Late Cretaceous orbitally-paced carbon isotope stratigraphy from the Bottaccione Gorge (Italy)

Palaeogeography, Palaeoclimatology, Palaeoecology 379–380 (2013) 81–94

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Palaeogeography, Palaeoclimatology, Palaeoecology journal homepage: www.elsevier.com/locate/palaeo

Late Cretaceous orbitally-paced carbon isotope stratigraphy from the Bottaccione Gorge (Italy) Mario Sprovieri a,⁎, Nadia Sabatino b, Nicola Pelosi c, Sietske J. Batenburg d, Rodolfo Coccioni b, Michele Iavarone c, Salvatore Mazzola c a

Istituto per l'Ambiente Marino Costiero, Consiglio Nazionale delle Ricerche (IAMC-CNR), Capo Granitola, Via del Faro 3, 91021, Campobello di Mazara (Tp), Italy Dipartimento di Scienze della Terra, della Vita e dell'Ambiente, Università degli Studi “Carlo Bo”, Campo Scientifico “E.Mattei”, Localita' Crocicchia, 61029, Urbino, Italy Istituto per l'Ambiente Marino Costiero, Consiglio Nazionale delle Ricerche (IAMC-CNR), Napoli, Calata Porta di Massa (Interno Porto di Napoli), 80133, Napoli, Italy d SEES, University of Portsmouth, Burnaby Road, PO1 3QL, Portsmouth, United Kingdom b c

a r t i c l e

i n f o

Article history: Received 13 December 2012 Received in revised form 8 April 2013 Accepted 9 April 2013 Available online 16 April 2013 Keywords: Late Cretaceous Bottaccione section Carbon isotopes stratigraphy Long-term cycles Astronomical tuning

a b s t r a c t A refined astronomical tuning of the upper Albian−lower Campanian is proposed from the Bottaccione reference section (Gubbio, central Italy). Long-term eccentricity cycles filtered from a high-resolution δ13C record were tuned to the highly stable 405 kyr cycles of the new La2010 astronomical solution for the Earth's orbital elements. The achieved orbital tuning provides a new precise, and accurate age model for dating biostratigraphic, magnetostratigraphic and carbon isotope events through a ~23 Myr long record. Cycles of ~ 8.0, 4.7, 3.4 and ~ 2.4 Myr modulate the entire δ13C record, thus extending their detection from the Cenozoic to ~ 100 Ma and represent primary and stable long-term oscillation modes of Earth's climate–ocean system. Although an ultimate driver of these long-term periodicities is lacking, we speculate that specifically the periodicity at 4.7 Myr, represents a homologue of the present eccentricity grand-cycles, evolved by the chaotic behaviour of solar system planets during the Mesozoic. The long-term periodicities potentially reflect an unexplored expression of the low-frequency response of the carbon cycle to global biogeochemical dynamics of major nutrients, particularly phosphorus, associated with modulation of inputs to the ocean in turn triggered by high-order marine transgressions and formation of highly productive shelf seas. This very long-term eccentricity control, modulated by periodic low-energy cycles, is suggested to play a crucial role in carbon cycling, controlling a chain of climate sensitive global biogeochemical processes on the Earth. Finally, these grand-cycles provide a potential tool for geological correlation and provide a robust constraint for accurate calculation of the orbital evolution of the Solar System. © 2013 Elsevier B.V. All rights reserved.

1. Introduction Astronomical forcing is a main driver of climate change, and orbital cycles recorded in sediments provide excellent high-resolution time control to investigate processes related to Earth system dynamics (e.g., House, 1985; Bradley, 1999; Shackleton et al., 1999; Hinnov, 2000; Muller and MacDonald, 2002; Weedon, 2003). However, astronomical theory is still a focus of research. Sophisticated climate models are used to test how insolation, and in particular its distribution over latitude through the year, forces climate. Also, recent detailed investigations of high-resolution multi-proxy records provided new insights on the role played by orbital forcing on the geological records and suggest that different aspects of climate may have different driving forces, and some may even be unrelated to insolation (Raymo and Nisancioglu, 2003; Huybers and Aharonson, 2010). Different proxies within a sedimentary record can be measuring completely different aspects of ⁎ Corresponding author. E-mail address: [email protected] (M. Sprovieri). 0031-0182/$ – see front matter © 2013 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.palaeo.2013.04.006

climate (e.g., Muller and MacDonald, 1997; Karner and Muller, 2000). This stimulates to develop a deeper understanding and appropriate investigation of time series from geological records and suggests cautiousness in interpreting the sedimentary cycles as a simple linear response to orbital forcing (e.g., Huybers and Aharonson, 2010 and references therein). In this study, mid- and long-term periodic variations in the carbon isotope record from the Mesozoic deep-sea sediments of the UmbriaMarche Basin (central Italy) are analysed and discussed in detail. These pelagic sedimentary sequences are among the best stratigraphic records of Mesozoic to Paleogene climate evolution in the Tethyan region (Herbert and Fischer, 1986; Tsikos et al., 2004; Coccioni et al., 2012). In particular, new and high-resolution carbon stable-isotope data for the upper Albian–lower Campanian stratigraphic interval at the Bottaccione section are here presented. A detailed and integrated biomagnetostratigraphy combined with δ13C stratigraphy provides the opportunity to compare the Late Cretaceous Bottaccione reference section with other key records from northern Europe (Jenkyns et al., 1994; Paul et al., 1994; Mitchell et al., 1996; Voigt and Hilbrecht, 1997; Voigt

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et al., 2007; Melinte-Dobrinescu and Bojar, 2010), in particular with the English Chalk reference curve (Jarvis et al., 2006). Application of integrated methodologies of non-stationary/ non-linear signal analysis, specifically Intrinsic Mode Functions (Huang et al., 1998), the WWZ Wavelet (Foster, 1996), the “REDFIT” periodogram (Lomb, 1976; Scargle, 1982; Schulz and Mudelsee, 2002), filtering-reconstruction techniques and direct tuning with the new solution of the astronomical curve (Laskar et al., 2011), give us the opportunity to achieve a new orbital tuning for a Late Cretaceous ~23 Myr long record. Moreover, exploration of long-term cycles (~405 kyr to ~8.0 Myr) in the δ13C signal throughout the entire record offers an unprecedented chance to investigate unexplored global carbon cycle dynamics and responses to orbital forcing, biogeochemical cycles and sea level changes. 2. Geological and stratigraphic setting The Cretaceous pelagic sequence of the Umbria-Marche Basin was deposited in a complex basin and swell topography along the continental

margin of the Apulian block, which moved with Africa relative to northern Europe (Channell et al., 1979; Centamore et al., 1980). The basement of the Umbria-Marche Apennines is continental and the Upper Jurassic through Paleocene pelagic succession overlies a subsiding Triassic to Early Jurassic carbonate platform. The Cretaceous pelagic sequence of the Umbria-Marche Basin is well exposed in the classical Tethyan reference section of the Bottaccione Gorge (Gubbio area) (Coccioni, 1996). The studied section (Figs. 1 and 2) is part of the Scaglia Bianca and Scaglia Rossa Formations exposed in the Bottaccione Gorge (Arthur and Fischer, 1977; Cresta et al., 1989) and starts with a lower unit of light grey to light green limestones with bands of black chert of late Albian to early Turonian age. In the uppermost part of the Scaglia Bianca there is a distinctive ~1 m-thick level of interbedded black laminated shales and grey radiolarian sands which is known as the Bonarelli Level and is dated as late Cenomanian (Arthur and Premoli-Silva, 1982). Organiccarbon values in these shales are very high, ranging up to 30% TOC (Tsikos et al., 2004). 5 to 8 m above this carbon-rich horizon, both limestones and cherts change to pink in colour and the succession contains conspicuous clay-rich interbeds. This is the Scaglia Rossa that extends

Fig. 1. Location map of the Bottaccione section and schematic picture of stratigraphic stage boundaries identified along the succession cropping out along the road.

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from the lower Turonian into the middle Eocene. Both the Scaglia Bianca and Scaglia Rossa are dominantly composed of solution-welded calcareous nannofossils with foraminifers, chiefly planktonic, accompanied by rare benthic specimens (e.g., Premoli Silva and Sliter, 1995, 1999; Gardin et al., 2001; Tremolada, 2002; Coccioni and Luciani, 2004; Petrizzo et al., 2011). 2.1. Bio- and chronostratigraphy of the Upper Cretaceous Bottaccione section Lithologic, biostratigraphic and magnetostratigraphic informations from the Bottaccione section are synthetised in Fig. 2. In particular, the first occurrence (FO) of the planktonic foraminifer Thalmanninella globotruncanoides was selected to recognise the basal boundary of the Cenomanian Stage (Kennedy et al., 2004), located at 438.1 m at the Bottaccione section. The base of the Turonian stage is placed at the lowest occurrence of the ammonite Watinoceras devonense near the expression of the global oceanic anoxic event OAE 2 at Pueblo (Kennedy et al., 2005). Following Premoli Silva and Sliter (1995), Tremolada (2002), Tsikos et al. (2004), Kennedy et al. (2005), Caron et al. (2006) and Scopelliti et al. (2008) and on the basis of the chemo- and biostratigraphic correlation between the Gubbio section and the GSSP (Global Boundary Stratotype Section and Point) at Pueblo provided by Tsikos et al. (2004), the base of the Turonian Stage at the Bottaccione section is placed just above the FO of the calcareous nannofossil Quadrum gartneri that marks the base of the CC11 Zone of Sissingh (1977) at 487.25 m. Accordingly, the Cenomanian− Turonian boundary is placed at 487.47 m (just above the FO of the calcareous nannofossil Q. gartneri at 487.25 m) about 0.8 m above the top of the Bonarelli horizon and within the planktonic foramineral zone of Whiteinella archaeocretacea. A proposal for a composite GSSP for the base of the Coniacian Stage is in preparation and will definitively use two candidate boundary stratotypes: the Salzgitter–Salder Quarry section (SW of Hannover, Lower Saxony province, Northern Germany) and the Słupia Nadbrzeżna section in central Poland (Walaszczyk et al., 2010). The proposed marker for the base of the Coniacian Stage is the FO of the inoceramid bivalve Cremnoceramus deformis erectus that is above the FO occurrence of the ammonite Forresteria (Harleites) petrocoriensis in Europe and correlates approximately to the base of the Scaphites preventricosus ammonite Zone of the North American Western Interior basin (Cobban et al., 2006). In the most recent planktonic foraminiferal biostratigraphy (Robaszynski et al., 1990; Premoli Silva and Sliter, 1995; Robaszynski and Caron, 1995; Ogg et al., 2004, 2008, 2012), the Turonian–Coniacian boundary is placed within the Dicarinella concavata Zone. However, this species has not been recorded from the Salzgitter–Salder Quarry section and only rare specimens of this species were reported from the Slupia Nadbrzezna section (Walaszczyk et al., 2010). The calcareous nannofossil Marthasterites furcatus, with the FO defining the CC12/ CC13 zonal boundary of Sissingh (1977), appears just above the base of the Słupia Nadbrzeżna section. The calcareous nannofossil Micula decussata (M. staurophora of some authors) that marks the CC13/CC14 zonal boundary of Sissingh (1977) is absent from the reference successions. Therefore, the Turonian–Coniacian boundary falls within the planktonic foraminiferal D. concavata Zone and the calcareous nannofossil CC13 Zone of Sissingh (1977), which at the Bottaccione section, following Premoli Silva and Sliter (1995) and Tremolada (2002), span from metre levels 523.85 to 555.85 and from metre levels 524.85 to 547.85, respectively. Following Walaszczyk et al. (2010), the δ 13C carbon stable isotope Navigation Event (Jarvis et al., 2006) occurs very close to the supposed Turonian–Coniacian boundary at Salzgitter–Salder Quarry section, and is, therefore, a useful additional marker for the identification of this boundary. From all of the above, at the Bottaccione section the Turonian–Coniacian boundary is placed at metre level 526, 2 m

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above the location reported by Premoli Silva and Sliter (1995). The FO of the inoceramid Platyceramus undulatoplicatus in the Olazagutia Quarry section (northern Spain) (Lamolda, 2002; Lamolda and Paul, 2007; Lamolda et al., 2007) was selected by the Santonian Working Group as GSSP for the base of the Santonian Stage. The FO of the calcareous nannofossil Lucianorhabdus cayeuxii that marks the base of the CC16 Zone of Sissingh (1977) occurs significantly below this level, with the FO of the planktonic foraminifer Globotruncana linneiana just above the Coniacian–Santonian boundary. Following Premoli Silva and Sliter (1995) and Tremolada (2002), at the Bottaccione section the FO of L. cayeuxii and the FO of G. linneiana are recognised at metre levels 555.85 and 560.85, respectively. Accordingly, also in agreement with Premoli Silva and Sliter (1995) and Tremolada (2002), at the Bottaccione section the Coniacian– Santonian boundary is assumed to be placed at metre level 560.8, just below the FO of G. linneiana and within the CC16 Zone of Sissingh (1977). The base of the Campanian Stage is generally defined by the base of the Scaphites leei III ammonite Zone of the North American Western Interior (e.g., Cobban et al., 2006). At the Waxahachie Dam Spillway section, the Santonian–Campanian boundary, as defined by the FO of M. testudinarius, is located just below the FO of the calcareous nannofossil Broinsonia parcus parcus that marks the base of the calcareous nannofossil CC18 and NC18 Zones respectively of Sissingh (1977) and Roth (1978) and the last occurrence (LO) of D. asymetrica (Gale et al., 2008). Following Premoli Silva and Sliter (1995), Gardin et al. (2001) and Tremolada (2002), at the Bottaccione section these events lie at metre levels 595.35 and 596.25, respectively. Therefore, at the Bottaccione section the Santonian–Campanian boundary is assumed to be placed at the base of Chron C33r, that is at 596.85 m, following Alvarez et al. (1977) and Premoli Silva and Sliter (1995). According to Alvarez et al. (1977) and Premoli Silva and Sliter (1995) the top of Chron C33r lies at 618.85 m. 3. Materials and analytical methods 3.1. Carbon and oxygen isotopes The Bottaccione section was unevenly sampled with a frequency of about one sample every ~ 20 cm for the entire upper Albian– lower Campanian stratigraphic interval. A total of 1063 bulk sample stable isotope analyses were carried out with an automated continuous flow carbonate preparation GasBench II device (Spötl and Vennemann, 2003) and a ThermoElectron Delta Plus XP mass spectrometer at the IAMC-CNR (Naples) isotope geochemistry laboratory. Acidification of samples was performed at 50 °C. Every 6 samples, an internal standard (Carrara Marble with δ 18O = − 2.43‰ versus VPDB and δ 13C = 2.43‰ vs. VPDB) was run and every 30 samples the NBS19 international standard was measured. Standard deviations of carbon and oxygen isotope measures were estimated at 0.1 and 0.08‰, respectively, on the basis of ~ 200 samples measured three times. All the isotope data are reported in δ‰ versus VPDB. 3.2. Methods for spectral analysis Exploration of the highly complex and non-stationary δ 13C signals was carried out by applying the Empirical Mode Decomposition algorithm (EMD) of Huang et al. (1998) in order to decompose non-stationary, multi-component signals into a series of amplitude and frequency modulation (AM-FM) components, each with slowly varying amplitude and phase. The major advantage of EMD is that the “origin” functions are derived from the signal itself, hence the analysis is adaptive in contrast to the traditional methods where the “origin” functions are fixed as sine and cosine (e.g., for the Fourier transform like methods) and the “mother” wavelet functions for wavelet analysis. This decomposition technique is derived from the simple assumption that any

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complex signal can be decomposed into a finite and often small number of components called “Intrinsic Mode Functions” (IMF), each of them representing an embedded characteristic simple oscillation on a separated time-scale. Huang et al. (1998) defined an IMF as a function that satisfies two criteria: i) the number of local extreme points and of zero-crossings must either be equal or differ at most by one; ii) at any instant, the mean of the envelope defined by the local maxima and the envelope corresponding to the local minima must be zero. The algorithm to compute IMFs is based on the concept of “sifting” defined by Huang et al. (1998) (see Appendix 1). The decomposition result of EMD depends on the choice of two important parameters: the stopping criterion and the used interpolation technique (see Appendix 1). The result of the decomposition aims to successively and hierarchically remove the first IMFs produced from the raw signal, containing the highest frequency components, and maintaining the lowest frequency components. The last IMF, or residual, should reveal the longest-term trends in the data. The final result of this process is the creation of a bank of adaptive sub-signals whose sum produces the original signal. The δ 13C signal of the Bottaccione record and the IMF components were analysed without interpolation, keeping the original unevenly sampling intervals, with: i) “REDFIT”, an evolution of the Lomb-Scargle periodogram (Lomb, 1976; Scargle, 1982; Schulz and Mudelsee, 2002). Also, to check the appropriateness of the first-order autoregressive model, (AR(1)—linear stochastic model, where each estimated data point depends only on its own immediate past value plus a random component), the “REDFIT” equivalent of theoretical and data spectrum tests based on Monte-Carlo simulations (Schulz and Mudelsee, 2002) was run. If the r_test value falls outside the acceptance region [rcritlo, rcrithi], the null hypothesis that the spectrum is consistent with an AR(1) model is rejected. ii) Foster's (1996) weighted wavelet Z-transform (WWZ). To analyse irregularly sampled non-stationary signals (frequency changes along time) an extension of the classic wavelet formalism was adopted. Such an extension was developed by Foster (1996), which defines the WWZ as a suitable weighted projection method which re-orthogonolises the three basic functions (real and imaginary part of the Morlet wavelet and a constant) by rotating the matrix of their scalar products. The author furthermore introduces statistical F-tests to distinguish between periodic components and a noisy background signal. 4. Stable isotope stratigraphy 4.1. Diagenetic overprint The range of variability recorded in the measured carbon-isotope curve corresponds to biogenic calcite precipitated under open marine conditions during the late Cretaceous (e.g., Stoll and Schrag, 2000; Jarvis et al., 2006). Also, the carbon isotope signal shows very similar absolute values and general trends previously recorded by coeval bulk δ13C curves (e.g. Jenkyns et al., 1994; Mitchell et al., 1996; Voigt and Hilbrecht, 1997;

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Jarvis et al., 2006; Voigt et al., 2007). This is a very strong indication that the carbon isotope signature has not been considerably affected by meteoric or burial diagenesis. Conversely, the δ18O values (Supplementary Fig. S1A) scattering between −5.3‰ and −2.0‰ appear depleted by ~2–5‰ relative to diagenetically unaltered marine calcite (e.g., van de Schootbrugge et al., 2000 and references therein). The data suggest that the oxygen isotope composition of the measured samples reflects elevated temperatures during burial diagenesis and/or effects of meteoric diagenesis. Although post-depositional carbonate precipitation, resulting from the breakdown of organic matter and consequently having a very low (~−20‰) δ13C value, may influence the original marine carbon isotope record, the contribution to the δ13C record can be ignored because of the overwhelming dominance of pelagic carbonate. The δ13C vs δ18Ο cross-plot (Supplementary Fig. S1B) with a low correlation (R2 = 0.08) between the two signals supports the hypothesis of a conservative behaviour of the carbon isotope values in the studied samples. The interpretation of the bulk carbonate isotope data may also be compromised by a combination of factors, including changes in nannofossil species composition, changes in the size distribution of planktonic foraminifera, and variations in diagenetic alteration amongst lithologic cycles. However, such changes do not coincide with observable shifts or trends in the carbon isotope record, suggesting that the carbon isotope ratios reflect changes in the isotopic composition of the global oceanic carbon reservoir. Effects of diagenesis could explain the quite large δ13C variability (~0.5‰) at a decametric scale, which is present in most parts of the collected record. Long-term variability could be biased by such a short-term variability in carbon isotopes. However, the outstanding mid- to long-term resemblance with the Contessa Highway and the coeval English Chalk records (see Fig. 2) confirms the potential of the Bottaccione record to reliably capture long-eccentricity cycles. In particular, due to the long residence time of carbon in the marine environment, the carbon isotope composition of many Cretaceous deep-water records (Sprovieri et al., 2006; Voigt et al., 2007; Giorgioni et al., 2012) commonly respond only to longer term forcing mechanisms, such as the ~100 kyr and 405 kyr cycles of eccentricity modulated precession, and not to individual precessional cyclicity. Finally, the adopted sampling density is sufficient to reliably detect long-term eccentricity cycles. Actually, as demonstrated by Eyer and Bartholdi (1999) and Babu and Stoica (2010), it is possible, for an unevenly sampled signal, to properly identify periodic components with frequencies much higher than the classic Nyquist frequency calculated for time-series with the same number of uniformly sampled signals. Thus, unevenly sampled signals attenuate aliased spectral components and facilitate the identification of components beyond the Nyquist limit in the associated periodogram. Specifically, for the unevenly sampled δ13C, the associated Nyquist frequency is 5 cycle/cm for the upper part and 0.5 cycle/cm for the bottom part of Bottaccione section. These values are much higher than the 0.025 cycle/cm calculated for evenly spaced signals.

4.2. Carbon isotope stratigraphy The carbon-isotope profile generated from the Bottaccione section is presented in Fig. 2. Key markers are provided by the well known δ13C excursions: Albian/Cenomanian Boundary Event, Mid-Cenomanian Event I,

Fig. 2. Carbon-isotope stratigraphy from the Bottaccione section. High resolution carbon-isotope stratigraphy from the Bottaccione section (this study, the thick black line is a 10 point Savitzky– Golay smoothing average illustrating long term trends) compared to the δ13C curve generated from the Contessa Quarry section (Stoll and Schrag, 2000) and the age calibrated English Chalk reference curve (Jarvis et al., 2006). δ13C curves from the Bottaccione and Contessa Quarry sections are reported on the same height scale. The position of key isotope events is indicated. The lithologic log is mainly based on Premoli Silva and Sliter (1995). Members are after Coccioni (1996). The position of the Pialli (or Breistroffer) (Coccioni, 2001) and the Bonarelli Levels, which are the sedimentary expression respectively of OAE 1d and OAE 2, is shown. Planktonic foraminiferal zonation is after Premoli Silva and Sliter (1995, 1999), Coccioni and Luciani (2004), and Petrizzo et al. (2011). Calcareous nannofossil zonation is after Gardin et al. (2001) and Tremolada (2002). Agglutinated foraminiferal zonation is by Kaminski et al. (2011). Abbreviation: LCE I–III—Lower Cenomanian I–III; R. (Ps.) t.—Rotalipora (Pseudothalmanninella) ticinensis; R. (Th.) appen.—Rotalipora (Parathalmanninella) appenninica; R. (Th.) globotr.—Rotalipora (Thalmanninella) globotruncanoides; R. (Th.) r.—Rotalipora (Thalmanninella) reicheli; R. (Th.) gree.—Rotalipora greenhornensis; W. archaeocr.—Whiteinella archaeocretacea; Helvet.—Helvetoglobotruncana; Marginotr.—Marginotruncana; D.—Dicarinella; C.—Contusotruncana. This work introduces a new height scale for the Cretaceous Umbria-Marche basin including the Bottaccione section. According to this scale, the meter level zero corresponds to the base of Berrasian, and the Cretaceous–Paleogene boundary corresponds to meter level 757.45, which in the height scale after Premoli Silva and Sliter (1995) is placed at meter level 382.6.

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Cenomanian/Turonian Boundary Event, and Late Turonian events (Jarvis et al., 2006). At the base of the section (upper Albian) carbon isotopes oscillate around an average value of ~2.0‰. In the upper part of the Rotalipora (Pth.) appenninica zone a positive carbon isotope excursion occurs, with separate peaks, reaching a maximum δ13C value of ~2.6‰. The first peak coincides with the Pialli/Breistroffer Level (Coccioni, 2001). The Albian−Cenomanian stage boundary lies above the second peak (see Fig. 2). In the Cenomanian, the δ13C values oscillate between 1.4‰ and 2.6‰, with a prominent positive shift, reaching a value of 2.9‰, correlated with the Mid Cenomanian Event I of Coccioni and Galeotti (2003). The Bonarelli Level is the stratigraphic expression of the most pronounced and best studied oceanic anoxic event at the Cenomanian–Turonian boundary (OAE 2 or CTBE, Jenkyns, 1980, 1985; Schlanger et al., 1987; Tsikos et al., 2004) and consists mainly of black laminated shales alternating with radiolarian sands. The total absence of carbonate in this level precludes generation of a bulk carbonate carbon isotope curve. The Cenomanian−Turonian boundary (at metre level 487.47), located 0.8 m above the top of the Bonarelli horizon, is characterised by a positive δ13C shift from 1.9‰ to 2.8‰. In the middle Turonian the δ13C values decrease from 2.9‰ to 1.0‰ to culminate in a positive shift in the late Turonian. In the lower and middle Coniacian, the δ 13C values oscillate around a mean value of ~2.0‰. From the upper Coniacian to the early Santonian the δ13C curve shows prominent positive peaks and a δ13C maximum values of 3.4‰. Above, up to the lower Campanian, the δ13C curve shows a sequence of slight negative and positive trends with δ13C fluctuating between 1.7‰ and 2.9‰. The positive shift at the top of the section appears to coincide with the C33r−C33n limit. In Fig. 2, a comparison between the studied carbon isotope record and the δ13C curves from the Contessa Quarry and the English Chalk composite reference succession is shown. The excellent correspondence among the three carbon isotope records, in both shape and absolute values, demonstrates that the isotopic variability is reproducible in coeval sections. Nonetheless, slight differences can be observed once very short-term isotope variability is compared among the records. This could be explained with differential effects of local diagenesis affecting part of the sedimentary successions and also by imperfect sub-sampling of the collected rocks, possibly including diagenetically altered carbonate components. In the Cenomanian and Turonian stages the carbon–isotope correlation offers robust and consistent support to the Tethyan planktonic foraminiferal and Boreal macrofossil biozonations (Tsikos et al., 2004; Gale et al., 2005; Ogg et al., 2008). In the English Chalk composite record and elsewhere the CTBE is characterised by a wide δ13Ccarb excursion (+2‰) which is not detectable in the Bottaccione section since the Bonarelli Level is carbonate free. δ13C values above the Bonarelli Level have an average value of 2.6‰ and do not show the sharply falling values seen in England (see Fig. 2). The values remain low in the middle Turonian up to the positive shift of the Late Turonian Events, similar in magnitude in all the three correlated sections. The δ13C minimum at the Turonian− Coniacian boundary (Navigation Event) offers a crucial correlation for the different Tethyan and Boreal sedimentary records. From the upper Coniacian to the middle Santonian, a broad positive δ13C shift, with four separate peaks, is more strongly developed in the two Tethyan, Bottaccione and Contessa Quarry, successions than in the English Chalk, showing an excursion of ~1‰. This could reflect the hemipelagic setting of the English Chalk sections where signals could be damped by diagenetic processes. However, above the δ13C minimum, defining the East Cliff Event, the correlation, in terms of carbon isotope stratigraphy and integrated biostratigraphy, between Tethyan and Boreal records becomes more problematic. Assuming a synchronous nature of carbon-isotope events among the different sedimentary records and considering the lower sampling resolution from the English Chalk (Jarvis et al., 2006) the proposed integrated stratigraphy from the Late Cretaceous Bottaccione succession offers a good reference record for the Tethyan realm.

5. Orbital tuning of the Bottaccione section 5.1. Cenomanian/Turonian–Campanian interval: frequency analysis and lithologic hierarchical pattern organisation The IMF spectra calculated for the δ 13C signal (Supplementary Figs. S2, S3) show that the first mode is characterised by the highest frequency values of the signal and is generally associated with a significant component of noise. The IMF2 component documents short-term oscillations, about 234–265 cm long, which do not correspond to any kind of lithologic expression/response. On the other hand, the IMF3 records δ 13C variability of 344–488 cm (Fig. 3) that definitively corresponds to a hierarchical organisation of thick-thin limestone beds throughout the stratigraphic interval. In particular, to get a robust interpretation of this frequency spectrum, we explored the sedimentary record in a number of “lithologic windows”, one for each stratigraphic stage, which appear organised in regular alternations of thick and thin limestone beds, sometimes intercalated by very thin marly beds and with thicker limestone beds generally lacking intercalation of marls (Supplementary Fig. S4). The number of analysed “lithologic windows” is limited to short intervals of the record (about 20% of the entire explored section), due to difficulty in identifying other well-exposed and sufficiently long-sedimentary intervals throughout the outcrop, free of metallic nets and vegetation coverage. However, they allow to i) investigate modes of sedimentary response to orbital forcing, ii) capture changes in sedimentation rates through the section and iii) primarily constrain depth/age relationships. Following de Boer (1982) and Schwarzacher (1994), this pattern of hierarchical organisation of thick-thin limestone beds has been interpreted as the robust expression of short- and long-term eccentricity forcing. Also, the thicker limestone beds (interpreted as minima of the 405 kyr cycle) are characterised by relative higher δ 13C isotope values (Supplementary Figs. S4, S5) which suggests intervals of deposition with higher ocean productivity and more dynamic circulation. This is also supported by the common occurrence of chert layers within these thicker limestones, which are indicative of radiolarian blooms (e.g. Herbert and Fischer, 1986; Leckie, 1989; Bralower et al., 1994; Erbacher et al., 1996; Lanci et al., 2010). 5.2. Orbital forcing and δ 13C phase-relationship Previously, Herbert and D'Hondt (1990), Herbert (1997) and Zachos et al. (2001) demonstrated that short- and long-term eccentricity forcing on carbon isotope records from the Late Cretaceous to the Miocene are clearly detectable in marine records. These eccentricity paced cycles, interpreted by Pälike et al. (2006) as likely caused by a long memory of carbon in the ocean, have been considered as one component of the “ heartbeat” of the Earth' s climate system. Experiments with energy balance models demonstrated how the climate effects of eccentricity are enhanced via its modulation of precession and thus differential land–sea heating and seasonal climate cycles occur mainly at low latitudes (Short et al., 1991; Crowley et al., 1992). This primarily influences precipitation patterns and intensity that during eccentricity maxima induce a “monsoon-like” precipitation mode with intense, but short wet seasons, and prolonged dry seasons, whereas during eccentricity minima more uniform annual precipitation occurs. Accordingly, the thinner limestone beds, in the hierarchical pattern organisation of the Bottaccione section (Supplementary Fig. S4), have been interpreted as sedimentary response to eccentricity maxima, associated with i) frequent and intense water column stratification, ii) reduced yearly integrated primary productivity, and iii) relative minima in δ13C. During eccentricity minima, a relative higher primary productivity, stimulated by limited column water stratification and a good bottom water ventilation induced enhanced CaCO3 deposition, leading to thicker limestone beds often containing levels of chert nodules, and stimulated maxima in δ13C. A possible global amplifier to this mechanism could also be

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Fig. 3. Spectral analysis of IMF3 from the Cenomanian/Turonian−lower Campanian interval. A) Raw δ13C signal (the thick black line is a 10 point Savitzky-Golay smoothing average) from the upper part of the studied record (Cenomanian/Turonian to the top of the section). B) Bias-corrected “REDFIT” power spectrum of the IMF 3 extracted by EMD algorithm from the raw δ13C signal. REDFIT parameters: Number of Welch overlapped segment averaging (WOSA), with 50% overlap: n50 = 3; −6dB Bandwidth = 3.52E−07. Brown and black continuous lines indicate the 95% and 80% confidence levels, respectively, while the red continuous line indicates the AR(1) red noise level. C) Foster's WWZ Wavelet spectrum, parameter used: c = 0.005. The decay constant c defines the width of the wavelet window then the number of cycles of a given frequency expected within the window. Reasonable values of c are between 0.001 and 0.0125. The WWA (Weighted Wavelet Amplitudes) is reported in the colour chromatic scale. The brown and black contour lines indicate the 95% and 80% confidence levels. Small red boxes indicate the stratigraphic position of the “lithologic windows” throughout the Bottaccione section (see Supplementary Fig. S4) and the associated periodicity of the 405 kyr cycle in the depth domain. D) REDFIT power spectra calculated for each of the five selected intervals extracted from the raw δ13C signal.

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associated to an eccentricity controlled carbon cycle and accumulation of organic carbon not in the ocean realm but on land (Zachos et al., 2010). Once again, eccentricity minima associated with a more seasonal distribution of precipitation should be favourable for maintaining wetlands and accumulating peat, and thus increasing organic carbon burial on land. Conversely, a return to prolonged dry seasons during eccentricity maxima would reverse the trend reducing the accumulation of organic carbon, while also promoting oxidation of existing peat (or lignite) through desiccation. Thus negative δ13C excursions could be associated with the release of isotopically depleted CO2, and on the opposite, rising δ13C driven by sequestration of isotopically depleted CO2. These patterns could easily be achieved by oscillating the transfer of relatively 12C enriched carbon into a reduced carbon sediment reservoir, such as continental peats and marine sediments.

5.3. The Cenomanian/Turonian–Campanian orbital tuning With a robust phase relationship between long-eccentricity forcing and sedimentary δ 13C response established, it is crucial for the orbital tuning of the Bottaccione record, to tightly constrain changes in the age–depth profile of the long sedimentary record and identify significant variations in the sedimentation rate. Sedimentation rates calculated for each separate 405 kyr long “lithologic window” provide reliable constraints to capture changes in deposition modes and provide an opportunity to translate the dominant periodicities from the depth to the time domain. This allows to interpret the REDFIT periodogram and WWZ wavelet (Foster, 1996) calculated for IMF3 which explores the evolution of mid-term periodicities along the record and documents the 405 kyr periodicity band in the δ13C signal (Fig. 3). The derived

Fig. 4. Orbital tuning of the Cenomanian/Turonian–lower Campanian interval. Raw and 405 kyr filtered δ13C record (wavelet band pass filter frequencies: 2.94 ∗ 10−3/cm b ν486.8–517 m b 3.08 ∗ 10−3/cm; 2.60 ∗ 10−3/cm b ν517–545 m b 2.82 ∗ 10−3/cm; 1.76 ∗ 10−3/cm b ν545–575 m b 1.92 ∗ 10−3/cm; 2.20 ∗ 10−3/cm b ν575–600 m b 2.35 ∗ 10−3/cm; 3.60 ∗ 10−3/cm b ν600–625 m b 3.75 ∗ 10−3/cm) of the Bottaccione section tuned to the 405 kyr cycles (band filter: 2.4 ∗ 10−6/year b ν b 2.6 ∗ 10−6/year) of the La2010d astronomical solution (Laskar et al., 2011) from the Cenomanian–Turonian boundary to the top of the succession. Cenomanian–Turonian boundary age is from Meyers et al. (2012b).

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AR(1) process fits well with the IMF3 signal. Also, the “REDFIT” equivalent of theoretical and data spectrum runs test based on Monte-Carlo simulations (Schulz and Mudelsee, 2002) applied to the IMF3 provided a value of 2 within the acceptance interval [0, 3]. However, a direct application of IMF3 for orbital tuning of the record is problematic due to the presence of higher-frequencies that in some intervals (specifically in the Turonian–lower Coniacian, see Supplementary Fig. S2) produce an unreasonable numbers of cycles. Thus, the WWZ wavelet calculated for IMF3 (Fig. 3C) was used as a guideline to subdivide the sedimentary record in a number of intervals characterised by nearly-constant values of sedimentation rates, well-constrained by values calculated in the different lithologic windows (Supplementary Fig. S4). Power spectra, calculated for each interval (for the raw δ13C signal, Fig. 3D), were used to accurately select the frequency bands for filtering the δ 13C record in the long-term eccentricity band (Fig. 4). Then, following the proposed phase-relationship between orbital-forcing and sedimentary response, a final correlation between relative highs in the 405-kyr filtered cycles of δ13C and minima in the insolation forcing and minima in δ 13C with 405-kyr eccentricity maxima, using the new astronomical solution La2010d (Laskar et al., 2011), provides the orbital tuning of the record (Fig. 4). Currently, none of the full La2010 solutions are valid beyond 50 Ma, although the 405-eccentricity component is stable over longer time-scales (Laskar et al., 2011). This allows for astronomical tuning of our studied time-interval, but only to the extracted 405-kyr component. The La2010d solution was selected, as it remains stable the longest, and differs maximally 40 kyr in the position of 405-kyr minima with the second most stable solution, La2010a. Starting point for tuning the record was the Cenomanian–Turonian boundary level recently dated by Meyers et al. (2012b) at 93.90 ± 0.15 Myr. Cubic spline interpolation between consecutive tuned 405 kyr cycles provided an accurate age for key stratigraphic events recognised in between eccentricity extremes. The achieved orbital tuning compares well (within the ranges of associated errors) to estimates based on curve fitting between radiometric dates (Table 1) for the Coniacian−Santonian stages and is in agreement with time durations recently calculated on a cyclostratigraphic basis by Locklair and Sageman (2008) from the Niobrara Formation of the Western Interior basin, USA. 5.4. Orbital tuning of the succession from the onset of the Bonarelli Level down to the upper Albian interval Recently, several efforts to date the onset of the Bonarelli Level, based on orbital tuning of lithological and proxy records, have been reported (Mitchell et al., 2008; Lanci et al., 2010). Sageman et al. (2006) proposed a detailed orbital tuning of the OAE 2 at Pueblo and calculated a duration

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of 520–560 kyr between the Cenomanian–Turonian boundary and the base of the δ13Corg positive excursion which in the Furlo and Bottaccione sections coincides with the base of the Bonarelli Level (Tsikos et al., 2004; Mort et al., 2007; Scopelliti et al., 2008). This duration of 554 kyr compares well to the range of 450 to 550 kyr recently determined by Meyers et al. (2012a) through detailed time series analysis and recognition of obliquity forcing in the Tarfaya and Demerara Rise records. Consequently, assuming an age of 93.90 ± 0.15 Myr for the Cenomanian– Turonian boundary (Meyers et al., 2012b), we calculated a date of ~94.42–94.46 Myr for the onset of the Bonarelli Level. The IMF spectra calculated for the δ13C signal (Supplementary Figs. S6, S7) again show that the first mode is characterised by the highest frequency values of the signal and is generally associated with a significant component of noise. The IMF2 and IMF3 components document short-term oscillations, about 223–321 cm long, which do not have any kind of lithologic expression/response. On the other hand, the IMF4 records δ13C variability of 354–590 cm which corresponds to a hierarchical organisation of thick-thin limestone beds recorded along this lower part of the section (Supplementary Fig. S4) and again to the 405 kyr orbital forcing. Moreover, the WWZ Foster's wavelet (Foster, 1996) calculated for IMF4, indicate a nearly-continuous occurrence of this periodic forcing in the δ13C signal (Fig. 5). Consequently, we tuned the IMF4 record to the long-term eccentricity cycles of the La2010d insolation curve (Fig. 6) and interpolated between consecutively tuned points to precisely date all the stratigraphic events recognised in this lower part of the section (Table 1). 6. Long-term and grand-cycles During the Cretaceous, the carbon cycle was affected by major perturbations related to intense volcanic outgassing and possibly to sudden methane release (Larson and Erba, 1999; Jahren et al., 2001; Weissert and Erba, 2004; Méhay et al., 2009). However, some studies suggest that orbital forcing may have played an important role in carbon cycling (Pälike et al., 2006; Mitchell et al., 2008; Lanci et al., 2010; Batenburg et al., 2012, Giorgioni et al., 2012). The orbitally calibrated record of the Bottaccione section offers an unprecedented opportunity to explore very long-term cyclicity in the δ13C signal (Fig. 7). In particular, long-term ~8.0, 4.7, 3.4 and 2.4 Myr periodicities appear as primary pacers of the whole interval (Fig. 7). Once again, following Schulz and Mudelsee (2002) the results of spectral analysis for these four low frequency bands, due to their extremely reduced associated variance, were validated by Monte Carlo simulations. Specifically, the “REDFIT” equivalent of theoretical and data spectrum runs test based on Monte-Carlo simulations provided values that confirm statistical

Table 1 Ages and durations of the Cenomanian–Santonian stratigraphic intervals at the Bottaccione section compared to data reported in literature. Radiometry Obradovich (1993) Boundary ages (Ma) Santonian–Campanian Coniacian–Santonian Turonian–Coniacian Cenomanian–Turonian Albian–Cenomanian

83.5 86.3 88.7 93.3 98.5

Duration (my) Santonian Coniacian Turonian Cenomanian

2.8 2.4 4.6 5.2

Magnetostratigraphic Zones C33R a

± ± ± ± ±

Orbital Cande and Kent (1995)

0.5 0.5 0.5 0.2 0.5

3.9

C/T boundary time scale from Meyers et al. (2012b).

Palmer and Geissman (1999)

Ogg et al. (2004)

Ogg et al. (2008)

Ogg et al. (2012)

83.5 ± 1 85.8 ± 1 89 ± 1 93.5 ± 4 99 ± 1

83.5 85.8 89.3 93.6 99.6

83.5 85.8 88.6 93.6 99.6

83.6 ± 86.3 ± 89.8 ± 93.9 ± 100.5

2.3 3.2 4.5 5.5

2.3 3.4 4.3 6.0

2.3 2.8 5.0 6.0

2.7 3.5 4.1 6.6

3.9

3.7

3.7

± ± ± ± ±

0.7 0.7 1.0 0.8 0.9

± 0.7 ± 0.7 ± 0.8 ± 0.9

Locklair and Sageman (2008)

0.2 0.5 0.3 0.15

This study 83.06 86.00 89.46 93.90a 99.30

2.39 3.40

2.94 3.46 4.44 5.40

3.24

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Fig. 5. Spectral analysis of the interval from the base of the Bonarelli level down to the upper Albian. A) Raw δ13C signal (the thick black line is a 10 point Savitzky-Golay smoothing average). B) Bias-corrected “REDFIT” power spectrum of the IMF 4 extracted by an EMD algorithm from the raw δ13C signal. REDFIT parameters: Number of Welch overlapped segment averaging (WOSA), with 50% overlap: n50 = 3; −6dB; Bandwidth = 2.56E−04. Brown and black continuous lines indicate the 95% and 80% confidence levels, respectively, while the red continuous line indicates the AR(1) red noise level. C) Foster's WWZ Wavelet spectrum, parameter used: c = 0.005. The decay constant c defines the width of the wavelet window then the number of cycles of a given frequency expected within the window. Reasonable values of c are between 0.001 and 0.0125. The WWA (Weighted Wavelet Amplitudes) is reported in the colour chromatic scale. Brown and black contour lines indicating the 95% and 80% confidence levels. Small red boxes as in Fig. 3.

reliability for the three periodicities. Firstly, Olsen (2001) proposed that ~8, 3.4 and 1.7 Myr, detected from a depth-rank series of the Upper Triassic Newark supergroup, correspond to 4.8 and 2.4 Myr eccentricity cycles, and also Ikeda et al. (2010), in their cyclostratigraphic reconstruction of the Inuyama Triassic bedded chert sequence, speculated that the approximately 3.6 and 1.8 Myr periodicities manifested as amplitude modulation of the 405-kyr cycle and could be correlated to present-day long-term eccentricity cycles evolved by chaotic behaviour of the solar system. Time series investigation by Boulila et al. (2012) of the ~65 Myr long Cenozoic compilation δ13C data of Zachos et al. (2008) confirmed

that two periodicities around 8.1 and 4.5 Myr paced the carbon cycle as a response to orbital eccentricity as a consequence of enhanced sensitivity of marine carbon cycling to longer forcing periods. Time series analysis of the La2010d astronomical solution by Boulila et al. (2012) suggested that cycles of ~ 4.5 Myr are present in the orbital eccentricity solution and correspond to a long-term eccentricity amplitude modulation derived from the resonant argument 2(g4 − g3) − (s4 − s3) where g3, g4 are related to the precession of perihelion of the Earth and Mars, and s3, s4 are related to the

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Fig. 6. Orbital tuning of the interval from the base of the Bonarelli Level down to the upper Albian. Raw and IMF4 extracted δ13C record of the Bottaccione section, tuned to the 405 kyr cycles of the La2010d astronomical solution (Laskar et al., 2011).

precession of the nodes of those planets. Also, amplitude modulation analysis of astronomical models shows that the ~ 2.4 Myr eccentricity cycles are mainly modulated by ~ 8.1 Myr periods. However, observed phase shifts between ~ 8.1 Myr eccentricity orbital forcing and δ 13C cycles during the Cenozoic (Boulila et al., 2012) suggest that additional non-orbital processes (e.g., tectonics, sea level changes, non-linear biogeochemical dynamics, etc.) could drive δ 13C variations in a non predictable way at this long-term periodicity. A primary and nearly-permanent long-term eccentricity control modulated by periodic ~8.0 and 4.5 Myr low-energy cycles appears to play a crucial role in carbon cycling and to control a chain of climate sensitive and global biogeochemical processes on the Earth. Föllmi (1996) and later Handoh and Lenton (2003) discussed long-term 5–6 Myr cycles documented in the phosphorus signal of a composite record of the Late Cretaceous. In particular, Handoh and Lenton (2003) provided a biogeochemical model to demonstrate that periodic phosphorus burial and OAEs may be triggered by cyclic changes in phosphorus input from the continents. Transgressions could generate transient increases in phosphorus input to the ocean as well as creating highly productive shelf seas. The authors suggested that low-frequency harmonics of oscillations in the climate–ocean dynamics could be linked to the third-order

eustatic sea-level change. Accordingly, a primary long-term orbital control is suggested as crucial driver of the global biogeochemical dynamics driving the carbon cycle in the Late Cretaceous oceans at periodicities corresponding to ~8.0 and 4.5 Myr. However, although an ultimate forcing of these long periodicities is lacking and major climate perturbations, unrelated to orbital drivers of climate and triggered by intense volcanic activity and tectonics, primarily regulated the evolution of ocean/climate dynamics at such a long geological time-scale, Herbert et al. (1999) suggested that low-frequency harmonics of the Milankovitch (orbital) cycles could generate up to a 2–4 million year cycles that could trigger third-order eustatic sea-level change. This forcing frequency cascade is still of a markedly higher frequency with respect to the here reported 4.5–8.0 Myr cycles, but in principle, a higher frequency forcing (sea-level variation) could force a lower frequency response, given the right nonlinearities in the Earth system as well as an appropriate envelope curve of the higher frequency oscillation and/or stochastic forcing. A more complex model is required to see whether this is a feasible alternative hypothesis. Causal chains of variations in ocean nutrient supply and changing rates of oceanic turnover may also have played an additional role in controlling long-term carbon-isotope variations.

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Fig. 7. Long-term δ13C cycles through the whole record. A) Bias-corrected “REDFIT” power spectrum of the whole 405 kyr-cycle tuned record. REDFIT parameters: Number of Welch overlapped segment averaging (WOSA), with 50% overlap: n50 = 3; −6dB Bandwidth = 7.88E−8. Brown and black continuous lines indicate the 95% and 80% confidence levels, respectively, while the red continuous line indicates the AR(1) red noise level. B) Foster's WWZ Wavelet spectrum, parameter used: c = 0.005. The decay constant c defines the width of the wavelet window and the number of cycles of a given frequency expected within the window. Reasonable values of c are between 0.001 and 0.0125. The WWA (Weighted Wavelet Amplitudes) is reported in the colour chromatic scale. Brown and black contour lines indicating the 95% and 80% confidence levels. C) Long-term 8.0 Myr (violet dashed line, wavelet band pass filter frequencies: 1.20 ∗ 10−7/year b ν b 1.30 ∗ 10−7/year), composite (8.0 + 2.4) Myr (wavelet band pass filter frequencies: 1.20 ∗ 10−7/year b ν b 1.30 ∗ 10−7/year and 4.08 ∗ 10−7/year b ν b 4.26 ∗ 10−7/year, respectively) and 4.7 Myr cycles (wavelet band pass filter frequencies: 2.10 ∗ 10−7/year b ν b 2.20 ∗ 10−7/year) filtered from the tuned δ13C record (the thick black line is a 10 point Savitzky–Golay smoothing average). OAE indicates the deposits coeval to Oceanic Anoxic Events throughout the record as single and/or groups of black shales.

Detection of these grand-cycles during the Late Cretaceous could represent an appropriate low-frequency tool for geological correlation and a robust constraint for accurate calculation of the orbital evolution of the Solar System. Supplementary data to this article can be found online at http:// dx.doi.org/10.1016/j.palaeo.2013.04.006.

Acknowledgements We warmly thank F. Hilgen for his key and constructive comments on a preliminary version of the manuscript. We also thank the three anonymous reviewers that with their comments and suggestions definitively improved a first version of the manuscript. H.C. Jenkyns is

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thanked for providing isotope data from the English Chalk. P. Sclafani is acknowledged for English revision of the manuscript. References Alvarez, W., Arthur, M.A., Fischer, A.G., Lowrie, W., Napoleone, G., Premoli Silva, I., Roggenthen, W.M., 1977. Upper Cretaceous–Paleocene magnetic stratigraphy at Gubbio, Italy V. Type section for the Late Cretaceous–Paleocene geomagnetic reversal time-scale. Geological Society of America Bulletin 88, 383–389. Arthur, M.A., Fischer, A.G., 1977. Upper Cretaceous–Paleocene magnetic stratigraphy at Gubbio, Italy. I. Lithostratigraphy and sedimentology. Bulletin of the Geological Society of America 88, 367–371. Arthur, M.A., Premoli-Silva, I., 1982. Development of widespread organic carbon-rich strata in the Mediterranean Tethys. In: Schlanger, S.O., Cita, M.B. (Eds.), Nature and Origin of Cretaceous Carbon-rich Facies. Academic Press, London, pp. 7–54. Babu, P., Stoica, P., 2010. Spectral analysis of nonuniformly sampled data—a review. 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