Contemporaneous massive subaerial volcanism and late cretaceous Oceanic Anoxic Event 2

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Earth and Planetary Science Letters 256 (2007) 211 – 223 www.elsevier.com/locate/epsl

Contemporaneous massive subaerial volcanism and late cretaceous Oceanic Anoxic Event 2 Junichiro Kuroda a,⁎, Nanako O. Ogawa a , Masaharu Tanimizu b , Millard F. Coffin c , Hidekazu Tokuyama c , Hiroshi Kitazato a , Naohiko Ohkouchi a a

b

Institute for Research on Earth Evolution, Japan Agency for Marine-Earth Science and Technology, 2-15 Natsushima-cho, Yokosuka, 237-0061, Japan Kochi Institute for Core Sample Research, Japan Agency for Marine-Earth Science and Technology, B200 Monobe, Nankoku, 783-8502, Japan c Ocean Research Institute, University of Tokyo, 1-15-1 Minamidai, Nakano, Tokyo, 164-8639, Japan Received 6 June 2006; received in revised form 22 January 2007; accepted 22 January 2007 Available online 30 January 2007 Editor: M.L. Delaney

Abstract Oceanic Anoxic Events (OAEs) are geological time intervals characterized by extremely high burial rates of organic carbon that led to deposition of organic-rich “black shales” in the global ocean. It has been proposed that oceanic anoxic events are ultimately caused by massive volcanism associated with formation of large igneous provinces (LIPs) because of chronological similarities, but no general consensus has developed yet. To investigate the possibility of LIP volcanism instigating OAEs, we measured stable isotopic compositions of bulk organic carbon (δ13Corg) and lead (Pb) isotopic compositions in the silicate sediment fraction across the Bonarelli black shale (Italy), a type stratigraphic section for the end-Cenomanian OAE (OAE-2; 94 Ma). Ultra-high-resolution δ13Corg records determined every 1.5 mm capture a 3‰ sharp negative shift at the base of the Bonarelli. At the same stratigraphic level, Pb isotopic compositions in the silicate sediment fraction exhibit significant shifts toward characteristic values of volcanic rocks from contemporaneous LIPs (Caribbean and Madagascar flood basalts). These data suggest a rapid, substantial increase in the relative supply of silicate minerals from the two LIPs. Massive subaerial volcanism associated with LIP formation provides a simple explanation for these two isotopic geochemical signals via release of a huge amount of carbon dioxide (∼ 105 Gt CO2) and particulate materials into the atmosphere, which resulted in a rapid negative shift of δ13C in sea water and changes in Pb isotopic compositions in the silicate sediment fraction, respectively. We interpret that massive volcanism triggered significant climatic changes, inducing biotic crises and oceanic anoxia. © 2007 Elsevier B.V. All rights reserved. Keywords: Oceanic anoxic event; Black shale; Carbon isotopes; Lead isotopes; Large igneous provinces

1. Introduction Cretaceous time is punctuated by multiple depositional episodes of organic-rich “black shale” throughout ⁎ Corresponding author. Tel.: +81 46 867 9620; fax: +81 46 867 9315. E-mail address: [email protected] (J. Kuroda). 0012-821X/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2007.01.027

the global ocean (i.e., Tethys, Atlantic, and equatorial Pacific) that have become known as Oceanic Anoxic Events (OAEs) [1,2]. An anoxic event that occurred immediately before the Cenomanian–Turonian (C–T) boundary (93.6 Ma) [3], termed “OAE-2”, has been considered to be typical of OAEs because of its worldwide occurrence [1,2,4] and highly elevated organic carbon

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content [5]. It has been proposed that OAEs are ultimately caused by rapid, voluminous release of methane from gas hydrates in marine sediment [6,7], or by massive volcanic events associated with the emplacement of large igneous provinces (LIPs) and sea-floor spreading at mid-ocean ridges [8–13]. The latter has been particularly well investigated because of chronological similarities between LIPs and OAEs [8–13], e.g., the timings of OAE-1a (Early Aptian) and OAE-2 are close to those of emplacements of the Ontong Java Plateau [14–16] and Madagascar/Caribbean flood basalts [13,17,18], respectively. Hydrothermal activity associated with emplacement of the Caribbean LIP has been proposed as a trigger for OAE-2 because of rapid enhancement of hydrothermally derived trace elements in black shales [13]. However, general consensus on a causal mechanism has not yet been achieved. Stable isotopic compositions of organic and carbonate carbon in marine sediments are useful for tracing the carbon cycle in the ocean-atmosphere system in the geological past. LIP volcanism likely perturbed the global carbon cycle by releasing huge amounts of CO2 into the system. For instance, a short-term negative shift at the onset of OAE-1a [19,20] suggests rapid release of isotopically light carbon into the ocean–atmosphere system. Although the cause of the negative spike is not yet fully understood, it has been attributed either to methane release or volcanic CO2 degassing. In contrast, such an isotopic event has not been considered for OAE-2. However, a small but sharp negative shift in δ13C has been reported from some sites, including southern England [21], the southeastern North Atlantic [22], the western Pacific (Japan) [23,24] and the Western Interior Seaway (WIS) in North America [25]. More detailed and highresolution studies are essential to assess carbon isotopic variations during OAE-2. In addition to stable isotopic ratios of sedimentary carbon, some radiogenic isotopes in sediment may provide information on temporal relationships between volcanic and paleoenvironmental events. In particular, Pb isotopes in the silicate fraction of sediment appear suitable for tracing changes in relative contributions of volcano-derived silicates in marine sediments, as Pb isotopic compositions in volcanic rock have been relatively well investigated [14–16,26–31]. In this study we present ultra-high-resolution isotopic records of total organic carbon (δ13 Corg) measured every 1.5 mm as well as Pb isotopic compositions of silicates in the stratigraphic section containing the “Bonarelli” black shale that was deposited on a pelagic marginal shelf in western Tethys during OAE-2 [32,33] (Fig. 1). We use these geochemical data

to investigate temporal relationships between LIPs and OAE-2. 2. Samples and methods 2.1. Bonarelli black shale An almost continuous sequence of Triassic to Miocene strata is found in the Umbrian Apennines, central Italy [34]. The Upper Cretaceous sedimentary succession is mainly composed of fine-grained chalk with bedded and/ or nodular chert. The distinctive Bonarelli black shale, averaging 1 m thick, is intercalated in the upper part of Scaglia Bianca Formation (Albian to basal Turonian). The top of the planktonic foraminiferal Rotalipola cushmani Zone lies approximately 15 cm below the base of the Bonarelli in the Botaccione section [35], and the Bonarelli itself is enclosed within Whiteinella archaeocretacea Zone [35,36]. Tsikos et al. [36] reported that the first appearance of a marker nannofossil Quadrum gartneri, which essentially corresponds to the C–T boundary, was observed ca. 1 m above the Bonarelli. The Bonarelli is characterized by millimeter-scale alternations of dark and light layers (Fig. 1). Light layers mainly contain abundant radiolaria, whereas dark layers are enriched in organic carbon [5]. The scale of the alternating light and dark layers generally ranges from 1.5 to 50 mm. On the basis of microscopic observations, there is no clear evidence for bioturbation or trace fossils within each layer [5]. The sedimentary succession in the Umbrian Apennines was deposited on the pelagic marginal shelf of the Apulian Platform (western Tethys), and isolated from both significant terrigenous input and redeposited shallow-water sediment [33]. Redeposited sediment is rarely observed in the Bonarelli [5]. Intervals below and above the Bonarelli are mainly composed of fine-grained chalk with discontinuous chert layers (Fig. 1). Near-continuous samples of the Bonarelli and adjacent sedimentary rocks were taken at the Gorgo Cerbara outcrop (Fig. 1). The thickness of the Bonarelli at the sampling site was approximately 104 cm. Detailed sampling protocols are given in Kuroda et al. [5]. Samples for stable isotopic analysis of bulk organic carbon were cut into slices every 1.5 mm and then pulverized in an agate mortar after repeated cleaning with deionized water, ethanol and acetone. Rock samples for biomarker analyses and Pb isotopic measurements were obtained separately from the same outcrop. After removal of weathered surfaces, each sample was repeatedly cleaned with deionized water, ethanol and acetone, and then pulverized on a clean bench.

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Fig. 1. Study site, geological map [34] indicating sampling locality, lithologic column, and stratigraphic variations of total organic carbon (TOC) content and δ13Corg through the Bonarelli and underlying limestone. Detailed δ13Corg variations at the base of the Bonarelli. In δ13Corg profiles, data from organic-rich (N2% TOC) and -poor (b2%) sediment are indicated by filled and open squares, respectively. The 0 cm level corresponds to the base of the Bonarelli. A grey curve in the right panel indicates the simulated curves of δ13Corg using a box model (see text and Appendix B). The Cenomanian–Turonian boundary is approximately 1 meter above the top of the Bonarelli [36]. Arrows on the lithologic column indicate stratigraphic levels of samples analyzed for Pb isotopic composition.

2.2. Biomarker and carbon isotopic analyses Lipids were Soxhlet extracted with dichloromethane/ methanol (7:3, v/v). After saponification, the extracts were separated into neutral and acidic fractions and further separated into several sub-fractions by SiO2 column chromatography. Individual compounds were quantified and identified using an Agilent 6890 gas chromatograph (GC) equipped with an Agilent 5973 mass-selective detector. Carbon isotopic compositions of individual compounds were determined using a ThermoFinnigan Delta plus XP isotope-ratio mass spectrometer (IRMS) coupled to an Agilent 6890 GC. Determinations of total organic carbon contents and isotopic compositions were undertaken using the above IRMS connected to a Flash EA 1112 through a ConFlo III interface after removal of carbonates with HCl. The carbon isotopic compositions are expressed in delta (δ) notation against the Vienna

Peedee belemnite standard (VPDB) with a standard deviation better than ±0.2‰ (2σ). Pulverized sediments were Soxhlet extracted with dichloromethane/methanol (7:3, v/v). The total lipid extract was saponified with 0.5 M KOH/MeOH for two hours at 80 °C. Then the saponified extract was separated into neutral and acidic fractions and further separated into several sub-fractions by SiO2 column chromatography. Individual compounds were quantified and identified using an Agilent 6890 gas chromatograph (GC) equipped with DB-1 capillary column (30 m × 0.32 mm i.d., film thickness of 0.1 μm) and an Agilent 5973 mass-selective detector (MS). Helium was used as a carrier gas. The column oven temperature was programmed from 40° to 120 °C at 30 °C min− 1 and from 120° to 320 °C at 6 °C min− 1 with a final hold time of 25 min. Carbon isotopic compositions of individual compounds were determined using a ThermoFinnigan

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Delta plus XP isotope-ratio mass spectrometer (IRMS) coupled to an Agilent 6890 GC. The capillary column and column-oven temperature program were the same as that of the GC/MS. Carbon isotopic compositions are expressed in delta (δ) notation against the Vienna Peedee belemnite standard (VPDB). Repeated injection of a standard mixture of 13 n-alkanes indicated that analytical errors (1σ) of the n-alkanes are generally better than 0.3‰. Determinations of total organic carbon contents and isotopic compositions were undertaken using an on-line system of the above IRMS connected to a Flash EA 1112 Automatic Elemental Analyzer through a ConFlo III interface [37] after removal of carbonates with HCl. Analytical error (1σ) was estimated to be within 0.1‰ on the basis of repeated measurements of authentic and laboratory standards. 2.3. Lead isotopic compositions Lead isotopic compositions and quantifications of Pb, Th and U were performed partly following established techniques [38]. All chemical procedures were carried out in a class-1000 clean room and on class-100 clean benches. We used clay-rich samples from the Bonarelli interval (GCB04, 05, 06, 09, 11, 17 and 23; Fig. 1), from which sedimentary petrologic features were confirmed with an optical microscope, X-ray diffraction and electron microprobe (partly reported in [5]). To avoid contamination of Pb from pyrite, we selected pyrite-poor samples from the Bonarelli interval on the basis of element compositional mapping on thin sections by electron microprobe analysis (generally b0.9% as FeS2 concentration [5]). Thus we interpreted that Pb in these shale samples is mostly derived from silicate fraction. Rock powder containing from 200 to 400 ng Pb was weighed in a 7 mL PFA Teflon vial, and decomposed with 0.7 ml 20 M HF and 0.7 mL 8 M HBr at 130 °C. After drying at 140 °C, the samples were redissolved with 0.5 M HBr and then centrifuged. The supernatant was loaded onto a 0.1 mL anion exchange resin (Bio-Rad AG1-X8, 200–400 mesh, Br-form) packed in a TEF Teflon column. After elution of other elements with a 2.5 mL combined acid of 0.25 M HBr and 0.5 M HNO3, the Pb fraction was collected with 1 mL H2O. After drying at 120 °C, the Pb fraction was dissolved in 0.15 M HNO3. An appropriate volume of NIST SRM 997 Tl solution (0.01 mg/mL, 205Tl/203Tl = 2.3871) was added to obtain the final sample solution containing 200 ng/mL Pb and 20 ng/mL Tl to correct for instrumental mass fractionation. Pb isotopic compositions were measured using a ThermoFinnigan NEPTUNE

multi-collector ICP-MS with analytical uncertainties better than 0.090%-RSD (2σ) on the basis of repeated analyses of the NIST SRM981 standard solution. Operating conditions of the multi-collector ICP-MS have been fully described elsewhere [38]. Analytical accuracies represented as differences in standard data against literature values [38] were better than ±0.016%. Analytical reproducibilities of Pb isotopic compositions given as three separate measurements of unknown shale samples (GCB04, 05, 06, 09, 11, 17 and 23) were generally better than 0.036%-RSD (2σ; Table 2). The total procedural blank of Pb ranges between 50 and 80 pg. In contrast to the Bonarelli interval, clay-rich intervals are rare in the underlying limestone interval (Scaglia Bianca; Fig. 1). We used seven chalk samples from this interval (Fig. 1) for Pb isotopic analysis. In this case, Pb is mainly contained in carbonate and silicate phases. Chalk samples were decalcified using 1 M HNO3 to eliminate Pb in the carbonate fraction that was precipitated directly from seawater. Thus, we interpreted that most Pb in the residual fraction originated in the silicate fraction. Samples were prepared and measured in a similar manner to the black shales except for a Pb concentration in the sample solution of approximately 50 ng/mL, because of low Pb concentrations in the chalks. For these chalk samples, we used an APEX high-intensity inlet system (Elemental Scientific Inc.) to improve sensitivity. On the basis of Pb isotopic analysis with NIST SRM981 standard solution, we confirmed that analytical uncertainties when measured with the APEX were almost equal to those measured without the APEX. Concentrations of Pb, Th and U were analyzed using a PerkinElmer ELAN DRC-II ICP-MS to determine initial Pb isotopic ratios. We applied a standard addition method for Pb, Th and U quantification. Signal instability of the ICP was corrected with a Bi internal standard. Samples were dissolved with a 0.4 ml HF–HNO3 mixture (1:1, v/v). Analytical precision (RSD) and accuracy (compared to reference values [39]) on the basis of repeated measurements of Pb, Th and U concentrations in JA-3 rock reference material were better than 3%. Three aliquots were prepared per sample (except for GCB17, GCB23, GC99-03 and GC94-02) to determine uncertainties for measurement of U/Pb and Th/Pb ratios, which were better than 0.12 (2σ; Table 2). Uncertainties for initial Pb isotopic ratios are determined by error propagation of uncertainties in Pb isotopic measurements and that of U/Pb and Th/Pb analyses. Uncertainties in Pb isotopic measurements were estimated for each sample in the Bonarelli interval (n = 3), whereas those from underlying limestone (Scaglia Bianca) were

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Table 1 Concentrations and carbon isotopic compositions of organic compounds extracted from organic-rich samples in the Bonarelli GC010801 −1

n-Alkanes C18 Alkane C19 Alkane C20 Alkane C21 Alkane C22 Alkane C23 Alkane C24 Alkane C25 Alkane C26 Alkane C27 Alkane n-Alkanols C14 Alkanol C15 Alkanol C16 Alkanol C17 Alkanol C18 Alkanol C19 Alkanol C20 Alkanol C21 Alkanol C22 Alkanol C23 Alkanol C24 Alkanol n-Fatty acids C14 Fatty acid C15 Fatty acid C16 Fatty acid C17 Fatty acid C18 Fatty acid C19 Fatty acid C20 Fatty acid C21 Fatty acid C22 Fatty acid C23 Fatty acid C24 Fatty acid C25 Fatty acid C26 Fatty acid C27 Fatty acid C28 Fatty acid C30 Fatty acid Hopanoids 17β(H),21β(H)−Hopane 17β(H),21β(H)-Homohopane 17β(H),21β(H)-Homohopanoic acid 17β(H),21β(H)-Bishomohopanoic acid Sterols Cholesterol Total organic carbon Stratigraphic level

GCBR09 −1

GC010809

δ C (‰)

(ug g C )

δ C (‰)

(ug g C )

δ C (‰)

(ug g C− 1)

δ13C (‰)

n.d. n.d. 0.19 0.36 0.44 0.58 0.30 0.27 0.14 0.18

n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

0.78 1.8 1.9 1.9 2.0 2.4 2.2 2.4 1.7 1.1

− 28.6 − 29.4 − 28.9 − 29.3 − 29.1 − 29.0 − 28.9 − 29.3 n.d. n.d.

n.d. n.d. 0.28 0.37 0.43 0.55 0.23 0.25 0.28 0.32

n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

n.d. 0.47 1.0 1.7 1.5 2.5 0.97 2.4 0.46 1.9

n.d. −27.3 −28.9 −27.7 −26.5 −27.4 −28.5 −25.2 n.d. n.d.

0.67 n.d. 2.3 0.58 3.7 0.32 2.6 0.33 0.81 0.12 0.37

n.d. n.d. − 23.0 n.d. − 26.5 n.d. − 25.8 n.d. − 24.6 n.d. − 25.4

3.4 8.2 14 4.7 12 3.2 12 3.2 5.9 1.4 2.0

n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

2.6 1.1 6.0 2.4 10 1.2 12 0.59 3.3 0.55 1.0

n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

1.7 1.1 2.1 0.92 2.2 0.70 1.1 0.71 0.62 0.19 0.34

n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a. n.a.

n.d. n.d. 11 1.5 10 0.99 2.0 0.72 1.0 0.33 0.55 0.092 0.16 n.d. n.d. n.d.

n.d. n.d. − 27.0 − 27.3 − 27.0 − 29.0 − 28.0 − 29.8 − 29.2 − 29.6 − 28.6 − 29.4 − 28.6 n.d. n.d. n.d.

34 41 90 42 51 29 33 30 37 28 32 16 19 5.4 11 10

− 28.8 − 28.9 − 27.3 − 30.1 − 29.2 − 30.4 − 30.1 − 30.5 − 29.9 − 30.3 − 29.7 − 30.5 − 29.6 n.d. − 29.6 − 30.6

n.d. n.d. 7.0 1.1 14 0.80 1.6 0.48 1.2 0.27 1.5 0.26 0.44 n.d. 0.066 0.081

n.d. n.d. − 27.9 − 26.1 − 26.4 − 27.1 − 26.8 − 27.7 − 27.6 − 27.3 − 27.0 − 27.6 − 27.8 n.d. − 28.2 − 27.7

23 39 65 47 47 31 24 19 22 16 20 7.1 16 3.8 9.0 6.7

−27.6 −26.7 −26.5 −27.0 −27.9 −27.2 −26.9 −27.3 −25.8 −26.9 −25.4 −26.7 −25.1 n.d. −25.4 −26.9

0.93 3.9 1.0 3.0

− 29.4 − 29.8 − 30.3 − 30.7

3.2 15 38 138

n.d. n.d. − 29.8 − 29.9

2.1 10 1.9 11

− 27.1 − 27.7 − 28.7 − 28.5

4.18 12.3 20.7 74.2

−24.8 −26.9 −26.4 −27.2

6.6 2.0 wt.% 0 cm

− 25.1 − 24.8

11 6.4 wt.% 21 cm

n.a. − 26.0

25 3.3 wt.% 43 cm

n.a. − 22.7

n.d. 22 wt.% 76 cm

n.a. −22.7

13

−1

GCB21

(ug g C )

13

13

Abbreviations: n.d., not detected; n.a., not analyzed.

not determined for each sample because they were not analyzed three times independently (see above). For these chalk samples, we adopted the largest standard deviations in Pb isotopic measurements (results of GCB23) as the

representative standard deviations of Pb isotopic measurements. Although the concentration of Pb in analytical solution was different between shale and chalk samples, we confirmed that precision and accuracy as well as

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signal intensities were almost equal when the high-intensity inlet system was used for chalk solutions. Thus, we consider that standard deviations in Pb isotopic measurements for the shale solution are applicable to chalk samples. On the basis of triplicate measurements of separately prepared sample aliquots, uncertainties in determination of U/Pb and Th/Pb ratios were determined for each sample except for two shale samples (GCB17 and GCB23) and two chalk samples (GC99-03 and GC94-02). We adopted standard deviations in analysis of GC95-04 as the representative precisions for these four samples, because GC95-04 has similar U/Pb and Th/Pb ratios as well as signal intensities of Pb, Th and U to those in these four samples, and the standard deviations of GC95-04 are the largest among measured samples. 3. Results The 104-cm-thick Bonarelli black shale is characterized by millimeter-scale alternations of organic carbonrich and -poor intervals, whereas adjacent limestone is consistently depleted in organic carbon (Fig. 1 and Appendix A). The organic-rich intervals roughly correspond to black laminae. Values of δ13Corg shift abruptly from −24 to −27‰ at the base of the Bonarelli (Fig. 1 and Appendix A). The 3‰ decrease of δ13Corg occurs within the basal 2 cm. The rapid negative shift is followed by a gradual positive excursion in the basal 40 cm of the Bonarelli. Relative abundances and carbon isotopic compositions of organic compounds including n-alkanes, alcohols, fatty acids, hopanoids, and sterols (Table 1) indicate that stratigraphic variations of compound-specific δ13C records parallel patterns in the δ13Corg profile (Fig. 2). Measured isotopic ratios of Pb, concentrations of Pb, Th and U, and initial isotopic compositions (93.6 Ma) calculated on the basis of U/Pb and Th/Pb ratios are listed in Table 2. Stratigraphic variations of initial Pb isotopic compositions and cross plots among 206Pb/ 204 Pb, 207Pb/204 Pb and 208 Pb/204Pb ratios (Fig. 3) indicate a shift at the base of the Bonarelli black shale, coincident with the sharp negative shift of δ13Corg, and a more prominent systematic isotopic shift in 208Pb/ 204 Pb ratios than in the 206 Pb/204Pb and 207 Pb/204Pb ratios, respectively. 4. Discussion The mean sedimentation rate of the Bonarelli has traditionally been estimated to be 1.3 to 3.1 m my− 1 [32,33], which yields a duration of 350 to 800 ky for the Bonarelli interval. As described above, the Bonarelli

Fig. 2. Carbon isotopic ratios of total organic carbon (δ13Corg) and organic compounds (n-alkanols, n-alkanes, n-fatty acids, hopanoids, and sterols) extracted from four organic carbon-rich intervals in the Bonarelli.

black shale roughly corresponds to a biostratigraphic interval between the top of the R. cushmani Zone and the bottom of the Q. gartneri Zone [35,36]. A recent study addressing timescales of the OAE-2 interval in the WIS section, employing cyclostratigraphy of the Bridge Creek limestone in conjunction with 40Ar/39Ar dating of bentonite, calculated the duration of this biostratigraphic interval to be approximately 540 ky [40]. A similar duration has been reported by investigations of foraminiferal biostratigraphy from the WIS [41] and cyclostratigraphy of a North African section [42]. Since the previously estimated duration of the Bonarelli (350 to 800 kyr [32,33]) encompasses these recent estimates from other basins [40–42], we consider that a sedimentation rate of 1.3–3.1 m my− 1 is realistic and apply it in the following discussion. The sharp negative shift of δ13Corg in the basal 3 cm of the Bonarelli occurred within 7–15 ky. This rapid negative shift is followed by a gradual positive excursion in the basal 40 cm of the Bonarelli, the duration of which corresponds to 130–310 kyr. The isotopic excursion is paralleled by δ13C records of various organic compounds (Fig. 2). Since relative abundances of these major compounds do not vary significantly among samples (Table 1), δ13Corg variations do not appear to reflect changes in relative contributions of 13C-depleted organisms. Instead, they mainly reflect isotopic variations of dissolved inorganic carbon in the surface ocean (δ13CDIC). This interpretation is supported by the duration of gradual rebound (130 to 310 ky) after the rapid negative shift of δ13Corg in the Bonarelli, which is

Table 2 Measured Pb isotopic ratios, concentrations of Pb, Th and U, weight ratios of U/Pb and Th/Pb, and age corrected Pb isotopic composition (t = 93.6 Ma) in the Bonarelli black shale and underlying limestones 207

208

Pb/204Pb ± 2σ

Pb/204Pb ± 2σ

Pb/204Pb ± 2σ Pb ± 2σ

Measured ratio

Measured ratio

Measured ratio

(μg g− 1) (%−RSD)

(μg g− 1) (μg g− 1) (%-RSD) (%-RSD)

Bonarelli black shale GCB23 86.2 19.2064 ± 0.0146 GCB17 50.1 19.3298 ± 0.0036 GCB11 27.6 19.2901 ± 0.0065 GCB09 19.3 18.9340 ± 0.0038 GCB06 12.0 18.9355 ± 0.0035 GCB05 8.5 18.8741 ± 0.0007 GCB04 4.9 18.9142 ± 0.0060

15.6754 ± 0.0036 15.6875 ± 0.0028 15.6854 ± 0.0011 15.6582 ± 0.0018 15.6555 ± 0.0017 15.6522 ± 0.0009 15.6591 ± 0.0024

38.7352 ± 0.0139 38.7548 ± 0.0077 38.7784 ± 0.0056 38.7376 ± 0.0060 38.7301 ± 0.0050 38.7140 ± 0.0025 38.7289 ± 0.0066

8.44

2.06

5.54

0.66

9.58

2.02

6.50

0.68

Cenomanian limestone GC99-02 − 1.5 19.7869

15.7092

GC99-03 GC99-04b

− 2.5 19.3884 − 3.5 19.2889

GC99-04a

Pb/204Pb ± 2σ a

Pb/204Pb ± 2σ a

Pb/204Pb ± 2σ a

206

207

208

Initial ratio (93.6 Ma)

Initial ratio (93.6 Ma)

Initial ratio (93.6 Ma)

0.24

18.5256 ± 0.0824 b

15.6428 ± 0.0053 b

38.6543 ± 0.0435 b

0.21

18.6260 ± 0.0812 b

15.6538 ± 0.0048 b 38.6850 ± 0.0419 b

20.7 ± 0.63 3.73 ± 4.1 13.3 ± 2.5 0.64 ± 0.02 19.4 ± 13 1.45 ± 13 2.17 ± 14 0.11 ± 0.02 8.12 ± 5.1 1.88 ± 3.4 2.92 ± 2.6 0.36 ± 0.02 18.4 ± 19 3.45 ± 11 6.57 ± 6.7 0.36 ± 0.05 13.1 ± 7.6 1.68 ± 8.3 4.06 ± 8.5 0.31 ± 0.01

0.18 ± 0.01 0.07 ± 0.01 0.23 ± 0.01 0.19 ± 0.02 0.13 ± 0.00

18.6249 ± 0.0217

15.6536 ± 0.0015

38.7186 ± 0.0060

18.8182 ± 0.0176

15.6527 ± 0.0019

38.7128 ± 0.0065

18.5623 ± 0.0176

15.6376 ± 0.0019

38.6532 ± 0.0058

18.5023 ± 0.0483

15.6344 ± 0.0025

38.6516 ± 0.0060

18.5938 ± 0.0125

15.6438 ± 0.0025

38.6866 ± 0.0067

39.2126

0.25 ± 21

0.24 ± 11

18.4841 ± 0.0481 c

15.6468 ± 0.0042 c

38.9150 ± 0.0159 c

15.6934 15.6813

39.3272 39.2088

0.61 0.44 ± 1.5

0.61 0.61 0.37 ± 4.9 0.37 ± 3.5

− 3.5 19.3538

15.6880

39.2731

0.38 ± 8.8

0.36 ± 5.7 0.35 ± 7.6

GC99-06

− 5.5 19.6110

15.7045

39.5997

0.25 ± 2.5

0.38 ± 3.7 0.32 ± 4.7

GC95-04

− 36.5 19.1625

15.6881

39.2587

0.70 ± 19

0.72 ± 9.3 0.51 ± 12

GC94-02

− 41.5 19.3964

15.7095

39.4996

0.49

0.65

0.90 ± 0.02 1.01 0.69 ± 0.02 0.97 ± 0.04 1.54 ± 0.02 1.04 ± 0.12 1.33

Level (cm)

Th ± 2σ

U ± 2σ

0.35 ± 11

0.45

U/Pb ± 2σ Th/Pb ± 2σ

1.26 ± 0.04 1.00 0.69 ± 0.02 0.95 ± 0.02 1.32 ± 0.03 0.73 ± 0.08 0.91

18.3471 ± 0.0824 b, c 15.6435 ± 0.0053 b, c 38.9936 ± 0.0434 b, c 18.5693 ± 0.0216 c 15.6469 ± 0.0037c 38.9812 ± 0.0161 c 18.3734 ± 0.0270 c

15.6410 ± 0.0038 c

38.9524 ± 0.0201 c

18.2450 ± 0.0331 c

15.6391 ± 0.0039 c

39.0880 ± 0.0150 c

18.4021 ± 0.0824 c

15.6517 ± 0.0053 c

38.9151 ± 0.0434 c

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206

Sample

18.4517 ± 0.0824 b, c 15.6643 ± 0.0053 b, c 39.0577 ± 0.0434 c, c

a

Standard deviations of initial Pb isotopic ratios were determined by error propagation of Pb isotopic measurements and determination of U/Pb and Th/Pb ratios. The largest standard deviation associated with U/Pb and Th/Pb measurements (GC96-04) was adopted as the representative standard deviations for samples with single measurement of U/Pb and Th/Pb ratios. c The largest standard deviation associated with Pb isotopic measurement (GCB23) was adopted as the representative standard deviations for samples with single measurement of Pb isotopic ratios. b

217

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Fig. 3. Depth profiles of 206Pb/204Pb (a), 207Pb/204Pb (b), and 208Pb/204Pb (c), and cross plots of 206Pb/204Pb vs. 207Pb/204Pb (d), and 206Pb/204Pb vs. Pb/204Pb (e) in the Bonarelli (red squares) and underlying Cenomanian limestone (blue squares). All data are calculated at 93.6 Ma [3]. Error bars indicate 2σ. For comparison, Pb isotopic ratios of basaltic rock from the Caribbean (at 88–95 Ma) [26–29], Madagascar (at 88 Ma) [30] flood basalts, and from MORB (at present) from the Atlantic and Pacific (light gray) and Indian (dark gray) oceans [31] are also shown in (d) and (e). The 0 cm level corresponds to the base of the Bonarelli. A northern hemisphere reference line (NHRL) is shown in panel (d). Meshed domains in (d) and (e) indicate likely end member of Pb isotopic ratios of volcanogenic silicate minerals from the LIP.

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roughly 2–3 times the residence time of dissolved inorganic carbon in the global ocean [43]. In general, three times the residence time is necessary for a 95% achievement of steady state following removal of a perturbation. From 40 to 104 cm in the Bonarelli, δ13Corg values of organic-rich sediment are relatively constant, with a mean value of − 23.4‰ (Fig. 1), suggesting steady state conditions during that time. Although 3‰ negative shifts are also observed in several thin organic carbonpoor intervals, e.g., 70 and 85 cm (Fig. 1), these shifts return abruptly to pre-shift values of − 24‰, which may be explained by changes in sea surface ecology or diagenetic alteration of isotopic signatures. Values of δ13Corg in the underlying limestone (−24.1‰) are significantly higher than those in the lower Cenomanian interval of this section (−25.3 ± 0.8‰), suggesting that the widely recognized “positive excursion” due to enhanced burial rates of organic carbon during OAE-2 [2,4,21–25,36] had started before the onset of Bonarelli deposition. One possibility for the sharp negative spike is local upwelling of deep or intermediate water that contained isotopically light dissolved inorganic carbon [19]. Another possibility is that this sharp negative shift reflects a global signal. Slight negative spikes of δ13Corg and carbonate carbon isotopic compositions (δ13Ccarb) within the increasing C–T positive excursion are observed in

OAE-2 sediments from various paleogeographic sites including southern England (“trough interval” in Gale et al. [21]), the southeastern North Atlantic (isotopic interval “B” in Erbacher et al. [22]), the western Pacific (Japan) [23,24], and the WIS of North America [25,36] (Fig. 4). A significant difference in lengths of carbon isotopic excursions among sites (Fig. 4) reflects that of sedimentation rates: the Italian section in this study (1.3–3.1 m my− 1 [32,33]), southern England (15–31 m my− 1 [44]), the WIS in the North America (9–13 m my− 1 [41]) and the Demerara Rise (15 m my− 1 [22]). The length of the carbon isotopic excursion in each section is roughly proportional to the average sedimentation rate in each section, suggesting a contemporaneous carbon isotopic excursion, and supporting the idea that the δ13C excursion including the negative isotopic spike is a global signature. Importantly, negative spikes, at least in the WIS and South England sections are commonly observed around the last occurrence levels of Rotalipora cushmani and Axopodorhabdus albianus [21,36], key marker species of planktonic foraminifera and coccolithophorids, respectively. Because the base of the Bonarelli also corresponds to this biostratigraphic level [35,36], the negative shift of δ13Corg documented here represents a global isotopic variation in the ocean carbon reservoir. The biostratigraphic level associated with the carbon isotopic negative

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Fig. 4. Correlation of the carbon isotopic record from Bonarelli intervals to records from sections in southern England [21], the Demerara Rise in the southeastern North Atlantic (ODP Site 1261) [22], northern Japan [23] and the WIS in North America [25]. The shaded interval marks the sharp negative spike in carbon isotopic records in the sections. The broken line indicates the Cenomanian–Turonian boundary. Note that the C–T boundary is located 1 m above the Bonarelli in the Italian section [36]. Average sedimentation rates estimated previously are also shown in each panel (see text).

spike is important because 1) it corresponds to the onset of the mass-extinction level of benthic and deep-dwelling planktonic foraminiferal species [45], and 2) subsequently organic-rich sediment expanded throughout the global ocean [4]. These phenomena could be related to the negative δ13C event. We attribute the negative shift of δ13CDIC to a substantial increase in input of isotopically light carbon such as volcanogenic CO2 from the mantle (−5‰ [43,46] or more depleted down to −23‰ [46,47]) or methane (approximately −60‰ for biogenic CH4 [48] and −30‰ for thermogenic CH4 [49], respectively) from gas hydrates [6] or a geothermal reservoir [49]. Stratigraphic variations in initial Pb isotopic ratios in the silicate fraction show apparent shifts at the boundary between the Bonarelli and underlying limestone (Fig. 3). A significant change in source of silicate minerals is one of the likeliest causes for Pb isotopic shifts at the onset of the Bonarelli black shale. It is noteworthy that shifts in Pb isotopic compositions at the onset of Bonarelli deposition (Fig. 3d and e) are displaced to the fields of

basaltic rocks from the Caribbean [26–29] and Madagascar [30] LIPs erupted around OAE-2, in addition to mid-ocean ridge basalts (MORB) [31]. It clearly indicates a rapid increase in relative supply of volcanogenic silicate minerals at the onset of the Bonarelli black shale. The Pb isotopic shift may have resulted from a local change in transportation process (e.g., wind direction) and source provinces of detrital materials. However, because this site was located in a pelagic environment far from land [50], we consider that such local changes are less plausible as causes for the Pb isotopic change. Due to sluggish ocean circulation during mid-Cretaceous time [51], hydrothermal activity from deepwater eruptions of LIPs or MORBs would probably not disperse volcanic material far from their sources. Massive subaerial or very shallow water volcanism is most likely necessary to supply such a large amount of volcanogenic material. Although it is difficult to determine the accurate Pb isotopic ratios of silicate minerals from LIPs as an endmember for mixing calculations, we have assumed a range of endmembers

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(meshed areas in Fig. 3d and e), that encompasses values from basalts of the Caribbean LIP [26–29]. On the basis of Pb isotopic compositions of 1) the assumed endmembers of rocks from the LIPs and 2) the pre-Bonarelli (background) sediment represented by average Pb isotopic ratios in the underlying limestone, we estimate that approximately 20–40% of Pb in the Bonarelli was supplied as volcanogenic silicate minerals. Because the base of the Bonarelli corresponds to 94.6–94.1 Ma, assuming published sedimentation rate estimates [32,33], Bonarelli initiation coincides well with Caribbean LIP formation, which has most recently been determined to be 95.1–92.2 Ma on the basis of 40 Ar/39Ar dating of whole rock samples [13]. The Madagascar LIP has an advantage over the Caribbean LIP in that the former erupted primarily subaerially, which would have enabled release of huge amounts of CO2 and particulate materials into the atmosphere. However, we note that the time envelope of onshore Madagascar LIP formation has been estimated to be somewhat later (91.6–83.7 Ma; 40Ar/39Ar [17] and U–Pb [18] dating) than that of OAE-2, although interpreted volcanic basement of the extensive submarine Madagascar Ridge [52] remains unsampled and undated. Extensive aeolian dust produced by massive volcanism should induce a significant cooling due to reduction of solar insolation. In the case of the Laki eruption in Iceland (AD1783–1784), the second-largest known eruption of mafic magma during historic times, volcanic haze produced by the eruption led to a 1.3 °C cooling of annual mean temperature for 2–3 years [53]. Because the Caribbean and Madagascar LIPs [11] expelled magmas as much as five orders of magnitude more voluminous than that of Laki [53], far larger amounts of volcanic material were released, which would have resulted in a substantial, but short-term cooling event on a global scale. The short-term cooling would have been followed by global warming due to the greenhouse effect induced by elevated levels of atmospheric CO2. Sea surface temperatures (SST) of up to 25 °C, even in the ancestral Arctic Ocean [54], have been reported, with maximum temperatures at the C–T boundary. More recently, Forster et al. [55] have found an abrupt increase in SST up to 36 °C at the onset of OAE-2, consistent with our interpretation. Using a one-box kinetic model of the atmosphere-ocean system (Appendix B), we estimate that 7 to 12 × 104 Gt CO2 with an average δ13C of − 5‰ are required for a 3‰ negative shift of δ13CDIC. The simulated δ 13 C org values resemble those of the Bonarelli (Fig. 1b). If this degassing was completed within 7–15 ky, the mean discharge rate of excess CO2

during massive volcanism would have been 4.5–17.0 Gt CO2 y− 1, which corresponds to the emission rate of present anthropogenic CO2 (26.0 Gt CO2 y− 1). It should be noted that this value is the maximum estimate. A much more negative δ13C value (ca. − 23‰) has been proposed for CO2 from the mantle on the basis of bulk δ13C measurements of continental flood basalts (e.g., Kerguelen and Deccan Traps) and oceanic island basalts (e.g., Hawaii and Iceland) [47], as well as mantle xenoliths [46]. If this were the case for OAE-2, an approximately one order of magnitude smaller amount of CO2 is sufficient for the 3‰ negative shift. In the following discussion, however, we assumed the largest CO2 amount (with δ13C of −5‰) to evaluate if such a massive amount of CO2 could be supplied by a single LIP emplacement. The Caribbean and Madagascar LIPs each encompass approximately 4.5 × 106 km3 of basalt [11,12], which corresponds to a total mass of 1.3 × 107 Gt for each, assuming a mean density of 2.9 g cm− 3. If 20% of the total basaltic mass (2.6 × 106 Gt) of one of the provinces was erupted at that time, and if all of the CO2 calculated above (7–12 × 104 Gt) were released to the atmosphere-ocean system, the original magma would have contained 2.6– 4.6 wt% CO2. Recently, Thordarson and Self [56] have suggested that the source magma of the Columbia River basalt contained as much as 2.0 wt% CO2, which is comparable to our calculated values. Furthermore, Keppler et al. [57] have proposed that voluminous melting of carbonate-rich mantle can rapidly supply a large amount of CO2 into the atmosphere and ocean. In light of this work, subaerial eruption of each individual LIP could explain rapid degassing of anomalously large amounts of CO2 (on the order of 104 Gt). Isotopic shifts of both carbon and lead observed at the base of Bonarelli can be explained simply by a massive release of volcanogenic silicate minerals and CO2 associated with subaerial volcanism. Elevated SSTs due to massive CO2 degassing from subaerial eruptions would have prevented or slowed the formation of deep water, resulting in oceanic stratification [51], subsequent oceanic anoxia, and mass-extinction of deep-dwelling foraminifers [45]. Abrupt enhancement of CO2 in the ocean-atmosphere system would have altered ocean chemistry principally by lowering the pH and carbonate ion content in seawater [58]. These changes would be partially neutralized by a transient rise in the level of the calcite compensation depth (CCD), resulting in the widespread dissolution of sea-floor carbonate, which has been observed as short-term decrease in CaCO3 concentration in sediments deposited during Paleocene–Eocene boundary [58]. An abrupt decrease in concentration of CaCO3 is reported from OAE-2 intervals in central Italy

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[5,59], implying a tentative shoaling of the CCD during OAE-2. Our data do not fully exclude the alternative hypothesis of massive methane release for the negative spike in carbon isotopic composition. Recently, it has been suggested that thermogenic methane, which includes those degassed directly from the mantle, synthesized inorganically, or thermally decomposed from organic matter, is an important constituent of the geothermal reservoir [49]. Thus, the rapid release of thermogenic methane via a magmatic event would be an alternative possibility for the 3‰ negative shift of the ocean-atmosphere carbon budget, because thermogenic methane is depleted in 13 C to as low as − 30‰. However, determination of carbon species in fluid inclusions in mantle xenoliths has revealed that CO2 and carbonates are major carbon species in mantle xenoliths, whereas CH4 is only a trace component (reviewed by Deines [46] and references therein). Therefore, so far we interpret that thermogenic CH4 is less likely than CO2 to have been the cause of the 3‰ negative spike. If the methane hypothesis is valid, however, similar consequences, i.e., acceleration of global warming and subsequent oceanic stratification and shoaling of the CCD, would have ensued following methane injection because most methane would have been oxidized to CO2 shortly after release into the atmosphere-ocean system [58,60]. A negative excursion in strontium isotopic composition [61] and rare-metal anomalies (e.g., Sc and Cr) [13] have been observed in OAE-2 sediment, suggesting a substantial increase in magmatic activity during this event, possibly related to LIP formation. Since deep submarine volcanism cannot release large amounts of volcanic ash and CO2 to the hydrosphere or atmosphere, due to the confining pressure of water, geochemical signatures recorded in marine sediments should differ depending upon dominantly deep submarine or subaerial LIP emplacement. 5. Conclusions Ultra-high-resolution records of δ13Corg indicate a sharp negative shift at the base of the Bonarelli black shale. By comparing the δ13Corg record with carbon isotopic records from other sites and with those of extractable lipid compounds, we interpret that the sharp negative δ13Corg shift reflects a rapid decrease in isotopic composition of dissolved inorganic carbon in the global ocean. At the same stratigraphic level, we found systematic shifts in Pb isotopic compositions of the silicate fraction, which suggest an abrupt increase in relative supply of silicate minerals from volcanism. Although several scenarios may

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explain these geochemical signatures, a massive subaerial eruption potentially associated with LIP formation can explain them perhaps the most simply. OAEs in other geological time intervals may also have been triggered by LIPs, as suggested by age similarities, e.g., OAEs of early Aptian and early Toarcian age are temporally close to the ages of Ontong Java and Karoo flood basalts, respectively [10–12]. Further dense sampling and geochemical analyses of sediment sections containing black shales, geochemical fingerprinting of LIP events, and modelling of environmental change will contribute strongly to our understanding of interactions and feedbacks between the solid Earth and the Earth's environment. Acknowledgements We thank T. Ishikawa, K. Suzuki, H. Suga, A. Taira, Y. Sano, R. Tada, E. Tajika, K. Kaiho, M. Tejada and MUD Project members for helpful comments on the manuscript, and S. Kiyokawa, R. Coccioni and I. Premoli Silva for their help with sampling. The manuscript has benefited from reviews by C. Chauvel and an anonymous reviewer, and editing by M. Delaney. The senior author is supported by a Japan Society for the Promotion of Science (JSPS) post-doctoral fellowship. Appendix A. Supplementary data Supplementary data associated with this article can be found, in the online version, at doi:10.1016/j. epsl.2007.01.027. References [1] S.O. Schlanger, H.C. Jenkyns, Cretaceous oceanic anoxic events: causes and consequences, Geol. Mijnb. 55 (1976) 179–184. [2] S.O. Schlanger, M.A. Arthur, H.C. Jenkyns, P.A. Scholle, The Cenomanian–Turonian Oceanic Anoxic Event, I. Stratigraphy and distribution of organic carbon-rich beds and the marine δ13C excursion, in: J. Brooks, A.J. Fleet (Eds.), Marine Petroleum Source Rocks, Geol. Soc. London Spec. Publ., vol. 26, 1987, pp. 371–399. [3] F.M. Gradstein, J.G. Ogg, A.G. Smith, A Geologic Time Scale, Cambridge University Press, Cambridge, UK, 2004 589 p. [4] J. Kuroda, N. Ohkouchi, Implication of spatiotemporal distribution of black shales deposited during Cretaceous Oceanic Anoxic Event-2, Paleontol. Res. 10 (2006) 345–358. [5] J. Kuroda, N. Ohkouchi, T. Ishii, H. Tokuyama, A. Taira, Laminascale analysis of sedimentary components in Cretaceous black shales: Paleoceanographic implications for oceanic anoxic events, Geochim. Cosmochim. Acta 69 (2005) 1479–1494. [6] D.J. Beerling, M.R. Lomans, D.R. Gröcke, On the nature of methane gas-hydrate dissociation during the Toarcian and Aptian oceanic anoxic events, Am. J. Sci. 302 (2002) 28–49.

222

J. Kuroda et al. / Earth and Planetary Science Letters 256 (2007) 211–223

[7] H.C. Jenkyns, Evidence for rapid climate change in the Mesozoic– Palaeogene greenhouse world, Philos. Trans. R. Soc. Lond., A 361 (2003) 1885–1916. [8] C.W. Sinton, R.A. Duncan, Potential links between ocean plateau volcanism and global ocean anoxia at the Cenomanian–Turonian boundary, Econ. Geol. 92 (1997) 836–842. [9] A.C. Kerr, Oceanic plateau formation: a cause of mass extinction and black shale deposition around the Cenomanian–Turonian boundary, J. Geol. Soc. (Lond.) 155 (1998) 619–626. [10] R.L. Larson, E. Erba, Onset of the Mid-Cretaceous greenhouse in the Barremian–Aptian: igneous events and the biological, sedimentary, and geochemical responses, Paleoceanography 14 (1990) 663–678. [11] O. Eldholm, M.F. Coffin, Large igneous provinces and plate tectonics, in: M.A. Richards, R.G. Gordon, R.D. van der Hilst (Eds.), The History and Dynamics of Global Plate Motions, Geophys. Monogr., vol. 121, AGU, Washington, DC, 2000, pp. 309–326. [12] M.F. Coffin, O. Eldholm, Large igneous provinces, in: R.C. Selley, R. Cocks, I.R. Plimer (Eds.), Encyclopedia of Geology, Elsevier, Oxford, 2004, pp. 315–323. [13] L.J. Snow, R.A. Duncan, T.J. Bralower, Trace element abundances in the Rock Canyon Anticline, Pueblo, Colorado, marine sedimentary section and their relationship to Caribbean plateau construction and oxygen anoxic event 2, Paleoceanography 20 (2005), doi:10.1029/2004PA001093. [14] J.J. Mahoney, M. Storey, R.A. Duncan, K.J. Spencer, M.S. Pringle, Geochemisty and geochronology of the Ontong Java Plateau, in: M.S. Pringle, W.W. Sager, W. Sliter, S. Stein (Eds.), The Mesozoic Pacific. Geology, Tectonics, and Volcanism, Geophys. Monogr., vol. 77, AGU, Washington, DC, 1993, pp. 233–261. [15] M.L.G. Tejada, J.J. Mahoney, R.A. Duncan, M.P. Hawkins, Age and geochemistry of basement and alkalic rocks of Malaita and Santa Isabel, Solomon Islands, southern margin of Ontong Java Plateau, J. Petrol. 37 (1996) 361–394. [16] M.L.G. Tejada, J.J. Mahoney, C.R. Neal, R.A. Duncan, M.G. Petterson, Basement geochemistry and geochronology of Central Malaita, Solomon Islands, with implications for the origin and evolution of the Ontong Java Plateau, J. Petrol. 43 (2002) 449–484. [17] M. Storey, J.J. Mahoney, A.D. Saunders, R.A. Duncan, S.P. Kelly, M.F. Coffin, Timing of hot-spot related volcanism and the breakup of Madagascar and India, Science 267 (1995) 852–855. [18] T.H. Torsvik, R.D. Tucker, L.D. Ashwal, E.A. Eide, N.A. Rakotosolofo, M.J. de Wit, Late Cretaceous magmatism in Madagascar: palaeomagnetic evidence for a stationary Marion hotspot, Earth Planet. Sci. Lett. 164 (1998) 221–232. [19] A.P. Menegatti, H. Weissert, R.S. Brown, R.V. Tyson, P. Farrimond, A. Strasser, M. Caron, High resolution δ13C stratigraphy through the early Aptian “Livello Selli” of the Alpine Tethys, Paleoceanography 13 (1998) 530–545. [20] M. Dumitrescu, S.C. Brassell, Compositional and isotopic characteristics of organic matter for the early Aptian oceanic anoxic event at Shatsky Rise, ODP Leg 198, Palaeogeogr. Palaeoclimatol. Palaeoecol. 235 (2006) 168–191. [21] A.S. Gale, H.C. Jenkyns, W.J. Kennedy, R.M. Corfield, Chemostratigraphy versus biostratigraphy: data from around the Cenomanian–Turonian boundary, J. Geol. Soc. (Lond.) 150 (1993) 29–32. [22] J. Erbacher, O. Friedrich, P.A. Wilson, H. Birch, J. Mutterlose, Stable organic carbon isotope stratigraphy across Oceanic Anoxic Event 2 of Demerara Rise, western tropical Atlantic, Geochem. Geophys. Geosyst. 6 (2004), doi:10.1029/2004GC000850.

[23] T. Hasegawa, T. Saito, Global synchroneity of a positive carbon isotope excursion at the Cenomanian/Turonian boundary: validation by calcareous microfossil biostratigraphy of the Yezo Group, Hokkaido, Japan, The Island Arc 2 (1993) 181–191. [24] T. Hasegawa, T. Hatsugai, Carbon-isotope stratigraphy and its chronostratigraphic significance for the Cretaceous Yezo Group, Kotanbetsu area, Hokkaido, Japan, Paleontol. Res. 4 (2000) 95–106. [25] A.R. Bowman, T.J. Bralower, Paleoceanographic significance of high-resolution carbon isotope records across the Cenomanian– Turonian boundary in the western interior and New Jersey coastal plain, USA, Mar. Geol. 217 (2005) 305–321. [26] R.V. White, J. Tarney, A.C. Kerr, A.D. Saunders, P.D. Kempton, M.S. Pringle, G.T. Klaver, Modification of an oceanic plateau, Aruba, Dutch Caribbean: implications for the generation of continental crust, Lithos 46 (1999) 43–68. [27] F. Hauff, K. Hoernle, G. Tilton, D.W. Graham, A.C. Kerr, Large volume recycling of oceanic lithosphere over short time scales: geochemical constraints from the Caribbean Large Igneous Province, Earth Planet. Sci. Lett. 174 (2000) 247–263. [28] A.C. Kerr, J. Tarney, P.D. Kempton, P. Spadea, A. Nivia, G.F. Marriner, R.A. Duncan, Pervasive mantle plume head heterogeneity: Evidence from the late Cretaceous Caribbean–Colombian oceanic plateau, J. Geophys. Res. 107 (2002), doi:10.1029/2001JB000790. [29] P.M.E. Thompson, P.D. Kempton, R.V. White, A.D. Saunders, A.C. Kerr, J. Tarney, M.S. Pringle, Elemental, Hf–Nd isotopic and geochronological constraints on an island arc sequence associated with the Cretaceous Caribbean plateau: Bonaire, Dutch Antilles, Lithos 74 (2004) 91–116. [30] M. Storey, J.J. Mahoney, A.D. Saunders, Cretaceous basalts in Madagascar and the transition between plume and continental lithosphere mantle sources, in: J.J. Mahoney, M.F. Coffin (Eds.), Large Igneous Provinces: Continental, Oceanic, and Planetary Flood Volcanism, Geophys. Monogr., vol. 100, AGU, Washington, DC, 1997, pp. 95–122. [31] A.W. Hofmann, Mantle geochemistry: the message from oceanic volcanism, Nature 385 (1997) 219–229. [32] N. Ohkouchi, K. Kawamura, Y. Kajiwara, E. Wada, M. Okada, T. Kanamatsu, A. Taira, Sulfur isotopic records around Livello Bonarelli (Northern Apennines, Italy) black shale at the Cenomanian–Turonian boundary, Geology 27 (1999) 535–538. [33] M.A. Arthur, I. Premoli Silva, Development of widespread organic carbon-rich strata in the Mediterranean Tethys, in: S.O. Schlanger, M.B. Cita (Eds.), Nature and Origin of Cretaceous Carbon-Rich Facies, Elsevier, New York, 1982, pp. 7–54. [34] S. Cresta, S. Monechi, G. Parisi, Mesozoic–Cenozoic stratigraphy in the Umbria–Marche Area. Geological field trips in the Umbria–Marche Apennines (Italy), Mem. Descr. Carta Geol. Ital. 39 (1989) (185 pp.). [35] R. Coccioni, V. Luciani, Planktonic foraminifers across the Bonarelli Event (OAE2, latest Cenomanian): The Italian record, Palaeogeogr. Palaeoclimatol. Palaeoecol. 224 (2005) 167–185. [36] H. Tsikos, H.C. Jenkyns, B. Walsworth-Bell, M.R. Petrizzo, A. Forster, S. Kolonic, E. Erba, I. Premoli Silva, M. Baas, T. Wagner, J.S. Sinninghe Damsté, Carbon-isotope stratigraphy recorded by the Cenomanian–Turonian Oceanic Anoxic event; correlation and implications based on three key localities, J. Geol. Soc. 161 (2004) 711–719. [37] N. Ohkouchi, Y. Nakajima, H. Okada, N.O. Ogawa, H. Suga, K. Oguri, H. Kitazato, Biogeochemical processes in the saline meromictic Lake Kaiike, Japan: implications from molecular isotopic evidences of photosynthetic pigments, Environ. Microbiol. (2005) 1009–1016.

J. Kuroda et al. / Earth and Planetary Science Letters 256 (2007) 211–223 [38] M. Tanimizu, T. Ishikawa, Development of rapid and precise Pb isotope analytical techniques using MC–ICP–MS and new results for GSJ rock reference samples, Geochem. J. 40 (2006) 121–133. [39] A. Makishima, E. Nakamura, Suppression of matrix effects in ICP–MS by high-power operation of ICP: application to precise determination of Rb, Sr, Y, Cs, Ba, REE, Pb, Th and U at ng g-1 levels in milligram silicate samples, Geostand. Newsl. 21 (1997) 307–319. [40] B.B. Sageman, S.R. Meyers, M.A. Arthur, Orbital time scale and new C-isotope record for Cenomanian–Turonian boundary stratotype, Geology 34 (2006) 125–128. [41] G. Keller, A. Pardo, Age and paleoenvironment of the Cenomanian–Turonian global stratotype section and point at Pueblo, Colorado, Mar. Micropaleontol. 51 (2004) 95–128. [42] W. Kuhnt, A. Nederbragt, L. Leine, Cyclicity of Cenomanian– Turonian organic-carbon-rich sediments in the Tarfaya Atlantic Coastal Basin (Morocco), Cretac. Res. 18 (1997) 587–601. [43] L.R. Kump, M.A. Arthur, Interpreting carbon-isotope excursions: carbonates and organic matter, Chem. Geol. 161 (1999) 181–198. [44] G. Keller, Q. Han, T. Adatte, S.J. Burns, Paleoenvironment of the Cenomanian–Turonian transition at Eastbourne, England, Cretac. Res. 22 (2001) 391–422. [45] K. Kaiho, Planktonic and benthic foraminiferal extinction events during the last 100 m.y. Palaeogeogr. Palaeoclimatol. Palaeoecol. 111 (1994) 45–71. [46] P. Deines, The carbon isotopic geochemistry of mantle xenoliths, Earth-Sci. Rev. 58 (2002) 247–278. [47] H.J. Hansen, Stable isotopes of carbon from basaltic rocks and their possible relation to atmospheric isotope excursions, Lithos 92 (2006) 105–116. [48] M.J. Whiticar, Carbon and hydrogen isotope systematics of bacterial formation and oxidation of methane, Chem. Geol. 161 (1999) 291–314. [49] G. Etiope, R.W. Klusman, Geologic emissions of methane to the atmosphere, Chemosphere 49 (2002) 777–789. [50] C.R. Scotese, J. Golonka, Paleogeographic Atlas, University of Texas, Arlington, 1992.

223

[51] M.A. Arthur, B.B. Sageman, Marine black shales: depositional mechanisms and environments of ancient deposits, Annu. Rev. Earth Planet. Sci. 22 (1994) 499–551. [52] M. Sinha, K. Louden, B. Parsons, The crustal structure of the Madagascae Ridge, Geophys. J. R. Astron. Soc. 66 (1981) 351–377. [53] T. Thordarson, S. Self, Atmospheric and environmental effects of the 1783–1784 Laki eruption: A review and reassessment, J. Geophys. Res. 108 (2003), doi:10.1029/2001JD002042. [54] H.C. Jenkyns, A. Forster, S. Schouten, J.S. Sinninghe Damsté, High temperatures in the Late Cretaceous Arctic Ocean, Nature 432 (2004) 888–892. [55] A. Forster, S. Schouten, K. Moriya, P.A. Wilson, J.S. Sinninghe Damsté, Tropical warming and intermittent cooling during the Cenomanian/Turonian Oceanic Anoxic Event (OAE2): sea surface temperature records from the equatorial Atlantic, Paleoceanography, in press. [56] T. Thordarson, S. Self, Sulfur, chlorine and fluorine degassing and atmospheric loading by the Roza eruption, Columbia River Basalt Group, Washington, USA, J. Volcanol. Geotherm. Res. 74 (1996) 49–73. [57] H. Keppler, M. Wiedenbeck, S.S. Shcheka, Carbon solubility in olivine and the mode of carbon storage in the Earth's mantle, Nature 424 (2003) 414–416. [58] J.C. Zachos, U. Röhl, S.A. Schellenberg, A. Sluijs, D.A. Hodell, D.C. Kelly, E. Thomas, M. Nicolo, I. Raffi, L.J. Lourens, H. McCarren, D. Kroon, Rapid acidification of the ocean during the Paleocene– Eocene Thermal Maximum, Science 308 (2005) 1611–1615. [59] S. Turgeon, H.-J. Brumsack, Anoxic vs dysoxic events reflected in sediment geochemistry during the Cenomanian–Turonian Boundary Event (Cretaceous) in the Umbria–Marche Basin of central Italy, Chem. Geol. 234 (2006) 321–339. [60] J.A. Higgins, D.P. Schrag, Beyond methane: Towards a theory for the Paleocene–Eocene Thermal Maximum, Earth Planet. Sci. Lett. 245 (2006) 523–537. [61] T.J. Bralower, P.D. Fullagar, C.K. Paull, G.S. Dwyer, R.M. Leckie, Mid-Cretaceous strontium-isotope stratigraphy of deep-sea sections, Geol. Soc. Amer. Bull. 109 (1997) 1421–1442.

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