata Sphenoceramus lingua/Gonioteuthis quadrata Offaster pilula Offaster pilula/Galeola senonensis Galeola senonensis Galeola papillosa Echinocorys conica/Gonioteuthis quadrata gracilis Gonioteuthis quadrata gracilis/Belemnitalla mucronata/Belemnitalla senior Echinocorys conica/Belemnitalla mucronata/ Belemnitalla senior Galeola papillosa basiplana/Patagiosites stobaei/ Trachyscaphites spiniger spiniger Galerites vulgaris Nostoceras polyplocum Belemnitalla langei Micraster grimmensis/Cardiaster granulosus Belemnella lanceolata Belemnella pseudobtusa Belemnella obtusa Belemnella sumensis/Acanthoscaphites tridens Belemnella cimbrica Belemnella fastigata Spyridoceramus tegulatus/Belemnitella junior Tenuipteria argentea/Belemnitella junior Oxytoma danica/Tenuipteria argentea Tylocidaris baltica/Oxytoma danica

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