EDP Sciences logo
Submit your paper
Search
Menu
Home All issues Volume 192 (2021) BSGF - Earth Sci. Bull., 192 (2021) 14 Full HTML
Open Access
Issue
BSGF - Earth Sci. Bull.
Volume 192, 2021
Article Number 14
Number of page(s) 19
DOI https://doi.org/10.1051/bsgf/2020026
Published online 30 March 2021

© L. Le Callonnec et al., Published by EDP Sciences 2021

Licence Creative CommonsThis is an Open Access article distributed under the terms of the Creative Commons Attribution License (https://creativecommons.org/licenses/by/4.0), which permits unrestricted use, distribution, and reproduction in any medium, provided the original work is properly cited.

1 Introduction

Large paleo-climatic and paleo-oceanographic disturbances occurred during the Late Cretaceous. Most of these perturbations, such as oceanic anoxic events (OAEs; Schlanger and Jenkyns, 1976; Jenkyns, 2010), are recorded by the pronounced carbon isotope excursions (CIE) recognised in marine and terrestrial environments. The largest of these perturbations is the OAE 2, which took place at the Cenomanian-Turonian transition (e.g., Tsikos et al., 2004; Grosheny et al., 2006; Erba et al., 2013; Aguado et al., 2016). This event is regionally expressed by particular facies that are rich in organic matter and have local names such as the “Black Band” in Yorkshire (England), the “Thomel level” in the Vocontian Basin (France) or the “Bonarelli level” in Umbria-Marche (Italy). In the Anglo-Paris Basin, the event is prominently expressed in the chalky succession by a dark-grey, marly interval called the “Plenus Marls”.

The associated pronounced positive CIE is synchronous throughout the Tethyan and Atlantic Oceans as well as in the epicontinental seas (Arthur et al., 1988). The detail of the δ13C profile of OAE 2 shows several distinct maxima (a, b, and c) which seem to be related to anoxic conditions in most oceanic sites (Voigt et al., 2007; Joo and Sageman, 2014; Jarvis et al., 2015). Other younger peaks, in the lowermost Turonian, are thought to be related to suboxic or oxic conditions (Takashima et al., 2009).

The aim of this work is to highlight these paleo-oceanographic events in an epicontinental sea domain, in the central part of the Paris basin, which is in an intermediate position between the English chalk domain and the Tethyan realm (SE of France and central Italy). In 1999, two deep boreholes (Craie 701-Poigny, N 48°32’6”-E 3°17’33” and Craie-702-Sainte-Colombe, N 48°32’20”-E 3°15’25”) were drilled in the chalk of the Paris Basin, near Provins (Hanot, 2000, Fig. 1). It is in this eastern part of the Paris Basin that the chalk deposits reached a maximum thickness of 700 m. Both holes recovered an apparently complete (and nearly 700 m thick succession) that extends from the Late Cenomanian to Campanian, which could potentially be used for the correlation of biostratigraphic and geochemical events between the Tethyan and Boreal realms (Mégnien and Hanot, 2000; Robaszynski et al., 2005). In the Poigny borehole, we sampled a 270 m thick interval spanning the Cenomanian through to the Turonian-Coniacian boundary.

By generating a long-term carbon isotope profile from the Early Cenomanian to the Early Coniacian, we aimed to detect δ13C events and to test their potential for intra- and interbasinal correlations. In order to place the Poigny δ13C record at a large scale, we compared it with reference sections located along a North-South transect (Fig. 1), including the Culver Cliff section in England (Paul et al., 1994; Jarvis et al., 2001; Jarvis et al., 2006) and Hyèges, Angles, Vergons and Lambruisse sections of the Vocontian Basin (Mort et al., 2007; Takashima et al., 2009; Gyawali et al., 2017; Danzelle et al., 2018; Gale et al., 2018). A comparison with the pelagic Tethyan realm (Gubbio section; Jenkyns et al., 1994) is also proposed. Thus, such transects across the European epicontinental sea and the Tethyan Ocean provide a better global understanding of paleo-environmental changes for the Cenomanian and Turonian stages. Climatic changes during the studied interval are also discussed on the basis of the δ18O record.

thumbnail Fig. 1

Late Cenomanian paleogeography of Western Europe and a part of the Tethys (modified from Philip and Floquet, 2000). A rough location of some key sections: 1. Craie 701 borehole (this study); 2. Eastbourne, Sussex, UK (Paul et al., 1994); 3. Lambruisse, Vocontian Basin, France (Takashima et al., 2009); 4. Gubbio, Umbria-Marche Basin, Italy (Tsikos et al., 2004).

2 Geological setting

2.1 The chalk in the Paris Basin

Since the 19th century, chalk from the Anglo-Paris Basin has been the subject of hundreds of stratigraphic studies (Mortimore, 1983; Pomerol, 1983; Mortimore, 2011; and references therein).

The formation of the Paris Basin began at the end of the Permian and its Mesozoic evolution was largely influenced by global plate tectonic events (e.g. the opening of the Atlantic Ocean or closing of the Tethys Ocean) and eustatic variations (e.g., Guillocheau et al., 2000). The European chalk deposits correspond to the maximum of the second first-order eustatic cycle (Hallam, 1984, 1992; Haq et al., 1987; Haq, 2014). However, throughout its history, the Paris Basin has remained an epicratonic sea bordered by the Variscan massifs: the London-Brabant Massif to the East, the Armorican Massif to the West, and the Massif Central to the South (Fig. 1).

After a major regression and the emergence of the basin at the end of the Jurassic, the first Lower Cretaceous deposits were essentially continental (Wealden facies), argillaceous and sandy due to the erosion of the surrounding continents. Marine conditions returned progressively from the East (during the Aptian) to the West (during the Albian). During the Late Cretaceous, very high sea levels allowed the invasion of this epicontinental platform by pelagic organisms, such as coccolithophoridea, leading to the homogeneous chalk deposits. Sedimentation was driven by several transgressive-regressive third-order cycles (Robaszynski et al., 1998; Lasseur, 2007; Amédro and Robaszynski, 2014), and the highest sea level occurred near the Cenomanian-Turonian boundary (Haq, 2014). Tectonic episodes, related to the regional geodynamic evolution of the NW European basins, also had an influence on the depositional systems during the Middle Cenomanian to Lower Turonian and in the Upper Turonian intervals (Deconinck et al., 1991a; Vandyckes and Bergerat, 1992; Bergerat and Vandycke, 1994; Lasseur, 2007; Mortimore, 2011). For example, more subsiding areas have been identified locally (Boulonnais, Aube) but, on the contrary, other areas show relative uplift and erosion (Pays de Caux, Normandy, Touraine).

The studied borehole (Fig. 1) is located in the centre of the Paris Basin between the North Bohemian Massif to the East and the whole range of the Armorican Massif (West) and Massif Central to the South, which is an area of several major oceanic circulations (Voigt and Wiese, 2000). Cooler water masses from the north may have flowed down along the Bohemian Massif; a north-westward shift of boreal currents is defined in the southern part of the basin and, also, in the North Atlantic water circulation along the England margin (Voigt and Wiese, 2000). The most subsiding part of the Paris Basin, during the Cenomanian and Turonian, is found near Provins (Seine-et-Marne), which was also a most open area to the pelagic Tethyan realm (Lasseur, 2007) between the north-west infra-tidal and the south-east bathyal domains throughout the Anglo-Paris Basin.

2.2 Lithology of the Craie 701-Poigny borehole

The two boreholes at Poigny and Sainte-Colombe were drilled as part of the Craie 700 programme, initiated by the Paris Basin Geologists Association and the French Compagnie Générale de Géophysique. These cores are stored at the core library of Rennes1 University (France). The initial objective of this project was to understand the variations of the seismic wave velocity in the chalk which had remained unexplained after more than 50 years of petroleum exploration (Hanot, 2000). Despite the fact that the chalk turned out to be one of the hardest rocks to core due to rapid equipment breakage caused by flints, nodular chalks and fractures (Mortimore, 2014), recovery reached 98%. From an applied geology perspective, the main result was the occurrence of a 17 m thick, dolomitic unit, located at a depth of 168 to 185 m (from the surface) in the Craie 702-Sainte-Colombe hole (Pomerol, 2000), whereas the same stratigraphic interval drilled at Poigny (Craie 701 borehole, Robaszynski, 2000; Robaszynski et al., 2005) showed the usual chalk facies. Therefore, in order to prevent diagenetic artefacts, the Craie 701 borehole was chosen for the geochemical analysis.

The Craie 701 borehole displays an apparently continuous 650 m thick chalk succession from the Early Cenomanian to the Campanian-Maastrichtian boundary (Robaszynski and Bellier, 2000). The detailed lithology (Fig. 2) was described by Barrier (2000), Robaszynski (2000) and Robaszynski et al. (2005). Eleven lithological units (LU) have been described. The oldest eight were studied here. Briefly, the facies encountered in these LU are the following, from older to younger strata:

  • LU1 (700.00 to 674.20 m): hard grey cherty limestone with phosphatic and glauconite conglomeratic levels;

  • LU2 (674.20 to 654.60 m): grey flaser limestone composed of calcispheres cemented with calcite and partly with dolomite. A 40 cm thick conglomerate marks the base of this unit;

  • LU3 (654.60 to 646.60 m): grey marly chalk. This unit can correspond to the Anglo Paris Basin Plenus Marls Formation (Jefferies, 1963), considering the highest occurrence of the benthic foraminifera Rotalipora cushmani at 651.80 m (Robaszynski et al., 2005);

  • LU4 (646.60 to 622.10 m): nodular flaser chalk;

  • LU5 (622.10 to 520.7 m): calcispheres and hard flaser chalk with frequent marl beds.
    The three marly layers, corresponding to bentonite levels, are described: the Bridgewick Marl (at 523 m), the Caburn Marl (at 550.5 m) and the Southerham Marl (at 566.70 m; Deconinck et al., 2005; Robaszynski et al., 2005; Lasseur, 2007; Mortimore, 2011);

  • LU6 (520.70 to 508.10 m): chalk with tubular flints;

  • LU7 (508.10 to 448.40 m): “marbled” and grey chalk with large bioturbation structures and some flints at the base of this unit;

  • LU8 (448.40 to 427.60 m): white chalk with hard nodules.

The observed chalk succession is comparable to those described in the Boulonnais and the Aube areas, as well as in England, especially for the Cenomanian/Turonian transition where the eight characteristic levels of Plenus Marls were observed in the Aube (Amédro et al., 1978; Amédro et al., 1979; Robaszynski et al., 1987; Amédro et al., 1997; Amédro and Robaszynski, 2000; Amédro and Robaszynski, 2008).

thumbnail Fig. 2

Carbon and oxygen isotopic profiles of the Craie 701 borehole. Lithology (bentonite levels are identified by a green dashed line) and biostratigraphy are from Robaszynski et al. (2005). Subdivisions of the δ13C profile into eight carbon isotope sequences (CIS) labelled I though VIb are also mentioned.

2.3 Biostratigraphy of the Craie 701-Poigny Borehole

A chronostratigraphic framework of the Craie 701 borehole was established for the Cenomanian-Campanian succession based on planktonic and benthic foraminifera, calcareous nannofossils and dinoflagellates (Janin, 2000; Masure, 2000; Robaszynski and Bellier, 2000; Robaszynski et al., 2000, 2005). Ammonites are relatively rare in this part of the basin. In contrast, inoceramids are abundant. Bioclasts are composed of echinoderms, inoceramids and siliceous sponges. Thus, each stage boundary is proposed on the basis of a set of biostratigraphic data and regional lithological markers recognised from outcrops in the Anglo-Paris Basin (Mortimore et al., 2001; Mortimore, 2014 and references therein; Fig. 2).

The Cenomanian-Turonian boundary should occur in the 651.8–640.3 m interval (Figs. 2 and 3). This boundary is located, from lithological arguments, above the Plenus Marls equivalent (LU3) and near the base of the nodular chalk (646.6 m). From biostratigraphic data, it could be located between the LADs (Last Appearance Datums) of R. cushmani (651.80 m) and Inoceramus pictus and the FAD (First Appearance Datum) of Mytiloides mytiloides (643.00 m) (Robaszynski et al., 2000, 2005). A Turonian age for the overlying sediments is confirmed by the first occurrences of Globorotalites sp. (609.9 m), Stensioeina granulata (526.6 m) and Reussella cf. kelleri (510.1 m).

The Turonian-Coniacian boundary is poorly constrained. Initial proposals ranged from approximately 530 m (Janin, 2000; taking the abundance of the nannofossil Eiffellithus gr. eximius) to 500 m (Robaszynski et al., 2000), then to 450 m (Robaszynski et al., 2005). Lasseur (2007) proposed to relocate this boundary to around 500 m. The lower and upper boundaries of the Turonian will be refined later (see section 5), on the basis of carbon isotope correlation.

thumbnail Fig. 3

Carbon and oxygen isotopes and carbonate content data of the Craie 701 borehole across the Cenomanian-Turonian boundary.

3 Samples and methods

The Cenomanian-Turonian interval was sampled (198 samples) from 700 m (at the base of the core) to 436 m (middle of LU8), with a sampling interval ranging from 50 cm to 3 m, depending on the lithology and stratigraphic boundaries.

Bulk samples were crushed using an agate mortar. Calcium concentration measurements were performed using a Calcimeter. The CO2 pressure generated by the reaction of hydrochloric acid (30%) and the powdered sample (100 mg) was converted to calcium carbonate content.

Carbon and oxygen stable isotopes were measured using a dual inlet system Delta V Advantage mass spectrometer, coupled with a Carbo-Kiel Device for automated CO2 preparation from carbonate samples (30 to 40 μg). The reaction was produced by adding phosphoric acid to individual samples at 70 °C. Isotopic data are reported in conventional delta (δ) notation, relative to the Vienna Pee Dee Belemnite (VPDB). An internal standard has been calibrated to the NBS-19 reference standard. Analytical uncertainties based on replication and standard analyses are ±0.05‰ for carbon isotopes and ±0.08‰ for oxygen isotopes.

4 Results

4.1 Calcium carbonate content

A comparison of the evolution of the carbonate content (% CaCO3) with respect to the LUs shows that LU1 corresponds to the lowest values (an average of 50 to 60%) with a gradual increase to 81% (Fig. 2). The boundary between LU1 and LU2 is marked by a sharp increase, up to 91%. The percentage of CaCO3 values then remain relatively stable in LU2. In LU3 (which corresponds to the Plenus Marls Formation, latest Cenomanian), the carbonate content first decreases (down to 73%) in the lower part of the unit, and then increases (from 70 to 100% in its upper part, to the base of LU4).

In the remaining interval, i.e. LU4 through LU8, the CaCO3 content remains high with a slight, long-term evolution from an average of about 92% (from the top of LU4 to LU5) to an average of 98–100% (within LU7 and LU8). This long-term increase is interrupted by three significant decreases: down to 63% at 576.2 m, just below a breccia hardground, 83% at 556.5 m, which coincides with pyritic chalk, and 80% at 532.5 m. This interval corresponds to a lower overall trend of values (down to 90%) in the upper part of LU5.

4.2 Bulk carbon isotopes

The evolution of the carbon isotope ratio (δ13C) shows high-amplitude fluctuations (from 1.6 to 5.2‰) that we used to define eight isotopic sequences, labelled CIS (Carbon Isotope Sequence) I to VI (Fig. 2).

In CIS I, δ13C values are low and very stable (1.7‰). This sequence corresponds to LU1. In CIS II, an increase of ∼ 1‰ is observed until an interval of relatively stable values (2.7‰) formed a plateau matching LU2.

CIS III corresponds to the Cenomanian-Turonian boundary interval, with a positive δ13C excursion (CIE) reaching 5.17‰, which documents the OAE 2. This interval is studied in detail by Boulila et al. (2020). Based on cyclostratigraphic interpretation, these authors have shown an exceptionally well-defined carbon isotopic event with distinct onset and end. We propose to divide this into two sub-intervals: CIS IIIa, corresponding to the increasing part of the CIE, and CIS IIIb to its decreasing part (down to 3.3‰). CIS IIIa corresponds to the main part of LU3 (Plenus Marls, Fig. 3) and CIS IIIb to the lower part of LU4. In detail (Fig. 3), the increase in CIS IIIa shows three peaks, labelled 1 to 3, and the decrease in CIS IIIb has three minor peaks, labelled 4 to 6.

CIS IV (Fig. 2), which shows a plateau of relatively high values (3.3‰), corresponds to the uppermost part of LU4 and the lower part of LU5. After this plateau, a progressive decrease down to 1.6‰ characterises the CIS V isotopic sequence ending at 536 m and corresponding to the middle part of LU5. This decrease is interrupted by two positive peaks: the first at 2.42‰ (557.15 m) and the second at 1.88‰ (540.15 m).

A small, positive excursion between 536 and 450 m (culminating at 510 m, 2.6%) defines CIS Va and b. These isotopic sequences correspond to the uppermost part of lithological sequence LU5 and sequences LU6 and LU7. The end of this excursion coincides with a low δ13C value of 1.82‰ (450.2 m), which matches the LU7-LU8 boundary.

4.3 Oxygen isotopes

In the Upper Cenomanian and basal Turonian sediments (CIS I to CIS III, Fig. 2), oxygen isotopic ratio (δ18O) is very unstable with the amplitude of variations ranging from −3.1‰ to −5.2‰. In the rest of the Turonian (up to 628 m, sequences CIS IV to VI), the values are more stable, and the δ18O variability becomes lower than 1‰. Two important and regular value plateaus could be distinguished. In the first, δ18O is around −4‰ (CIS IV and V), and in the second it is approximately −3‰ (CIS VI), with a steep transition between these two plateaus (from 535 to 515 m).

Despite the high δ18O variability in the base of the core, several detailed trends can be observed in the Cenomanian and Cenomanian-Turonian boundary intervals (Figs. 2 and 3). In the basal sequence LU1-CIS I, it is impossible to distinguish a clear evolution in the mean δ18O values. Trends become more evident from CIS II, where the mean δ18O values increase from approximately −4.8‰ to −3.2‰ (Fig. 2). The following sequence (LU3-CIS IIIa, Plenus Marls) shows a significant decrease down to −5.2‰, and are the lowest values observed for the whole of the studied succession. It should be noted that these negative values are synchronous with the positive excursion of δ13C (Figs. 2 and 3). A positive excursion occurs in the sequence CIS IIIb: values rise to −3.2‰ then decrease to around −4.8‰ near the boundary between CIS IIIb and CIS IV. At the base of CIS IV, the δ18O values gradually increase to reach 610 m, at the plateau of about −4‰ described above.

5 Discussion

5.1 Stratigraphy of the chalk succession in the Craie 701-Poigny Borehole

To analyse the sequence of Late Cenomanian and Turonian paleo-environmental changes and to test their global or regional characteristics, it is necessary to have a very precise stratigraphic framework. Due to the scarcity of biostratigraphic markers in the Paris Basin chalk, global stratigraphic correlations with Boreal and Tethyan realms are complex. Carbon isotope chemostratigraphy is a powerful tool to establish a correlation framework for the Cenomanian and Coniacian deposits (Accarie et al., 1996; Voigt and Hilbrecht, 1997; Stoll and Schrag, 2000; Amédro et al., 2005; Voigt et al., 2007; Joo and Sageman, 2014; Jarvis et al., 2015). δ13Ccarb and δ13Corg profiles during this time interval are similar for several marine and non-marine sections, implying a global response of the carbon cycle (Accarie et al., 1996; Amédro et al., 2005; Joo and Sageman, 2014; Jarvis et al., 2015).

Therefore, we refined the chronostratigraphy of the Craie 701 borehole by using bentonite levels, which have been defined as isochronous and recognised throughout the basin from England to Germany (Deconinck et al., 1991b; Wray, 1999; Vanderaveroet et al., 2000; Godet et al., 2003; Robaszynski et al., 2005). We compared the isotopic and biostratigraphic data with those of a stratigraphically well-constrained reference section with similar lithological deposits: the Culver Cliff section on the Isle of Wight (Scholle and Arthur, 1980; Paul et al., 1994; Jarvis et al., 2001; Mortimore et al., 2001). This section was chosen because the observed lithological succession completely covers the studied interval and contains five bentonite levels in the Turonian. In the following, we used the biostratigraphic, isotopic and radiochronological synthesis of this section, provided by Jarvis et al. (2006).

The δ13C fluctuations are very similar between the Culver Cliff section and the Craie 701 borehole, both in form and in absolute values (Fig. 4). At the base of the Craie 701 borehole, CIS I corresponds to the lower values observed in the Early Cenomanian in England. The first significant increase in δ13C (about +0.7‰ at 675.2 m) can be correlated with the mid-Cenomanian Event 1. The occurrence of Inoceramus crippsi and Inoceramus virgatus in the core supports this correlation (Figs. 2 and 4). Since the small peak observed at 694 m may correspond to the one at Culver Cliff at 5 m, the Albian-Cenomanian boundary should be close to the base of the core. In Italy, the Contessa Quarry section also exhibits a large positive excursion of 1‰ during the Cenomanian (Stoll and Schrag, 2000).

The δ13C plateau corresponding to CIS II is also present in the Culver Cliff curve, but its extent is greater (55 m). It corresponds to the Middle and Upper Cenomanian intervals. In detail, the end of the observed increase at 662 m and the beginning of the first plateau can be correlated with the Jukes Browne event defined in England. The first occurrence of Rotalipora spp. at 675 m allows us to confirm the Cenomanian age of these deposits.

The huge positive excursion (CIS IIIa and b), corresponding to the well-known OAE 2, is similar in both sections. It encompasses with the Plenus Marls and the Cenomanian-Turonian transition. This is consistent with the last occurrence of R. cushmani, observed between 651 and 652 m (Robaszynski and Bellier, 2000; Robaszynski et al., 2005). In detail, peaks 2 and 3 occur during the increase of δ13C (Fig. 3) and are located in the end part of and just above the Plenus Marls, respectively. They coincide with peaks a and b of the same formation, defined by Jarvis et al. (2006), at Culver Cliff (Fig. 4). The “trough interval” between these peaks is also observed in both sections.

The Cenomanian-Turonian boundary is positioned in England and in the Paris Basin (Boulonnais, Pays de Caux, Aube) above the Plenus Marls (in the Mead Marls) at the level of the first minor peak observed during the progressive decrease of the δ13C (peak c). This peak seems to correspond to the minor isotopic event labelled 4 in the Craie 701 borehole. If correct, the stage boundary would be at 644 m, i.e. 2.6 m above the lithological boundary (646.6 m). This coincides with the interval proposed by the biostratigraphic data: between the last occurrences of R. cushmani and of I. pictus (651.8 m) and the first occurrence of M. mytiloides (643 m), cf. supra. More recently, the boundary has been precisely replaced on the basis of a fine cyclostratigraphic correlation between records from Poigny and Eastbourne, together with additional calcareous nannofossil data from Poigny (Boulila et al., 2020). In the final part of the Cenomanian-Turonian δ13C excursion, minor peak 5 (638 m) should correspond to the Holywell event, defined at Culver Cliff. The same observations are made by Voigt et al. (2007) in Germany. An identical correlation between the isotopic signal and the CaCO3 trend is also observed, i.e. relatively high values for isotopic peak 1 and a drop for peaks 2 and 3.

The second isotopic plateau (IV) has its equivalent in England, in the Early Turonian and the lower part of the Middle Turonian (Fig. 4). In detail, the Lulworth and Round Down events can be recognised in the Craie 701 borehole at 618 m and 594 m, respectively. These correlations are consistent with the presence of Inoceramus cuvieri and the appearance of Globorotalites and E. gr.eximus (Robaszynski and Bellier, 2000; Robaszynski et al., 2005).

After this plateau, the progressive decrease in the δ13C, down to the lowest values, is similar in both sections but the isotopic correlations are not so evident. Nevertheless, specific lithological levels may support the proposed isotopic correlation. The breccia located at 574 m in the Craie 701 borehole can be correlated with the Ogbourne Hardground. Voigt et al. (2007) also recorded this isotopic decrease, which also coincides with a significant decrease in carbonate contents in the studied core (Fig. 2). Three other clay horizons, corresponding to bentonites (Deconinck et al., 2005; Robaszynski et al., 2005), located in the core at 566.7 m, 550.5 m and 523 m can be well correlated with those described in the Upper Turonian of Culver Cliff, respectively, as Southerham Marl (B2), Caburn Marl (B3) and Bridgewick Marl (B4). The upper part of the LU5 may be dated to the beginning of the Late Turonian. The isotopic increase observed at 510 m in the studied succession may be correlated with the Hitch Wood event from the Culver Cliff trend and the following decrease (up to the Navigation event) with the LU7-LU8 lithological boundary at 449 m. A similar increase is also observed in Italy (Stoll and Schrag, 2000). The Turonian-Coniacian boundary coincides with these low values and should be located in the Craie 701 borehole at approximately 448.5 m. This is also consistent with the position of the Shoreham Marl bentonite level (B6 at Culver Cliff), at 424.9 m in the Craie 701 succession (Robaszynski et al., 2005).

The absolute isotopic values from the Turonian deposits are consistent in both sections. This boundary location is close to that proposed (444 m) by Robaszynski et al. (2005). The biostratigraphic framework of Janin (2000), which locates the boundary at 530 m, and that of Lasseur (2007), at 500 m, cannot be retained in this case.

A comparison of the sedimentation thickness in both sections shows differences in the Cenomanian (50 m in the Craie 701 borehole versus 75 m in the Culver Cliff section) and in the Lower and Middle Turonian (70 m versus 55 m). In contrast, the thickness of the Upper Turonian deposits is very different between the two sites: about 120 m in the Craie 701 borehole and only 30 m in England. Sedimentation rate estimates, based on bentonite horizon correlations, have shown an important dilatation (up to three times greater) for the Turonian deposits from the Aube compared to the Boulonnais Series, which is seemingly an area with less subsidence (Vanderaveroet et al., 2000). The German chalk succession also shows a thicker series than the English one (Wray, 1999). The reduced thickness of the Culver cliff section can be explained by the occurrence of several hardgrounds within this interval, from the Ogbourne to the Navigation hardgrounds. A new episode of tectonic instability during the Upper Turonian controlled subsiding zones along a major NW-SE trending axis (Aube) and uplifted areas in England and Normandy (Deconinck et al., 1991a, b; Vandyckes and Bergerat, 1992; Bergerat and Vandycke, 1994; Mortimore et al., 1996; Mortimore and Pomerol, 1997; Lasseur, 2007).

thumbnail Fig. 4

Correlation of the Culver Cliff section (Paul et al., 1994; Jarvis et al., 2001, 2006) with Craie 701 borehole Cenomanian-Turonian δ13C curves. Abbreviations: Alb – Albian; Ce – Cenomanian; Sd – Stoliczkaia dispar; Mm – Mantelliceras mantelli; Md – Mantelliceras dixoni; Ci – Cunningtoniceras inerme; Ar – Acanthoceras rhotomagense; Aj – A. jukesbrownei; C – Calycoceras guerangeri; Mg – Metoicoceras geslinianum; N – Neocardioceras juddii; Wd – Watinoceras devonense; Fc – Fagesia catinus; Mn – Mammites nodosoides.

5.2 Carbon isotopic ratio evolution from epicontinental seas to pelagic realm during the Cenomanian-Turonian

After correlating the δ13C evolution curve for the epicontinental domain of the South-East Paris Basin with that of southern England, we can extend this approach to the Tethyan pelagic domain with the Vocontian Basin (Hyèges, Angles, Vergons and Lambruisse sections) using the data of Gyawali et al. (2017) and Takashima et al. (2009) and the Umbria-Marche Basin (Italy, Gubbio section) using the data of Jenkyns et al. (1994).

First of all, we must note the similarity of the general shape of the δ13C evolution curves at the four sites (Fig. 5). However, the equivalent thickness of the various isotopic events is variable from one basin to another due to very different sedimentation rates between sites (epicontinental and pelagic realm). So, the thickness of the studied interval is about 165 m in the south of England, 250 m in the Craie 701 borehole, about 700 m in the Vocontian Basin and 85 m in Umbria-Marche. In detail, if we examine the major positive excursion around the Cenomanian-Turonian boundary at each location, we observe that its thickness is 8 m in England, 20 m in the southeast Paris Basin, 40 m in the Vocontian Basin and 3 m in Umbria-Marche. Minor isotopic events in the condensed pelagic Italian section may be difficult to discern compared to the epicontinental sections or the hemipelagic basin.

Concerning the absolute values of δ13C, Figure 6 compares the isotopic ratios of the selected sections. We observe that during the Early Cenomanian, the δ13C values are relatively close at the four sites, around 1.5‰ in average in England and the Paris Basin. The lowest values with high variability correspond to the Vocontian Basin (from 0.8 to 1.7‰) and the strongest to Umbria-Marche (2.2‰). The carbon isotopic ratio then increases to reach the plateau preceding the major excursion of OAE 2. This plateau is present in the four curves and culminates around 2.6‰ but ranges from 2.1‰ (Vocontian) to 2.7‰ (the Craie 701 borehole).

The OAE 2 isotopic excursion is well marked at all sites however, its amplitude is variable. Two sites (south-east Paris Basin, Culver Cliff) have very strong amplitudes with a maximum of 4.8 and 5.2‰ respectively, while the excursion is less pronounced in Umbria Marche (maximum 3.2‰). So, we can see that the amplitude of the δ13C excursion is greater in the shelf seas than in the open ocean. The ratio is generally higher in the Vocontian Basin (4.7‰), than in the Umbria-Marche Basin (3.2‰), although the latter is considered to be more open.

Such an observation was already made by Kaiho et al. (2014) in a platform-basin transect in Spain and southeastern France (Grosheny et al., 2017). Deeper, more restricted basins are more favourable to anoxic conditions than shallower water realms. Ocean margin flooding could enhance the primary productivity which preferentially removed 12C, causing waters to become enriched in the heavier isotope and therefore higher in δ13C carbonates. The consequence in the deep part of the basin was the development of the oxygen minimum zone with more 12C-enriched buried organic carbon.

Generally, the δ13C carbonate depth transect from platform to basin shows lower values in the shallow-water sections and a higher ratio in the deep-water sections (Renard et al., 1982; Patterson and Walter, 1994; Jenkyns, 1995; Weissert et al., 2008; Schiffbauer et al., 2017). This is also observed for the Cenomanian and Upper Turonian in the present work (Fig. 6). This may reflect stronger, land-derived organic component inputs, remineralisation of marine and terrestrial organic carbon, different carbonate factory (aragonite or calcite and high-Mg-calcite), or fluctuating sea level leading to subaerial exposure of shallow water sediments and diagenetic overprinting.

Several studies have shown that greater amplitudes in δ13C are often recorded in shallow-water carbonate deposits than those seen in pelagic sections (Jenkyns, 1995; Vahrenkamp, 1996; Wissler et al., 2003; Millan et al., 2009; Weissert, 2018). In shallow-water settings, the isotopic composition of water, and, therefore, of carbonate sediment are less stable than in open seas. This could be due to numerous factors, such as marine productivity, nutrient input, temperature, atmospheric pCO2, etc.

The location of the δ13C excursion, relative to the sedimentological expression of OAE 2, also seems to be somewhat different. Its maximum occurs inside the Thomel level in the Vocontian Basin (Takashima et al., 2009), in the top part of the Plenus Marls in England (Jarvis et al., 2006) and immediately above this level in the south-east Paris Basin. For Umbria-Marche, isotope analyses of bulk carbonates (Fig. 5; Jenkyns et al., 1994) might suggest that the isotopic maximum occurs above the Bonarelli level. In fact, this pattern is an artefact due to the lithology of this level which is made of silica and organic matter and is completely free of carbonates (no isotopic analysis on such a mineralogy type is possible). Considering the isotopic curve for organic matter (Tsikos et al., 2004; Mort et al., 2007), the maximum δ13C lies within the Bonarelli level. In summary, relative to lithology, the OAE 2 isotopic excursion is expressed earlier (within the level rich in organic matter) in the oceanic facies than in the epicontinental facies (at the top of the organic level).

In the Early and Middle Turonian, the post-excursion plateau (around 3‰) is observed in England, in the Paris Basin and in the Umbria-Marche trends. This plateau is much less marked in the curve corresponding to the Vocontian Basin, where a large progressive decrease is observed (1.5‰). After this plateau, in the Late Turonian, the values drop in all sections to a minimum that is close to that of the Early Cenomanian. The curves then show a similar trend with a progressive excursion (around 2.5‰) leading to low values of δ13C (of around 2‰) at the Turonian-Coniacian boundary.

There is no relationship between the paleo-latitudinal position of the sites and the isotopic fluctuations and absolute isotopic values (Fig. 5), suggesting that the environmental parameters are the most significant control of carbonate δ13C, compared to the climate.

thumbnail Fig. 5

Correlation of Culver Cliff section (Paul et al., 1994; Jarvis et al., 2001, 2006), Craie 701 borehole, Vocontian Basin sections (Takashima et al., 2009; Gyawali et al., 2017) and Gubbio succession (Jenkyns et al., 1994) Cenomanian-Turonian δ13C curves.
thumbnail Fig. 6

Composite δ13C curves for epicontinental basin (Culver Cliff section and Craie 701 borehole) compared with pelagic basin (Vocontian Basin sections and Gubbio succession for the Cenomanian-Turonian interval. Comparison of the δ13C curves with the sea level changes (Haq, 2014). The red line represents a high sea level time and the blue arrow a huge regression.

5.3 Cenomanian-Turonian carbon isotope ratio: a proxy of the global carbon cycle modulated by the local environment

Globally, all isotopic fluctuations, from the Early Cenomanian to the Turonian-Coniacian boundary, are present in all marine settings. The main parameters controlling the fluctuations in the carbon isotopic ratio in the open ocean are marine productivity (organic and carbonate) and the preservation of organic carbon in sediments, in relation to the oxygen content of the seawater. On the marine platform, isotopically light, terrestrial organic matter may also influence the bulk signal of the sediments.

To distinguish the respective part of each of these parameters, it should be noted that the δ13C increase of the Cenomanian-Turonian transition develops over quite a long period of time (from Middle Cenomanian to Middle Turonian). The evolution curve of δ13C can be de-convolved into two events (Fig. 7). First, there is a large maximum corresponding to an increase of about 1 to 1.5‰ of the isotopic values, which covers the time interval between the Middle Cenomanian and Middle Turonian. This maximum is almost identical at all of the sites. Superimposed on this maximum is a sharp excursion just before the Cenomanian-Turonian boundary whose amplitude varies according to the sites (barely 1‰ in open pelagic areas, up to 2.5‰ in epicontinental areas).

The facies evolution in the Craie 701 borehole is consistent with this scheme. During the Early Cenomanian, low δ13C values correspond to chalk with phosphatic pebbles, glauconite and a high content of quartz and bioclasts which may reflect oligotrophic oceanic water and lower organic and carbonate productivity (weak CaCO3; Fig. 2). The increase in the isotopic carbon ratio at the beginning of the maximum (Middle Cenomanian) coincides with an increase in the calcium carbonate content in sediments and lower detrital inputs. Chalk is enriched in calcispheres, such as Pithonella, nannoconus (Amédro et al., 1978; Hart, 1991; Wendler et al., 2010). This association characterises a high level CaCO3 marine environment with oxygenated water and low terrestrial inputs. In numerous sections, the isotopic long-term maximum begins with a peak (called the mid-Cenomanian event 1) that is contemporary with an increase in the planktonic assemblages in sediments (Paul et al., 1994; Robaszynski et al., 1998; Jarvis et al., 2006).

In the Early Turonian, the last part of the isotopic plateau may also correspond to organic and carbonate productivities (Fig. 7). Abundant and diversified bioturbations have been recognised in the chalk just above the Plenus Marls, reflecting well-oxygenated water. The benthic diversity recovers rapidly in the higher beds of the Plenus Marls which coincide with lower siliciclastic content and the onset of the chalk deposits. High organic productivity in the beginning of the Turonian is confirmed by the abundance of pithonellids in the white chalk (Caus et al., 1997; Wendler and Willems, 2002; Wendler et al., 2002; Wilmsen, 2003; Wendler et al., 2010). High siliceous productivity indices were also recorded in the Umbria-Marche Basin, by the radiolarian-rich Bonarelli Level (Danelian et al., 2007; Musavu-Moussavou et al., 2007). Thus, for the various marine environments, the facies corresponding to the isotopic maximum, show indices of high productivity but no sign of oxygen deficiency in the water. During the maximum sea-level, increased mobilisation of nutrients through transgressive erosion have contributed to subsequently high marine productivity.

Therefore, we postulate (Fig. 7) that the Middle Cenomanian to Middle Turonian maximum should be part of a global marine productivity increase event of almost identical magnitude in the European and Tethyan domains. The positive excursion of the Cenomanian-Turonian boundary would correspond to an acme of productivity increase, reinforced by better preservation of organic carbon in relation to oxygen depletion of the seawater that varies according to the regions and the environmental context (epicontinental seas versus open ocean).

During OAE 2, most of the sites also show a low calcium carbonate content when the carbon isotopic ratio is high (Tsikos et al., 2004 and this study Figs. 3 and 7). An increase in volcanic emissions (from seafloor spreading and submarine intraplate volcanism) coincides with the OAE 2 event (Pitman, 1978; Schlanger et al., 1981; Larson, 1991). Reduced deep ocean ventilation has also been proposed by Monteiro et al. (2012). This sluggish oceanic circulation within the Atlantic should be responsible for a breakdown in the vertical structure of the oceanic water column and restricted deep-water exchange (Erbacher et al., 2001; Lüning et al., 2004; Voigt et al., 2004; Watkins et al., 2005; Forster et al., 2007; Tsandev and Slomp, 2009; Friedrich et al., 2012; Monteiro et al., 2012; Wagner et al., 2013; Goldberg et al., 2016). Enhanced CO2 emission in the atmosphere may also be responsible for an increase in the acidity of ocean surface waters and, thus, the partial dissolution of carbonate particles (Boulila et al., 2019).

The global evolution of the isotopic curve shows similarities with those of sea level fluctuations (Fig. 6) at the 3rd/2nd order scale (Haq et al., 1987; Haq, 2014). Notably, the first and the final parts of the isotopic maximum correspond to the two high sea level stands of Cenomanian and Turonian age. These transgressive phases have allowed the observation of a mixed Boreal and Tethyan fauna in the chalk series of the Aube, but also allowed North American fauna to be present as far as Tunisia (Amédro and Robaszynski, 2008).

This suggests that the productivities are induced by high sea levels. The important fast regression occurring at the Middle-Late Cenomanian boundary coincides with an interruption (the Jukes Browne event) in the gradual increase up to the maximum of high carbon isotopic values. A 40 cm thick conglomerate is observed in the studied core at this level. The sharp δ13C excursion of OAE 2 seems to be linked to a “minor” third order scale regression. This is also an argument for considering that another parameter added to productivity to drive this event. The low values observed in the Late Turonian would correspond to low sea levels (3rd scale) and the increase to the important transgressive phase at the Turonian-Coniacian boundary.

The sediments from the epicontinental sea (Anglo-Paris Basin) show higher carbonate and organic productivity from Late Cenomanian to Middle Turonian, which is recorded in the deposits through high δ13C values (Figs. 5 and 6). High detrital nutrient input from the emergent area, under a warm and humid climate (Arthur et al., 1988; Sarmiento et al., 1988; Jenkyns, 1999; Handoh and Lenton, 2003; Voigt et al., 2004; Forster et al., 2007; Jenkyns, 2010; Föllmi, 2012; Aguado et al., 2016; Nuñez-Useche et al., 2016), may induce high marine productivity. Considering the long-term sea level rise and the associated increase in the epicontinental sea surface, the carbonate deposition rate should be more elevated in this realm. A similar pattern has already been observed in Tunisia (Accarie et al., 1996; Amédro et al., 2005).

In Italy, the highest isotopic values in the Early Cenomanian and Late Turonian are explained by more stable trophic water conditions for the entire studied period. In the Anglo-Paris Basin at that time, low sea level and high terrigenous input from the continents disturbed the oceanic environment. More open conditions in the surface water and oligotrophic surface water may explain why, in the Early and Middle Turonian, Italian deposits show a similar but lower trend compared to the epicontinental sea.

The first important sedimentological events in the Turonian (breccia and marl deposits in the Middle Turonian) coincide with a huge decrease in sea level at the third order scale. This trend is subsequently confirmed by the numerous hardgrounds observed in the Upper Turonian (Amédro et al., 1997). Ammonite and echinoid migrations indicate a shift in the Boreal and Tethyan oceanic currents, such as the southward extension of cooler water masses (Voigt and Wiese, 2000). Lower organic productivity in surface water may be induced by these changes in ocean circulation and were recorded in sediments by an important decrease in δ13C.

The upper part of the Turonian and the Lower Coniacian show relatively high and stable δ13C. A return of high organic and calcium carbonate productivities may explain this trend. Maximum bioturbation, observed in the chalk interval (from 448 m to 508 m), confirms the existence of optimal living conditions throughout the water column.

thumbnail Fig. 7

Schematic trends and controls of the carbon isotopic ratio and calcium carbonate content of sediments across the Cenomanian-Turonian transition.

5.4 Oxygen isotopes and global climate changes

As previously mentioned, the oxygen isotope ratios (δ18O) in the Craie 701 borehole present a strong variability making their interpretation complicated. The two main factors affecting δ18O in marine carbonates are the temperature and salinity of the seawater. In the case of biogenic carbonates, the influence of the metabolic processes involved in the production of CaCO3 minerals (vital effect) is added. Due to the strong thermal dependency of δ18O, this parameter measured in bulk carbonate has often been considered to only represent the progression of recrystallisation during burial diagenesis. Over the past 30 years, numerous studies have shown that this is not the case and that burial diagenesis does not obliterate the original primary record of the physicochemical conditions prevailing at the time of deposition. For the chalk facies, the same result was demonstrated early on by Scholle and Arthur (1980) and, more recently, on the Craie 701 borehole by Chenot et al., 2016. Recent data obtained by separation of the various carbonate producers (Minoletti et al., 2005, 2007; Bojanowski et al., 2017; Tremblin and Minoletti, 2018) show that the evolution of the δ18O values of the bulk carbonate is similar to that of the different biogenic fractions (foraminifera, nannofossils) or the fraction of non-obvious origin (so-called micarb). Moreover, these results show that δ18O variability depends much more on the fluctuations of the relative percentages of the different producers (vital effect, different depths for the environments of organisms etc.) than on diagenetic effects.

In the Craie 701 borehole, no significant changes in fossil forms are observed in the chalk deposits on a large scale but, not having any data on the relative proportion of foraminifera, nannofossils and micarbs in the studied samples, we have chosen to take into account only the medium-term evolution of the δ18O trends. Relative to the diagenesis, in the chalk part (Turonian) of the core, Lasseur (2007) considered it to remain constant as was already demonstrated by Scholle and Arthur (1980) for the European outcrops. In contrast, for the grey limestone from the Cenomanian, the impact of the diagenetic processes of cementation and recrystallisation during burial cannot be excluded, especially since the absolute values of the δ18O in this level are low (close to −5‰, Fig. 2). Nevertheless, we can see on a δ13C-δ18O cross plot (Fig. 9), that there is no correlation (r2 = 0.16 for all data and r2 less than 0.3 for the Cenomanian sample) between the two signals. Chenot et al. (2016) and Le Callonnec et al. (2000) previously concluded that only the chalk at the top of the Poigny borehole had been altered by meteoric waters. This implies that diagenesis under the influence of groundwater is limited. In the case of strong diagenesis control, the δ13C of cement would become more depleted due to the oxidation of organic matter within sediments and, in the same way, the δ18O would be lower due to meteoric pore water cementation at an elevated temperature. So, even in Cenomanian facies, the δ18O variations can be explained, to a large degree, as primary environmental signal changes (i.e., seawater temperature and/or continental freshwater supplies).

The δ18O trends of the different sections show a consistent evolution (Fig. 9), from the Cenomanian to the Turonian-Coniacian boundary. Since no δ18O data is available from the Culver Cliff section, data from the East Kent section are used (Jenkyns et al., 1994), which is located a little further north on the east coast of England. During the Late Cenomanian, a decrease in the δ18O trend is observed, leading to the lowest isotopic values before the Cenomanian-Turonian boundary (Fig. 8). A global warming during the transition from a “warm greenhouse” (from Early Albian to Early Cenomanian) to a “hot greenhouse” (Middle Cenomanian) may explain this isotopic trend (Jenkyns et al., 1994). This warming could be due to enhanced CO2 emission in the atmosphere linked to several volcanic eruptions in the Caribbean, Ontong Java and Manihiki Plateaus (Ando et al., 2009; Selby et al., 2009; Ando et al., 2010; Jenkyns, 2010; Du Vivier et al., 2015 Nuñez-Useche et al., 2016). The increase in the carbonate 87Sr/86Sr ratio from the Early Albian (0.7072) to Late Cenomanian (0.7075; Bralower et al., 1997; Monnet, 2009) is consistent with an increase in ocean crust production during this period. The δ18O variations in some sections have shown a brief cooling during the global warming of OAE 2 (Morel, 1998; Jarvis et al., 2011; Grosheny et al., 2017; Jenkyns et al., 2017, Desmares et al., 2019) but this event is absent in the chalk of the Craie 701 borehole.

During the Turonian, the δ18O data show a progressive increase in two successive steps (Figs. 2 and 9) that seem to correspond to a long-term climatic cooling, as has been recorded in England (Jenkyns et al., 1994). Several macrofaunal migrations in the Turonian were related to a climatic cooling due to a change in ocean circulation and a major regression in the Late Turonian (de Graciansky et al., 1987; Clarke and Jenkyns, 1999; Stoll and Schrag, 2000; Voigt, 2000; Voigt and Wiese, 2000; Voigt et al., 2004; Jarvis et al., 2015).

Possible polar ice-sheet accumulation and glacio-eustatic effects have also been suggested (Stoll and Schrag, 2000; Voigt et al., 2004). The evidence comes from flora associations in the northern, high latitudes (Spicer and Parrish, 1990), δ18O values in marine macrofossils from Antarctica (Pirrie and Marshall, 1990) and the oceanic distribution of ice rafted debris (Frakes and Francis, 1988).

However, several remarks or reservations should be raised. First of all, the δ18O data during the Late Cenomanian have a much greater variability in the Paris Basin and the Vocontian Basin than in England and Italy. More importantly, there are large differences in the mean absolute values of the isotopic ratios recorded at different sites (Table 1).

It should be noted that the variations of the δ18O at the sites correlated do not reflect a paleo-latitudinal climate control. Diagenesis cannot be mentioned. Firstly, for the same facies (chalk) of environmental and burial conditions in England and the Paris Basin, the δ18O record is very different. Secondly, sediments from Umbria-Marche should be the most diagenetically altered (siliceous limestones with styloliths), but they show the highest δ18O, when it should be the opposite. In contrast, the lithological facies of the Paris Basin (soft chalk) have recorded lower δ18O. Therefore, we can consider that if the long-term evolution of oxygen isotopic ratios reflects warming during the Cenomanian and cooling during the Turonian, a regional paleo-environmental event must be added for the south-east part of the Paris Basin and the Vocontian Basin. That event is responsible for the variability and the lowest δ18O recorded in these sections. Since these regions represent a corridor of communication between the Boreal domain and the Tethyan Ocean, we can evoke the influence of continental freshwater and/or polar water with low δ18O and the competition between these waters and those of the Tethys to explain the high variability of δ18O.

England (close to the oceanic influence of the North Atlantic) and Umbria-Marche (in the open Tethys) would be excluded from the influence of these low δ18O waters. Such mass marine-water circulation (southward or northward during the Cenomanian and Turonian) has been proposed to explain the macrofauna migrations (Voigt and Wiese, 2000) and the changes in dinoflagellate cysts and the planktonic/benthic fauna ratio (Paul et al., 1994).

In detail, the Craie 701 isotopic curve (Fig. 3) shows that, during the deposition of the Plenus Marls, a clear downward δ18O trend can be interpreted as a warm climatic phase. As mentioned before, this warming could be related to an acme of volcanic and hydrothermal marine activity in major Large Igneous Provinces, releasing large amounts of greenhouse gases into the atmosphere during OAE 2 (Ando et al., 2009; Selby et al., 2009; Jenkyns, 2010; Erba et al., 2015; Nuñez-Useche et al., 2016). Such climatic warming with minimal vertical and latitudinal thermal gradients during OAE 2 has been suggested by numerous authors (Huber et al., 1999; Jenkyns, 1999; Voigt et al., 2004; Forster et al., 2007; Jenkyns, 2010; Aguado et al., 2016).

thumbnail Fig. 8

δ13C-δ18O cross-plot from bulk rock carbonate samples from the studied Craie 701 borehole. The correlation coefficients are reported for the different time interval. The global r2 is 0.1651.
thumbnail Fig. 9

Correlation of the East Kent section (Jenkyns et al., 1994), Craie 701 borehole, Vocontian Basin sections (Takashima et al., 2009; Gyawali et al., 2017) and the Gubbio succession (Jenkyns et al., 1994) Cenomanian-Turonian δ18O curves.
Table 1

Comparison of the mean oxygen stable isotope data (‰) from the studied core, East Kent (England), Vocontian Basin and Umbria-Marche sections for the key time intervals.

6 Conclusions

Stable carbon isotope chemostratigraphy carried out on the Craie 701 borehole allows for a high-resolution stratigraphic framework from the Cenomanian to the Turonian-Coniacian boundary. Thus, this borehole, located in the south-east Paris Basin, represents one of the best reference sections for correlating biostratigraphic and geochemical events between the Tethyan and Boreal realms. Comparison of the Craie 701 δ13C profile with those of various marine sites from northern European epicontinental seas and the Tethyan Ocean has shown similar and synchronous isotopic events which imply global processes. In particular, the middle-term increase in δ13C values (about 1.5‰ amplitude), which occurs from the Middle Cenomanian to Middle Turonian, corresponds to a global increase in organic and carbonate productivity. This event coincides with the mid-Cretaceous eustatic sea level rise.

OAE 2 is highlighted by a large, brief positive carbon isotope excursion (CIE), superimposed on the middle term trend that occurs during the Late Cenomanian. The CIE is present at all of the sites, whatever the sedimentary facies. However, its amplitude is variable from one site to another: it is higher (5.2‰) in the epicontinental sites, where anoxia is less pronounced, than in the open oceanic sites (3.2‰), where black shale intervals are recorded.

A regional and environmental modulation related to the preservation and burial of organic matter is thus superimposed on the global processes. The fluctuations of the calcium carbonate content in the latest Cenomanian sediments are likely to be linked to the massive volcanic CO2 degassing in the atmosphere and the acidification of seawater.

Variations in δ18O suggest that the Cenomanian-Turonian boundary represents a major turning point in the climatic history of the Earth, peaking at the OAE 2. In addition, δ18O data provide evidence of Turonian climate deterioration, characterised by synchronous stepped episodes of cooling throughout Europe, during the Early and Late Turonian in particular.

Acknowledgements

We would like to thank Damien Gendry (University of Rennes) for his assistance in sampling the Craie 701 borehole; and Nathalie Labourdette for the stable isotope measurements.

Special thanks to Maurice Renard and François Baudin for useful discussions that helped for the redaction of this manuscript; and to Jean-françois Deconinck and anonymous reviewer for their very helpful reviews.

This program research has been partially funded by the labex Matisse.

References

  • Accarie H, Emmanuel L, Robaszynski F, Baudin F, Amédro F, Caron M, et al. 1996. La géochimie isotopique du carbone (δ13C) comme outil stratigraphique. Application à la limite Cénomanien/Turonien en Tunisie Centrale. Comptes rendus des séances de l’Académie des sciences Paris 322(série II a): 579–586. [Google Scholar]
  • Aguado R, Reolid M, Molina E. 2016. Response of calcareous nannoplankton to the Late Cretaceous Oceanic Anoxic Event 2 at Oued Bahloul (central Tunisia). Palaeogeography Palaeoclimatology Palaeoecology 459: 289–305. [Google Scholar]
  • Ando A, Huber BT, MacLeod KG, Ohta T, Khim BK. 2009. Blake Nose stable isotopic evidence against the mid-Cenomanian glaciation hypothesis. Geology 37(5): 451–454. [Google Scholar]
  • Ando A, Huber BT, MacLeod KG. 2010. Depth-habitat reorganization of planktonic foraminifera across the Albian/Cenomanian boundary. Paleobiology 36: 357–373. [Google Scholar]
  • Amédro F, Robaszynski F. 2000. Les craies à silex du Turonien supérieur au Santonien du Boulonnais (France) au regard de la stratigraphie événementielle. Comparaison avec le Kent (U.K.). Géologie de la France 4: 39–56. [Google Scholar]
  • Amédro F, Robaszynski F. 2008. Zones d’ammonites et de foraminifères du Vraconien au Turonien : une comparaison entre les domaines boréal et téthysien (NW Europe/Tunisie centrale). Brest : Carnets de géologie/Notebooks on Geology, Note brève 2008/02-fr. [Google Scholar]
  • Amédro F, Robaszynski F. 2014. Le Crétacé du Bassin parisien. In Gely JP, Hanot F, eds.Le Bassin parisien, un nouveau regard sur la géologie. Bulletin information géologues bassin de Paris, Mémoire hors-série 9: 75–84. [Google Scholar]
  • Amédro F, Damotte R, Manivit H, Robaszynski F, Sornay J. 1978. Échelles biostratigraphiques dans le Cénomanien du Boulonnais (macro-micro-nanno fossiles). Géologie méditerranéenne 5(1): 5–18. [Google Scholar]
  • Amédro F, Manivit H, Robaszynski F. 1979. Echelles biostratigraphiques du Turonien au Santonien dans les Craies du Boulonnais (macro-micro-nanno fossiles). Annales de la Société géologique du Nord 98: 287–305. [Google Scholar]
  • Amédro F, Robaszynski F, Colleté C, Fricot C. 1997. Les craies du Cénomanien-Turonien de l’Aube et du Boulonnais : des événements litho- et biosédimentaires communs. Annales de la Société géologique du Nord 5: 189–197. [Google Scholar]
  • Amédro F, Accarie H, Robaszynski F. 2005. Position de la limite Cénomanien-Turonien dans la Formation Bahloul de Tunisie centrale : apports intégrés des ammonites et des isotopes du carbone (δ13C). Eclogae Geologicae Helvetiae 98: 151–167. [Google Scholar]
  • Arthur MA, Dean WE, Pratt LM. 1988. Geochemical and climatic effects of increased marine organic carbon burial at the Cenomanian/Turonian boundary. Nature 335: 714–717. [Google Scholar]
  • Barrier P. 2000. Etude microfaciologique de deux forages profonds dans la Craie de Provins (701 Poigny et 702 Sainte Colombe) : empilement des faciès, biodiversité et découpage séquentiel. Bulletin information géologues bassin de Paris 37(2): 33–43. [Google Scholar]
  • Bergerat F, Vandycke S. 1994. Palaeostress analysis and geodynamical implications of Cretaceous-Tertiary faulting in Kent and the Boulonnais. Journal of the Geological Society, London 151: 439–448. [Google Scholar]
  • Bojanowski M, Dubicka Z, Minoletti F, Olszewska-Nejbert D, Surowski M. 2017. Stable C and O isotopic study of the Campanian chalk from the Mielnik section (eastern Poland): Signals from bulk rock, belemnites, benthic foraminifera, nannofossils and microcrystalline cements. Palaeogeography Palaeoclimatology Palaeoecology 465A: 193–211. https://doi.org/10.1016/j.palaeo.2016.10.032. [Google Scholar]
  • Boulila S, Galbrun B, Driss S, Gardin S, Bartolini A. 2019. Constraints on the duration of the early Toarcian T-OAE and evidence for carbon-reservoir change from the High Atlas (Morocco). Global and Planetary Change 175. https://doi.org/10.1016/j.gloplacha.2019.02.005. [Google Scholar]
  • Boulila S, Charbonnier G, Spangenberg JE, Gardin S, Galbrun B, Briard J, et al. 2020. Unraveling short- and long-term carbon cycle variations during the Oceanic Anoxic Event 2 from the Paris Basin Chalk. Global and Planetary Change 186: 103126. [Google Scholar]
  • Bralower TJ, Fullagar PD, Paull CK, Dwyer GS, Leckie RM. 1997. Mid Cretaceous strontium-isotope stratigraphy of deep-sea sections. Geological Society of America Bulletin 109: 1421–1442. [Google Scholar]
  • Caus E, Teixell A, Bernaus JM. 1997. Depositional model of a Cenomanian-Turonian extensional basin (Sopeira basin, NE Spain): Interplay between tectonics, eustacy and biological productivity. Palaeogeography Palaeoclimatology Palaeoecology 129: 23–36. [Google Scholar]
  • Chenot E, Pellenard P, Martinez M, Deconinck JF, Amiotte-Suchet P, Thibault N, et al. 2016. Clay mineralogical and geochemical expressions of the “Late Campanian Event” in the Aquitaine and Paris basins (France): Palaeoenvironmental implications. Palaeogeography Palaeoclimatology Palaeoecology 447: 42–52. [Google Scholar]
  • Clarke LJ, Jenkyns HC. 1999. New oxygen isotope evidence for long-term Cretaceous climatic change in the Southern Hemisphere. Geology 27: 699–702. [Google Scholar]
  • Danelian T, Baudin F, Gardin S, Masure E, Ricordel C, Fili I, et al. 2007. The record of mid Cretaceous Oceanic Anoxic Events from the Ionian zone of southern Albania. Revue de micropaléontologie 50(3): 225–238. [Google Scholar]
  • Danzelle J, Riquier L, Baudin F, Thomazo C, Pucéat E. 2018. Oscillating redox conditions in the Vocontian Basin (SE France) during Oceanic Anoxic Event 2 (OAE 2). Chemical Geology. https://doi.org/10.1016/j.chemgeo.2018.05.039. [Google Scholar]
  • Deconinck JF, Amédro F, Fiolet-Piette A, Juignet P, Renard M, Robaszynski F. 1991a. Contrôle paléogéographique de la sédimentation argileuse dans le Cénomanien du Boulonnais et du Pays de Caux. Annales de la Société géologique du Nord 1(2): 57–66. [Google Scholar]
  • Deconinck JF, Amédro F, Desprairies A, Juignet P, Robaszynski F. 1991b. Niveaux repères de bentonites d’origine volcanique dans les craies du Turonien du Boulonnais et de haute-Normandie. Comptes rendus des séances de l’Académie des sciences Paris 312(Série II): 897–903. [Google Scholar]
  • Deconinck JF, Amédro F, Baudin F, Godet A, Pellenard P, Robaszynski F, et al. 2005. Late Cretaceous palaeoenvironments expressed by the clay mineralogy of Cenomanian–Campanian chalks from the east of the Paris Basin. Cretaceous Research 26(2): 171–179. [Google Scholar]
  • Desmares D, Testé M, Broche B, Tremblin M, Gardin S, Villier L, et al. 2019. High-resolution biostratigraphy and chemostratigraphy of the Cenomanian stratotype area (Le Mans, France). Cretaceous Research. https://doi.org/10.1016/j.cretres.2019.104198. [Google Scholar]
  • Du Vivier ADC, Selby D, Condon DJ, Takashima R, Nishi H. 2015. Pacific 187Os/188Os isotope chemistry and U–Pb geochronology: Synchroneity of global Os isotope change across OAE 2. Earth and Planetary Science Letters 428: 204–216. [Google Scholar]
  • Erba E, Bottini C, Faucher G. 2013. Cretaceous large igneous provinces: The effects of submarine volcanism on calcareous nannoplankton. Mineralogical Magazine 77: 1044. [Google Scholar]
  • Erba E, Duncan RA, Bottini C, Tiraboschi D, Weissert H, Jenkyns HC, et al. 2015. Environmental consequences of Ontong Java Plateau and Kerguelen Plateau Volcanism. Geological Society America Special Paper 511. https://doi.org/10.1130/2015.2511(15). [Google Scholar]
  • Erbacher J, Huber BT, Norris RD, Markey M. 2001. Increased thermo-haline stratification as a possible cause for an oceanic anoxic event in the Cretaceous period. Nature 409: 325–327. [Google Scholar]
  • Föllmi KB. 2012. Early Cretaceous life, climate and anoxia. Cretaceous Research 35: 230257. [Google Scholar]
  • Forster A, Schouten S, Moriya K, Wilson PA, Sinninghe Damsté JS. 2007. Tropical warming and intermittent cooling during the Cenomanian/Turonian oceanic anoxic event 2: Sea surface temperature records from the equatorial Atlantic. Paleoceanography 22: PA1219. https://doi.org/10.1029/2006PA001349. [Google Scholar]
  • Frakes LA, Francis JE. 1988. A guide to Phanerozoic cold polar climates from highlatitude ice-rafting in the Cretaceous. Nature 333: 547–549. [Google Scholar]
  • Friedrich O, Norris RD, Erbacher J. 2012. Evolution of middle to Late Cretaceous oceans – A 55 m.y. record of Earth’s temperature and carbon cycle. Geology 40: 107–110. [Google Scholar]
  • Gale AS, Jenkyns HC, Tsikos H, Van Breugel Y, Damsté JS, Bottini C, et al. 2018. High-resolution bio- and chemostratigraphy of an expanded record of Oceanic Anoxic Event 2 (Late Cenomanian-Early Turonian) at Clot Chevalier, near Barrême, SE France (Vocontian Basin, SE France). Newsletters on Stratigraphy, 29.05.2018. https://doi.org/10.1127/nos/2018/0445. [Google Scholar]
  • Goldberg T, Poulton SW, Wagner T, Kolonic SF, Rehkämper M. 2016. Molybdenum drawdown during Cretaceous Oceanic Anoxic Event 2. Earth and Planetary Science Letters 440: 81–91. [Google Scholar]
  • de Graciansky PC, Brosse E, Deroo G, Herbin JP, Muller C, Sigal J, et al. 1987. Organic-rich sediments and palaeoenvironmental reconstructions of the Cretaceous North America. In: Brooks J, Fleet AJ, eds. Marine petroleum source rocks. Geological Society London Special Publications 26: 317–344. [Google Scholar]
  • Grosheny D, Beaudoin B, Morel L, Desmares D. 2006. High-resolution biotratigraphy and chemostratigraphy of the Cenomanian/Turonian boundary event in the Vocontian basin, southeast France. Cretaceous Research 27(5): 629–640. [Google Scholar]
  • Grosheny D, Ferry S, Lécuyer C, Thomas A, Desmares D. 2017. The Cenomanian-Turonian boundary event (CTBE) on the southern slope of the Subalpine Basin (SE France) and its bearing on a probable tectonic pulse on a larger scale. Cretaceous Research 72: 39–65. [Google Scholar]
  • Guillocheau F, Robin C, Allemand P, Bourquin S, Brault N, Dromart G, et al. 2000. Mesocenozoic geodynamic evolution of the Paris Basin: 3D stratigraphic constraints. Geodinamica Acta 13: 189–246. [Google Scholar]
  • Godet A, Deconinck JF, Amédro F, Dron P, Pellenard P, Zimmerlin I. 2003. Enregistrement sédimentaire d’événements volcaniques dans le Turonien du Nord-Ouest du Bassin de Paris. Annales de la Société géologique du Nord 10(2): 147–162. [Google Scholar]
  • Gyawali BR, Nishi H, Takashima R, Herrle JO, Takayanagi H, Latil JL, et al. 2017. Upper Albian–upper Turonian calcareous nannofossil biostratigraphy and chemostratigraphy in the Vocontian Basin, southeastern France. Newsletters on Stratigraphy 50(2): 111–139. [Google Scholar]
  • Hallam A. 1984. Continental humid and arid zones during the Jurassic and Cretaceous. Palaeogeography, Palaeoclimatology, Palaeoecology 47: 195–223. [Google Scholar]
  • Hallam A. 1992. Phanerozoic sea-level changes. New-york: Columbia University Press. [Google Scholar]
  • Handoh IC, Lenton TM. 2003. Periodic mid-Cretaceous oceanic anoxic events linked by oscillations of the phosphorus and oxygen biogeochemical cycles. Global Biogeochemical Cycles 17(4): 1092. https://doi.org/10.1029/2003GB002039. [Google Scholar]
  • Hanot F. 2000. Apport industriel des forages du Programme CRAIE 700 pour la correction des variations latérales de vitesses dans la craie du Bassin de Paris. Bulletin information géologues bassin de Paris 37(2): 8–17. [Google Scholar]
  • Haq BU. 2014. Cretaceous eustasy revisited. Global and Planetary Change 113: 44–58. [Google Scholar]
  • Haq BU, Hardenbol J, Vail P.R. 1987. Chronology of fluctuating sea levels since the triassic. Science 235(4793): 1156–1167. [CrossRef] [PubMed] [Google Scholar]
  • Hart MB. 1991. The late Cenomanian calcisphere global bioevent. In: Grainger P, ed. Proceedings of the Annual Conference of the Ussher Society. Proceedings of the Ussher Society 7; 4. Ussher Society, Bristol, pp. 413–417. [Google Scholar]
  • Huber BT, Leckie RM, Norris RD, Bralower TJ, Cobabe E. 1999. Foraminiferal assemblage and stable isotopic change across the Cenomanian-Turonian boundary in the subtropical North Atlantic. Journal of Foraminiferal Research 29(4): 392–417. [Google Scholar]
  • Janin MC. 2000. Corrélations des forages Craie 700 d’après les nannofossiles calcaires. Bulletin information géologues bassin de Paris 37(2): 52–58. [Google Scholar]
  • Jarvis I, Murphy AM, Gale AS. 2001. Geochemistry of pelagic and hemipelagic carbonates: Criteria for identifying systems tracts and sea-level change. Journal of Geological Society of London 158: 685–696. [Google Scholar]
  • Jarvis I, Gale AS, Jenkyns HC, Pearce MA. 2006. Secular variation in Late Cretaceous carbon isotopes: A new δ13C carbonate reference curve for the Cenomanian–Campanian (99.6–70.6 Ma). Geological Magazine 143: 561–608. [Google Scholar]
  • Jarvis I, Lignum JS, Groecke DR, Jenkyns HC, Pearce MA. 2011. Black shale deposition, atmospheric CO2 drawdown, and cooling during the Cenomanian-Turonian Oceanic Anoxic Event. Paleoceanography 26: 1–17, Pa3201. https://doi.org/10.1029/2010pa002081. [Google Scholar]
  • Jarvis I, Trabucho JA, Gröcke DR, Ulicny D, Laurin J. 2015. Intercontinental correlation of organic carbon and carbonate stable isotope records: Evidence of climate and sea-level change during the Turonian (Cretaceous). The Journal of the International Association of Sedimentologists 1(2): 53–90. [Google Scholar]
  • Jenkyns HC. 1995. Carbon-isotope stratigraphy and paleoceanographic significance of the Lower Cretaceous shallow-water carbonates of resolution Guyot, Mid-Pacific Mountains. Proceedings of the Ocean Drilling Program Scientific Results 143: 99–104. [Google Scholar]
  • Jenkyns HC. 1999. Mesozoic anoxic events and palaeoclimate. Zentralblatt für Geologie Paläeontologie Teil I: 943949. [Google Scholar]
  • Jenkyns HC. 2010. Geochemistry of oceanic anoxic events. Geochemistry Geophysics Geosystems 11: Q03004. https://doi.org/10.1029/2009GC002788. [Google Scholar]
  • Jenkyns HC, Gale AS, Corfield RM. 1994. Carbon-isotope and oxygen-isotope stratigraphy of the English Chalk and Italian Scaglia and Its paleoclimatic significance. Geological Magazine 131(1): 1–34. [Google Scholar]
  • Jenkyns HC, Dickson AJ, Ruhl M, Van Den Boorn SH. 2017. Basalt-seawater interaction, the Plenus Cold Event, enhanced weathering and geochemical change: Deconstructing Oceanic Anoxic Event 2 (Cenomanian-Turonian, Late Cretaceous). Sedimentology 64: 16–43. [Google Scholar]
  • Joo YJ, Sageman BB. 2014. Cenomanian to Campanian carbon isotope chemostratigraphy from the Western Interior Basin, U.S.A. Journal of Sedimentary Research 54: 529–542. [Google Scholar]
  • Kaiho K, Katabuchi M, Oba M, Lamolda M. 2014. Repeated anoxia-extinction episodes progressing from slope to shelf during the latest Cenomanian. Gondwana Research 25: 1357–1368. [Google Scholar]
  • Larson RL. 1991. Latest pulse of Earth; evidence for a mid-Cretaceous super-plume. Geology 19: 547–550. [Google Scholar]
  • Lasseur E. 2007. La Craie du Bassin de paris (Cénomanien-Campanien, Crétacé supérieur). Sédimentologie de faciès, stratigraphie séquentielle et géométrie 3D (unpubl. PhD thesis). University of Rennes 1, 435 p. [Google Scholar]
  • Le Callonnec L, Renard M, Pomerol B, Janodet C, Caspard E. 2000. Données géochimiques préliminaires sur la série Cénomano-campanienne des forages 701 et 702 du programme Craie 700. Bulletin information géologues bassin de Paris 37(2): 112–119. [Google Scholar]
  • Lüning S, Kolonic S, Belhadj EM, Belhadj Z, Cota L, Baric G, et al. 2004. Integrated depositional model for the Cenomaniane-Turonian organic-rich strata in North Africa. Earth Science Reviews 64: 51–117. [Google Scholar]
  • Masure E. 2000. Les kystes de dinoflagellés en matière organique des forages du Programme Craie 700. Bulletin information géologues bassin de Paris 37(2): 44–51. [Google Scholar]
  • Mégnien C, Hanot F. 2000. Programme Craie 700. Deux forages scientifiques profonds pour étudier les phénomènes diagénétiques de grande ampleur dans la craie du Bassin de Paris. Bulletin information géologues bassin de Paris 37(2): 3–7. [Google Scholar]
  • Millan MI, Weissert HJ, Fernandez-Mendiola PA, Garcia-Mondéjar J. 2009. Impact of Early Aptian carbon cycle perturbations on evolution of a marine shelf system in the Basque-Cantabrian Basin (Aralar, N Spain). Earth and Planetary Science Letters 287: 392–401. [CrossRef] [Google Scholar]
  • Minoletti F, de Rafélis M, Renard M, Gardin S, Young J. 2005. Changes in the pelagic fine fraction carbonate sedimentation during the Cretaceous-Paleocene transition: Contribution of the separation technique to the study of Bidart section. Palaeogeography Palaeoclimatology Palaeoecology 216: 119–137. [Google Scholar]
  • Minoletti F, Hermoso M, Gressier V. 2007. Deciphering the geochemistry of calcareous pelagic producers: Beyond bulk carbonate analyses. Eos Trans AGU 88 Fall Meeting Supplementary, Abstract id. PP31C–0531. [Google Scholar]
  • Monnet C. 2009. The Cenomanian-Turonian boundary mass extinction (Late Cretaceous): New insights from ammonoid biodiversity patterns of Europe, Tunisia and the Western Interior (North America). Palaeogeography Palaeoclimatology Palaeoecology 282: 88104. [Google Scholar]
  • Monteiro FM, Pancost RD, Ridgwell A, Donnadieu Y. 2012. Nutrients as the dominant control on the spread of anoxia and euxinia across the Cenomanian‐Turonian oceanic anoxic event (OAE 2): Model‐data comparison. Paleoceanography 27(4): PA4209. https://doi.org/10.1029/2012PA002351. [Google Scholar]
  • Morel L. 1998. Stratigraphie à haute résolution du passage Cénomanien-Turonien. Thèse de l’Université Pierre et Marie Curie, Paris VI, 224. [Google Scholar]
  • Mort HP, Adatte T, Föllmi KB, Keller G, Steinmann P, Matera V, et al. 2007. Phosphorus and the roles of productivity and nutrient recycling during oceanic anoxic event 2. Geology 35(6): 483–486. [Google Scholar]
  • Mortimore R.N. 1983. The stratigraphy and sedimentation of the Turonian-Campanian in the Southern Province of England. Zitteliana 10: 27–41. [Google Scholar]
  • Mortimore R.N. 2011. A chalk revolution: What have we done to the chalk of England? Proceedings of the Geologists’ Association 122: 232–297. [Google Scholar]
  • Mortimore RN. 2014. Logging the Chalk. Scotland (UK): Whitles Publishing, 357 p. [Google Scholar]
  • Mortimore R, Pomerol B. 1997. Upper Cretaceous tectonic phases and end Cretaceous inversion in the Chalk of the Anglo-Paris Basin. Proceedings of the Geologists Association 108: 231–255. [Google Scholar]
  • Mortimore RN, Wood CJ, Pomerol B, Ernst G. 1996. Dating the phases of the Subhercynian tectonic epoch: Late Cretaceous tectonics and eustatics in the Cretaceous basins of northern Germany compared with the Anglo-Paris Basin. Zentralblatt für Geologie und Paläontologie Teil I(11/12): 1349–1401. [Google Scholar]
  • Mortimore RN, Wood CJ, Gallois RW. 2001. British Upper Cretaceous Stratigraphy. Geological Conservation Review Series 23. Peterborough: Joint Nature Conservation Committee, 558 p. [Google Scholar]
  • Musavu-Moussavou B, Danelian T, Baudin F, Coccioni R, Frölich F. 2007. The radiolarian biotic response during OAE 2. A high-resolution study across the Bonarelli level at Bottaccione (Gubbio, Italy). Revue de micropaléontologie 50: 253–287. [Google Scholar]
  • Nuñez-Useche F, Canet C, Barragan R, Alfonso P. 2016. Bioevents and redox conditions around the Cenomanian–Turonian anoxic event in Central Mexico. Palaeogeography Palaeoclimatology Palaeoecology 449: 205–226. [Google Scholar]
  • Patterson W, Walter L. 1994. Depletion of 13C in seawater ƩCO2 on modern carbonate platform: Significance for the carbon isotopic record of carbonates. Geology 22: 885–888. [Google Scholar]
  • Paul CRC, Mitchell SF, Marshall JD, Leafy PN, Gale AS, Duane AM, et al. 1994. Palaeoceanographic events in the Middle Cenomanian of Northwest Europe. Cretaceous Research 15(6): 707–738. [Google Scholar]
  • Philip J, Floquet M. 2000. Late Cenomanian. In: Dercourt J, Gaetani B, Vrielynck E, Barrier B, Biju Duval B, Brunet MF, et al., eds. Atlas Peri-Tethys, Palaeogeographical maps. Paris: CCGM/CGMW, map 14. [Google Scholar]
  • Pirrie D, Marshall JD. 1990. Diagenesis of Inoceramus and Late Cretaceous paleoenvironmental geochemistry: A case study from James Ross Island, Antarctica. Palaios 5: 336–345. [Google Scholar]
  • Pitman WC. 1978. Relation between eustasy and stratigraphic sequences of passive margins. Geological Society of America Bulletin 89: 1389–1403. [Google Scholar]
  • Pomerol B. 1983. Geochemistry of the Late Cenomanian-Early Turonian chalks of the Paris Basin: Manganese and carbon isotopes in carbonates as paleooceanographic indicators. Cretaceous Research 4: 85–93. [Google Scholar]
  • Pomerol B. 2000. Le forage de Sainte-Colombe (702) : description lithologique. Bulletin d’information des géologues du bassin de Paris 37: 27e32. [Google Scholar]
  • Renard M, Delacotte O, Létolle R. 1982. Le strontium et les isotopes stables dans les carbonates totaux de quelques sites de l’Atlantique et de la Téthys. Bulletin de la Société géologique de France 7, 24(3): 519–534. [Google Scholar]
  • Robaszynski F. 2000. Le forage de Poigny (701) : description lithologique. Bulletin d’information des géologues du bassin de Paris 37(2): 18–26. [Google Scholar]
  • Robaszynski F, Bellier J.P. 2000. Biostratigraphie du Crétacé avec les foraminifères dans les forages de Poigny et de Sainte-Colombe. Bulletin d’information des géologues du bassin de Paris 37(2): 59–65. [Google Scholar]
  • Robaszynski F, Amédro F, Colleté C, Fricot C. 1987. La limite Cénomanien-Turonien dans la région de Troyes (Aube, France). Bulletin d’information des géologues du bassin de Paris 24(4): 7–24. [Google Scholar]
  • Robaszynski F, Gale AS, Juignet P, Amédro F, Hardenbol J. 1998. Sequence stratigraphy in the upper Cretaceous of the Anglo-Paris basin exemplified by the Cenomanian Stage. In: Hardenbol J, Thierry J, Farley MB, Jaquin T, de Graciansky PC, Vail PR, eds. Mesozoic and Cenozoic sequence stratigraphy of European Mesozoic basins. SEPM special Publications 60: 363–386. [Google Scholar]
  • Robaszynski F, Pomerol B, Masure E, Janin MC, Bellier JP, Damotte R. 2000. Corrélations litho-biostratigraphiques et position des limites d’étages dans le Crétacé des sondages de Poigny et de Sainte-Colombe : une synthèse des premiers résultats. Bulletin d’information des géologues du bassin de Paris 37(2): 74–85. [Google Scholar]
  • Robaszynski F, Pomerol B, Masure E, Bellier JP, Deconinck JF. 2005. Stratigraphy and stage boundaries in reference sections of the Upper Cretaceous Chalk in the east of the Paris Basin: The “Craie 700” Provins boreholes. Cretaceous Research 26(2): 157–169. [Google Scholar]
  • Sarmiento JL, Herbert TD, Toggweiler JR. 1988. Causes of anoxia in the world ocean. Global Biogeochemical Cycles 2: 115–128. [Google Scholar]
  • Schiffbauer J, Huntley J, Fike D, Jeffrey M, Gregg J, Shelton K. 2017. Decoupling biogeochemical records, extinction, and environmental change during the Cambrian SPICE event. Science Advances 3: 1–7. [Google Scholar]
  • Schlanger SO, Jenkyns HC. 1976. Cretaceous oceanic anoxic events: Causes and consequences. Geologie Mijnbouw 55: 179–184. [Google Scholar]
  • Schlanger SO, Jenkyns HC, Premoli-Silva I. 1981. Volcanism and vertical tectonics in the Pacific Basin related to global Cretaceous transgression. Earth and Planetary Science Letters 52: 435–449. [Google Scholar]
  • Scholle PA, Arthur MA. 1980. Carbon isotope fluctuations in Cretaceous pelagic limestones; potential stratigraphic and petroleum exploration tool. American Association of Petroleum Geologists Bulletin 64: 67–87. [Google Scholar]
  • Selby D, Mutterlose J, Condon DJ. 2009. U–Pb and Re–Os geochronology of the Aptian/Albian and Cenomanian/Turonian stage boundaries: Implications for timescale calibration, osmium isotope seawater composition and Re–Os systematics in organicrich sediments. Chemical Geology 265: 394–409. https://doi.org/10.1016/j.chemgeo.2009.05.005. [Google Scholar]
  • Spicer RA, Parrish JT. 1990. Late Cretaceous-early Tertiary palaeoclimates of northern high latitudes: A quantitative view. Journal of the Geological Society London 147: 329341. [Google Scholar]
  • Stoll HM, Schrag DP. 2000. High-resolution stable isotope records from the Upper Cretaceous rocks of Italy and Spain: Glacial episodes in a greenhouse planet? Geological Society America Bulletin 112: 308–319. [Google Scholar]
  • Takashima R, Nishi H, Hayashi K, Okada H, Kawahata H, Yamanaka T, et al. 2009. Litho-, bio- and chemostratigraphy across the Cenomanian/Turonian boundary (OAE 2) in the Vocontian Basin of southeastern France. Palaeogeography Palaeoclimatology Palaeoecology 273(1): 61–74. [Google Scholar]
  • Tremblin M, Minoletti F. 2018. The meridional temperature gradient under greenhouse climatic state: A data-models discrepancy? Geophysical Research Abstracts 20: 9559. [Google Scholar]
  • Tsandev I, Slomp CP. 2009. Modeling phosphorus cycling and carbon burial during Cretaceous Oceanic Anoxic Events. Earth and Planetary Science Letters 286: 71–79. [Google Scholar]
  • Tsikos H, Jenkyns HC, Walsworth-Bell B, Petrizzo MR, Forster A, Kolonic S, et al. 2004. Carbonisotope stratigraphy recorded by the Cenomanian–Turonian Oceanic Anoxic Event: Correlation and implications based on three key localities. Journal of the Geological Society 161(4): 711–719. [Google Scholar]
  • Vahrenkamp VC. 1996. Carbon isotope stratigraphy of the Upper Kharaib and Shuaiba Formations: Implications for the Early Cretaceous Evolution of the Arabian Gulf Region. American Association of Petroleum Geologists 80(5): 647–662. [Google Scholar]
  • Vanderaveroet P, Amédro F, Colleté C, Deconinck JF, Récourt P, Robaszynski F. 2000. Caractérisation et extension de niveaux repères de bentonites dans le Turonien supérieur du Bassin de Paris (Boulonnais, Aube). Geodiversitas 22(3): 457–469. [Google Scholar]
  • Vandyckes S, Bergerat F. 1992. Tectonique de failles et paléo-contraintes dans les formations crétacées du Boulonnais (France). Implications géodynamiques. Bulletin de la Société géologique de France 163(5): 553–560. [Google Scholar]
  • Voigt S. 2000. Cenomanian–Turonian composite δ13C curve for Western and Central Europe: The role of organic and inorganic carbon fluxes. Palaeogeography, Palaeoclimatology, Palaeoecology 160: 91–104. [Google Scholar]
  • Voigt S, Hilbrecht H. 1997. Late Cretaceous carbon isotope stratigraphy in Europe: Correlation and relations with sea level and sediment stability. Palaeogeography, Palaeoclimatology, Palaeoecology 134: 39–59. [Google Scholar]
  • Voigt S, Wiese F. 2000. Evidence for Late Cretaceous (Late Turonian) climate cooling from oxygen-isotope variations and palaeobiogeographic changes in Western and Central Europe. Journal of the Geological Society 57(4): 737–743. [Google Scholar]
  • Voigt S, Gale AS, Flogel S. 2004. Midlatitude shelf seas in the Cenomanian-Turonian greenhouse world: Temperature evolution and North Atlantic circulation. Paleoceanography 19: PA4020. https://doi.org/10.1029/2004PA001015. [Google Scholar]
  • Voigt S, Aurag A, Leis F, Kaplan U. 2007. Late Cenomanian to Middle Turonian high-resolution carbon isotope stratigraphy: New data from the Münsterland Cretaceous Basin, Germany. Earth and Planetary Science Letters 253: 196–210. [Google Scholar]
  • Wagner T, Hofmann P, Flögel S. 2013. Marine black shale deposition and Hadley Cell dynamics: A conceptual framework for the Cretaceous Atlantic Ocean. Marine and Petroleum Geology 43: 222–238. [Google Scholar]
  • Watkins DK, Cooper MJ, Wilson PA. 2005. Calcareous nannoplankton response to late Albian oceanic anoxic event 1d in the western North Atlantic. Paleoceanography 20(2): PA2010. https://doi.org/10.1029/2004PA001097. [Google Scholar]
  • Weissert HJ. 2018. Jurassic‐Cretaceous Carbon Isotope Geochemistry–Proxy for Paleoceanography and Tool for Stratigraphy. In: Sial AN, Gaucher C, Ramkumar M, Ferreira VP, eds. Chemostratigraphy Across Major Chronological Boundaries. Geophysical Monograph Series 240: 211–221. [Google Scholar]
  • Weissert H, Joachimski M, Sarnthein M. 2008. Chemostratigraphy. Newsletters on Stratigraphy 42(3): 145–179. [Google Scholar]
  • Wendler J, Willems H. 2002. Distribution pattern of calcareous dinoflagellate cysts across the Cretaceous–Tertiary boundary (Fish Clay, Stevns Klint, Denmark); implications for our understanding of species-selective extinction. In: Koeberl C, MacLeod KG, eds. Catastrophic events and mass extinctions; impacts and beyond. Boulder(Colorado): Geological Society of America (GSA), 356, pp. 265–275. [Google Scholar]
  • Wendler J, Graefe KU, Willems H. 2002. Palaeoecology of calcareous dinoflagellate cysts in the mid-Cenomanian Boreal Realm; implications for the reconstruction of palaeoceanography of the NW European shelf sea. Cretaceous Research 23: 213–229. [Google Scholar]
  • Wendler JE, Lehmann J, Kuss J. 2010. Orbital time scale, intra-platform basin correlation, carbon isotope stratigraphy and sea-level history of the Cenomanian-Turonian Eastern Levant platform, Jordan. In: Homberg C, Bachmann M, eds. Evolution of the Levant Margin and Western Arabia platform since the Mesozoic. Geological Society, London, Special Publication 341: 171–186. [Google Scholar]
  • Wilmsen M. 2003. Sequence stratigraphy and palaeoceanography of the Cenomanian Stage in northern Germany. Cretaceous Research 24: 525–568. [Google Scholar]
  • Wissler L, Funk H, Weissert HJ. 2003. Response of Early Cretaceous carbonate platforms to changes in atmospheric carbon dioxide levels. Palaeogeography Palaeoclimatology Palaeoecology 200: 187–205. [Google Scholar]
  • Wray D. 1999. Identification and long-range correlation of bentonites in Turonian-Coniacian (Upper Cretaceous) chalks of northwest Europe. Geological Magazine 136(4): 361–371.and Floquet [Google Scholar]

Cite this article as: Le Callonnec L, Briard J, Boulila S, Galbrun B. 2021. Late Cenomanian-Turonian isotopic stratigraphy in the chalk of the Paris Basin (France): a reference section between the Tethyan and Boreal realms, BSGF - Earth Sciences Bulletin 192: 14.

All Tables

Table 1

Comparison of the mean oxygen stable isotope data (‰) from the studied core, East Kent (England), Vocontian Basin and Umbria-Marche sections for the key time intervals.

In the text

All Figures

thumbnail Fig. 1

Late Cenomanian paleogeography of Western Europe and a part of the Tethys (modified from Philip and Floquet, 2000). A rough location of some key sections: 1. Craie 701 borehole (this study); 2. Eastbourne, Sussex, UK (Paul et al., 1994); 3. Lambruisse, Vocontian Basin, France (Takashima et al., 2009); 4. Gubbio, Umbria-Marche Basin, Italy (Tsikos et al., 2004).
In the text
thumbnail Fig. 2

Carbon and oxygen isotopic profiles of the Craie 701 borehole. Lithology (bentonite levels are identified by a green dashed line) and biostratigraphy are from Robaszynski et al. (2005). Subdivisions of the δ13C profile into eight carbon isotope sequences (CIS) labelled I though VIb are also mentioned.
In the text
thumbnail Fig. 3

Carbon and oxygen isotopes and carbonate content data of the Craie 701 borehole across the Cenomanian-Turonian boundary.
In the text
thumbnail Fig. 4

Correlation of the Culver Cliff section (Paul et al., 1994; Jarvis et al., 2001, 2006) with Craie 701 borehole Cenomanian-Turonian δ13C curves. Abbreviations: Alb – Albian; Ce – Cenomanian; Sd – Stoliczkaia dispar; Mm – Mantelliceras mantelli; Md – Mantelliceras dixoni; Ci – Cunningtoniceras inerme; Ar – Acanthoceras rhotomagense; Aj – A. jukesbrownei; C – Calycoceras guerangeri; Mg – Metoicoceras geslinianum; N – Neocardioceras juddii; Wd – Watinoceras devonense; Fc – Fagesia catinus; Mn – Mammites nodosoides.
In the text
thumbnail Fig. 5

Correlation of Culver Cliff section (Paul et al., 1994; Jarvis et al., 2001, 2006), Craie 701 borehole, Vocontian Basin sections (Takashima et al., 2009; Gyawali et al., 2017) and Gubbio succession (Jenkyns et al., 1994) Cenomanian-Turonian δ13C curves.
In the text
thumbnail Fig. 6

Composite δ13C curves for epicontinental basin (Culver Cliff section and Craie 701 borehole) compared with pelagic basin (Vocontian Basin sections and Gubbio succession for the Cenomanian-Turonian interval. Comparison of the δ13C curves with the sea level changes (Haq, 2014). The red line represents a high sea level time and the blue arrow a huge regression.
In the text
thumbnail Fig. 7

Schematic trends and controls of the carbon isotopic ratio and calcium carbonate content of sediments across the Cenomanian-Turonian transition.
In the text
thumbnail Fig. 8

δ13C-δ18O cross-plot from bulk rock carbonate samples from the studied Craie 701 borehole. The correlation coefficients are reported for the different time interval. The global r2 is 0.1651.
In the text
thumbnail Fig. 9

Correlation of the East Kent section (Jenkyns et al., 1994), Craie 701 borehole, Vocontian Basin sections (Takashima et al., 2009; Gyawali et al., 2017) and the Gubbio succession (Jenkyns et al., 1994) Cenomanian-Turonian δ18O curves.
In the text

Current usage metrics show cumulative count of Article Views (full-text article views including HTML views, PDF and ePub downloads, according to the available data) and Abstracts Views on Vision4Press platform.

Data correspond to usage on the plateform after 2015. The current usage metrics is available 48-96 hours after online publication and is updated daily on week days.

Initial download of the metrics may take a while.