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Early Aptian paleoenvironmental evolution of the Bab Basin at the southern Neo-Tethys margin: Response to global carbon-cycle perturbations across Ocean Anoxic Event 1a

Kazuyuki Yamamoto

Institute of Geology and Paleontology, Graduate School of Science, Tohoku University, Aobayama, Sendai 980-0856, Japan

Present address: Abu Dhabi Project Division, INPEX Corporation, Akasaka Biz Tower 5-3-1, Akasaka, Minato-ku, Tokyo 107-6332, Japan

Masatoshi Ishibashi

Abu Dhabi Oil Company, Ltd (Japan), P.O. Box 630, Abu Dhabi, United Arab Emirates

Hideko Takayanagi

Department of Earth and Planetary Sciences, Graduate School of Environmental Studies, Nagoya University, Furo-cho, Chikusa-ku, Nagoya 464-8601, Japan

Yoshihiro Asahara


Tokiyuki Sato

Institute of Applied Earth Sciences, Faculty of Engineering and Resource Science, Akita University, Tegata-Gakuencho 1-1, Akita 010-0852, Japan

Hiroshi Nishi

The Center for Academic Resources and Archives, Tohoku University Museum, Tohoku University, Aobayama, Sendai 980-0856, Japan

Yasufumi Iryu


Present address: Institute of Geology and Paleontology, Graduate School of Science, Tohoku University, Aobayama, Sendai 980-0856, Japan ([email protected])

This article has been accepted for publication and undergone full peer review but has not been through the copyediting, typesetting, pagination and proofreading process, which may lead to differences between this version and the Version of Record. Please cite this article as doi: 10.1002/ggge.20083

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Abstract

Lower Aptian carbonates in the Bab Basin at the southern Neo-Tethys margin record

significant environmental changes across Oceanic Anoxic Event 1a (OAE1a). A long-lasting

negative shift of carbon-isotope ratios (δ13C) associated with a distinct decrease in oxygen-

isotope ratios (δ18O) in orbitolinid-rich carbonates characterizes the onset of OAE1a (Livello

Selli), supporting a hypothesis that a long-lasting volcanic CO2 emission is a main cause of

OAE1a, inducing global warming. A bloom of microencrusters (Lithocodium and Bacinella)

across the proto-Bab Basin occurred synchronously at the beginning of the subsequent

positive δ13C excursion, responding to the global carbon-cycle perturbations. The carbonates,

formed during the OAE1a, show higher strontium-isotope ratios (87Sr/86Sr) compared with

those of global seawater; this was likely caused by a local influx of isotopically heavier

strontium, along with nutrients, into the proto-Bab Basin. These biotic proliferations were

triggered by an increased nutrient supply induced by intensified continental weathering due to

the global warming suggested by the increase in δ18O values. Spatial variations in the δ13C

values among sites in the Bab Basin and its surrounding platform are related to local

environmental factors, such as the degree of mixing of basin water with ocean water and local

removal of 12C by metabolic activity at the platform-top. The δ13C profile of the studied core

indicates global removal of organic carbon of OAE1a began during the early stage of the

second-order transgression and lasted until the early stage of the highstand after the OAE1a.

The Livello Selli corresponds to the early stage of this transgression.

1. Introduction

The Early Aptian represents a period of significant environmental changes under extreme

greenhouse conditions [e.g. Skelton et al., 2003]. This period is characterized by an episodic,

widespread accumulation of organic carbon-rich deposits in an anoxic marine setting. This

depositional event, known as “Oceanic Anoxic Event 1a” (OAE1a) [Schlanger and Jenkyns,

1976; Arthur et al., 1990], was associated with environmental changes, such as major

perturbations in global carbon cycling [e.g. Menegatti et al., 1998], global sea-level rise [Haq

et al., 1988], increases in continental weathering and runoff [Erba, 1994; Föllmi et al., 1994;

Misuimi et al., 2009; Tejada et al., 2009; Blättler et al., 2011; Najarro et al., 2011a; Bottini et

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al., 2012], significant changes in flora and fauna [Larson and Erba, 1999; Erba, 1994, 2004;

de Gea et al., 2003; Erba and Tremolada, 2004; Weissert and Erba, 2004; Erba et al., 2010],

and drowning of carbonate platforms [Föllmi et al., 1994; Graziano, 1999; Wissler et al.,

2003; Burla et al., 2008]. In both marine carbonates and organic matter, OAE1a is marked by

a positive excursion of carbon-isotope ratios (δ13C) that is preceded by a pronounced negative

excursion of δ13C values. This trend has been identified at various localities around the globe

[Jenkyns, 1995; Menegatti et al., 1998; Bralower et al., 1999; Jenkyns and Wilson, 1999;

Ando et al., 2002; de Gea et al., 2003; Danelian et al., 2004; Dumitrescu and Brassell, 2006;

Föllmi et al., 2006; van Breugel et al., 2007; Millán et al., 2009]. Although the driving

mechanisms for OAE1a has been much debated, it is generally accepted that dissociation of

methane hydrate [e.g. Jahren et al., 2001; Beerling et al., 2002] and/or volcanic CO2

emission [e.g. Méhay et al., 2009; Tejada et al., 2009; Kuroda et al., 2011] were causes of the

negative excursion of δ13C values at the onset of OAE1a. A numerical simulation of

atmospheric and oceanic biogeochemical cycles specified the most critical factors promoting

and sustaining the oceanic anoxia during OAE1a [Misumi et al., 2009]. Intensified

weathering under the elevated atmospheric CO2 condition increases phosphate concentration

and export production in the ocean, resulting in an increase in the burial flux of organic

carbon and in turn making deep water anoxic. This is supported by some studies that

separated volcanogenic phases and weathering spikes in the OAE1a interval [Ando et al.,

2008; Tejada et al., 2009; Mehay et al., 2009; Erba et al., 2010; Blättler et al., 2011; Bottini

et al., 2012].

Although many studies focused on pelagic deposits, their coeval shallow-water platform

carbonates were also investigated to assess effects of OAE1a in a shallow-marine

environment [Vahrenkamp, 1996, 2010; Grötsch et al., 1998; Jenkyns and Wilson, 1999; van

Buchem et al., 2002; Immenhauser et al., 2004, 2005; Burla et al., 2008; Huck et al., 2010;

Rameil et al., 2010]. The southern Neo-Tethys margin, especially in the southern Arabian

Gulf region, is an ideal area for such investigations, because well-developed shallow-water

carbonate platforms existed there during the Early Aptian (Figure 1A and B) [Grötsch et al.,

1998; Pittet et al., 2002; van Buchem et al., 2002, 2010; Hillgärtner et al., 2003;

Immenhauser et al., 2004, 2005]. The Lower Aptian deposits in this region include the

uppermost part of the Kharaib Formation and the Shu’aiba Formation (Figure 1B). During

deposition of the latter, an intra-shelf basin (Bab Basin) was formed that was surrounded by

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giant carbonate platforms (Figure 1A and B) [Murris, 1980; Sharland et al., 2001; Yose et al.,

2006, 2010; van Buchem et al., 2002, 2010]. Recently, the results of comprehensive studies

on the litho- and chronostratigraphy of the Shu’aiba Formation were published; these were

based on data from subsurface and outcrop sections in the southern Arabian Gulf region [e.g.

Al-Ghamdi and Read, 2010; Droste, 2010; Pierson et al., 2010; Schroeder et al., 2010;

Strohmenger et al., 2010; Vahrenkamp, 2010; van Buchem et al., 2010; Yose et al., 2010].

Chemostratigraphic data, such as carbon-isotope stratigraphy, were also reported by several

investigators [Vahrenkamp, 1996, 2010; van Buchem et al., 2002; Al-Ghamdi and Read,

2010; Droste, 2010; Strohmenger et al., 2010]. These data are derived mainly from the

proximal shallow-water platform and associated slope carbonates around the Bab Basin

(Figure 1A and B).

This article presents an integrated data set consisting of Lower Aptian high resolution,

bulk carboante carbon (δ13C)-, oxygen (δ18O)-, and strontium (87Sr/86Sr)-isotope

stratigraphies, together with results of paleontological analyses on calcareous nannofossils

and ammonites from a core drilled at a distal central site in the Bab Basin (Figures 1 and 2).

Although shallow-water platform deposits, with their associated high-temporal resolution, are

the best archive to use for delineating the response of shallow-marine organisms and

ecosystems to environmental changes, they are susceptible to alteration as a result of meteoric

diagenesis [Allan and Matthews, 1982], and the geochemical signatures of platform-top

deposits are not always representative of open-marine conditions [Patterson and Walters,

1994; Immenhauser et al., 2003, 2008; Swart and Eberli, 2005; Vahrenkamp, 2010]. In

contrast, such signals are more commonly and better preserved in pelagic deposits, although

their temporal resolution is relatively low because of their slower sedimentation rate. The

studied core, consisting of lower shallow-water and upper basin carbonates, was drilled at a

distal central site in the Bab Basin situated in an intermediate position between shallow-

marine platform and pelagic settings. Therefore, our data provide excellent insight into global

carbon-cycle perturbations and associated shallow-marine environmental changes, as well as

evolution of an intra-shelf basin at the southern Neo-Tethys margin during the Early Aptian.

2. Regional Geological Framework

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During the Early Cretaceous, the Arabian Plate was located in a tropical low-latitude

region and constituted part of the southern Neo-Tethys margin in a passive margin setting

[Murris, 1980; Hughes, 2000]. The eastern Arabian Plate was extensively flooded as a result

of a global rise in sea-level, during which shallow-water carbonates accumulated on the

stable craton (Figure 1A). The Kharaib and Shu’aiba formations represent part of these

autochthonous deposits (Figure 1B) [e.g. Murris, 1980; Sharland et al., 2001]. The Kharaib

Formation, ranging in age from Early Barremian to early Early Aptian, is characterized by

ramp-type carbonate deposits that show predominantly aggradational stacking patterns

[Murris, 1980; van Buchem et al., 2002]. The uppermost portion of the formation consists of

orbitolinid-dominated, argillaceous limestones (Hawar Member), which were deposited

during the early Early Aptian (Figures 1B and 3A) [Murris, 1980].

In contrast to the relatively monotonous lithology and laterally continuous sedimentary

features of the Kharaib Formation, the Shu’aiba Formation exhibits a sedimentary

architecture characterized by an intra-shelf basin surrounded by carbonate platforms (Figure

1A and B). After deposition of the Kharaib Formation, shallow-water platform carbonates

accumulated mainly in near-land areas in association with the Early Aptian rise in sea-level.

In contrast, the accumulation rate of carbonates decreased significantly around the center of

the previous shallow-ramp area. This disparity resulted in a topographic depression (Bab

Basin) surrounded by shallow-water platforms characterized by the aggradation of lower

facies dominated by problematic microencrusters, Lithocodium aggregatum and Bacinella

irregularis (hereafter referred to as “Lithocodium–Bacinella”) (Figure 3B), and upper facies

dominated by rudists [van Buchem et al., 2002, 2010; Yose et al., 2006, 2010; Droste, 2010;

Strohmenger et al., 2010]. The Bab Basin was eventually filled with a mixture of prograding

carbonate clinoforms and argillaceous basinal deposits during the Late Aptian sea-level fall

[Murris, 1980; van Buchem et al., 2002, 2010; Pierson et al., 2010]. The Shu’aiba Formation

is unconformably overlain by shallow-marine shales and argillaceous limestones of the

Albian Nahr Umr Formation (Figure 1B) [e.g. Murris, 1980; Sharland et al., 2001].

3. Materials and Analytical Techniques

The studied core was obtained from an oil field offshore Abu Dhabi, United Arab

Emirates (UAE) (Figure 1A). It is 53.03 m in total length, and the depth interval ranges from

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2667.91 to 2614.88 mbsf (meters below the seafloor) (Figure 2). The core recovery is 100%.

Because the field has a gentle anticlinal structure dipping less than one degree, the vertical

thickness of the formation is approximate to its actual thickness.

A total of 301 bulk carbonate samples were collected from the core for geochemical

analyses; efforts were made to avoid recrystallized large bioclasts, large cement crystals,

stylolites, and pressure solution seams.

To identify and exclude diagenetically altered samples, mineral abundance, and trace

element (strontium [Sr], manganese [Mn], and iron [Fe]) concentration were determined for

197 and 126 samples, respectively. The mineral abundance was determined following Suzuki

et al. [2006] with a Phillips X’pert-MPD PW3050 system at the Institute of Geology and

Paleontology, Tohoku University, Japan (IGPS) and a Rigaku MultiFlex system at the

Nagoya University Museum, Japan. Based on results of the XRD analysis, 14 samples were

excluded from the subsequent isotope analyses because of their relatively high content of

dolomite (<86wt%) and/or pyrite (<6.0wt%). The trace element concentrations were

determined with a Varian Vista Pro Radial Inductively Coupled Plasma-Atomic Emission

Spectrometer (ICP-AES) by the following method. 0.200 g of the powdered sample was

added to lithium metaborate/lithium tetraborate flux (0.90 g), and fused in a furnace at

1000°C. The obtained melt sample was dissolved in 100 ml of 4% nitric acid/2%

hydrochloric acid. The resulting solution was analyzed by ICP-AES. Oxide concentration

was calculated from the determined elemental concentration (%). The detection limit of each

element was 0.01%, and the analytical error was less than 5% of the measured concentration

for all measured elements. We express data as Sr, Mn and Fe (not as oxides).

To provide chronological constraint on the studied carbonate sequence, strontium-isotope

ratios (87Sr/86Sr) and calcareous nannofossil assemblages were analyzed on 33 and 60

samples, respectively. 87Sr/86Sr was measured with a VG Sector 54-30 thermal ionization

mass spectrometer at the Department of Earth and Planetary Sciences, Nagoya University

basically following Asahara et al. [1999, 2006] and Suzuki et al. [2012]. The external

precision determined by replicate analysis of the NIST (National Institute of Standards and

Technology) SRM (Standard Reference Materials) 987 was less than ±0.000026 (2σ).

Numerical ages were obtained by comparison with the 87Sr/86Sr evolution of global seawater

reported as the Look-up Table Version 4:08/04 [Howarth and McArthur, 1997; McArthur et

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al., 2001]. Standard smear slide methods were used to analyze calcareous nannofossil

assemblages [Sato et al., 2004]. Because calcareous nannofossils are absent or rare in the

studied samples, we did not collect (semi-)quantitative data of their stratigraphic distribution

but recorded presence/absence of the species listed in Table 1. We followed the Tethyan

nannofossil biostratigraphy by Bown et al. [1998], which was originally established by Roth

[1978] and later revised by Bralower et al. [1995].

To establish chemostratigraphy, the carbon (δ13C)- and oxygen (δ18O)-isotope ratios of

158 samples was measured following Yamamoto et al. [2010] with a Finnigan deltaS mass

spectrometer at IGPS or a Finnigan MAT 252 mass spectrometer at Technology Research

Center, Japan Oil, Gas and Metals National Corporation (JOGMEC/TRC), each coupled with

a ThermoQuest Kiel-III automated carbonate device. The external precision determined by

replicate analysis of the laboratory standard was less than ±0.04‰ for δ13C values and

±0.07‰ for δ18O values (1σ) at IGPS and ±0.03‰ for δ13C values and ±0.05‰ for δ18O

values at JOGMEC/TRC.

4. Results

4.1. Lithostratigraphy

The carbonate sequence in the studied core consists of 12 lithostratigraphic units that are

numbered sequentially from the base (unit 1) to the top (unit 12) (Figures 2 and 3). Detailed

lithologic descriptions of these units are given in Table 2. The studied core represents a

continuous depositional record as no erosional surface indicating a hiatus is observed. XRD

analysis showed that the carbonates are composed almost exclusively of calcite, with the

exception of a dolomitized interlayer in unit 12. Minor amounts of other minerals, such as

quartz, pyrite, and some clay minerals, were noted, especially in units 4, 5, and 12.

Units 1 through 5 correlate with the uppermost part of the Kharaib Formation. Of these,

units 4 and 5 are equivalent to the Hawar Member, which is characterized by orbitolinid-

dominated, argillaceous limestones in the southern Arabian Gulf region [e.g. van Buchem et

al., 2002, 2010]. Units 6 through 12 correspond to the lower part of the Shu’aiba Formation.

Unit 10 is characterized by the abundant planktonic foraminifers, which suggests a deepest

depositional environment of all the units in the studied core. The base of unit 10 corresponds

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to the maximum flooding surface of the second-order depositional sequence, “MFS K80”

defined by Sharland et al., [2001] (Figures 1B and 2).

4.2. Trace Element Concentration

The concentration of Sr in the core samples primarily ranges from 0.01 to 0.08wt%, with

a few outliers reaching 0.18wt% within unit 11. The concentration of Sr is low and relatively

constant (0.03–0.05wt%) in units 1 through 7. Above unit 7, the concentration gradually

increases upward, reaching 0.08wt% in the lower part of unit 10. The concentration typically

falls between 0.03 and 0.07wt% in units 10 through 12 and is associated with large

fluctuations, especially in the lower part of unit 11. In contrast, the concentration of Mn is

generally low throughout the core and ranges from <0.01wt% (below the detection limit) to

0.03wt%. No regular trend with depth is recognized. The concentration of Fe varies from

0.01 to 5.35wt%. Units 6 and 7 have relatively low and constant values of 0.10wt% or less,

whereas the concentration is relatively higher in the other units, showing irregular

fluctuations.

4.3. Carbon- and Oxygen-Isotope Stratigraphies

The δ13C values, which range from 1.6 to 4.2‰, show negative and positive excursions

upsection (Figure 2; Table 3). The δ13C values fall within a narrow range from 3.4 to 3.8‰ in

units 1 through 3, with a minor negative spike (<0.3‰) in unit 2. An abrupt negative shift in

the δ13C values occurs within unit 4: the values decrease from 3.3‰ at the base to 2.0‰ near

the top. The δ13C values are relatively constant (2.1–2.3‰) in the interval from the

uppermost horizon of unit 4 to the upper horizon of unit 5; this is followed by a minor

decrease, reaching 1.6‰ at the uppermost horizon of unit 5. This is the minimum δ13C value

found throughout the studied core. Minor fluctuations, with an amplitude <0.6‰, occur at the

base of unit 6; these are followed by a prolonged positive excursion that continues through

the upper horizon of unit 10 (4.2‰). This excursion intercalates with minor negative spikes

(<0.5‰) in the middle part of unit 7. Increases in the δ13C values are rapid and gradual below

and above the negative spikes in unit 7, respectively. The δ13C values vary between 3.3 and

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4.2‰ in unit 10; the maximum δ13C value of 4.2‰ is recorded from the uppermost part of

this unit. The δ13C values gradually decrease upward from 3.9 to 2.9‰ in units 11 and 12.

The δ18O values range from –7.1 to –3.7‰ in the studied section and exhibit many

fluctuations, with amplitudes <~2.0‰ (Figure 2; Table 3). The values vary in a range from –

6.4 to –5.2‰ in units 1 through 3. The δ18O profile displays an increase in the interval from

the base of unit 4 to the lower horizon of unit 5, followed by a decrease that terminate by a

negative spike at the upper horizon of unit 7, which reaches –7.1‰, the minimum δ18O value

found throughout the studied core. This decrease intercalates a positive spike in the upper

part of unit 5, which reaches –3.7‰, the maximum δ18O value found throughout the studied

core. Above the negative-spike horizon, the δ18O values increase up to –4.2‰ at the upper

horizon of unit 8. In the interval from the base of unit 9 to the middle of unit 11, the δ18O

values mostly vary between –7.0 and –5.5‰. The δ18O profile shows an increasing trend in

the upper part of unit 11 followed by a gradual decrease in unit 12.

4.4. Strontium-Isotope Ratio

87Sr/86Sr ranges from 0.70752 to 0.70736 in the studied core (Figure 2; Table 4). The

strontium-isotope profile displays little variation in units 1 through 4, with an outlier of

0.70752 in unit 3. Then, the ratios increase and reach the maximum value of 0.70751 at the

lower horizon in unit 7, which is followed by a gradual decrease through unit 9 and minor

fluctuations around the value of 0.70738 in units 10 through 12. The 87Sr/86Sr evolution of

global seawater displays a decreasing trend from 0.707490 in the Upper Barremian to

0.707217 immediately below the Aptian/Albian boundary [McArthur et al., 2001].

The 87Sr/86Sr data from the studied core show a similar trend, with the exception of units 5

through 7 (Figure 2).

Numerical ages obtained by comparison with the global 87Sr/86Sr record [Howarth and

McArthur, 1997; McArthur et al., 2001] fall within a range from 122 to 128 Ma (Table 4).

However, the ages do not decrease monotonically upsection, and age reversal is recognized in

units 5 through 7. Units 1 through 4 and 8 through 12 correlate with the Upper Barremian and

Lower to Upper Aptian, respectively [Gradstein et al., 2004]. The Early/Late Aptian ages

obtained from units 8 through 12 are supported by our biostratigraphic data.

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4.5. Biostratigraphy

Calcareous nannofossils are absent or rare throughout the studied core; however, they are

relatively more common in the upper interval (units 8 through 12). Of all analyzed samples,

34 samples yielded calcareous nannofossils, which were in a moderate to good state of

preservation (Figure 2; Table 1).

Calcareous nannofossil assemblages are characterized by low species diversity and

primarily represented by Watznaueria barnesae. This species occurs from units 4, 5, and 7

through 12. Rhagodiscus spp. (R. asper and R. angustus) and Nannoconus spp. are found

from units 8 through 12.

The first occurrence of R. angustus, along with that of Eprolithus floralis, is considered to

define the base of the nannofossil zone NC7A [Roth, 1978; Bralower et al., 1995; Bown et al.,

1998; Bellanca et al., 2002] or to be located in the middle of the Aptian nannofossil zone

NC6B [Masse, 2002]. In contrast, Prediscosphaera columnata, which appears at the

Aptian/Albian boundary, is absent from the studied core. These indicate that the studied

interval correlates mostly to the Aptian and is older than the Albian.

Many molds of ammonite shell fragments occur in unit 9. A large mold of Pseudosaynella

raresulcata is found at 2639.16 mbsf in this unit (Figure 2). P. raresulcata occurs from the

upper part of the Deshayesites weissi Zone to the Deshayesites deshayesi Zone in the upper

Lower Aptian [Grauges et al. 2010; Moreno-Bedmar et al., 2010]. Because the middle part of

the nannofossil zone NC6B and the lowest part of the nannofossil zone NC7A overlap with

the upper part of the D. deshayesi Zone [Bown et al., 1998], the horizon containing the first

occurrence of R. angustus at 2641.90 mbsf in unit 8 is considered to be close to the base of

NC7A in the upper Lower Aptian.

In the studied core, Micrantholithus hoschulzii occurs at 2629.30 and 2622.28 mbsf in

unit 11. Generally, the last occurrence of Micrantholithus spp. is correlated to the

NC7A/NC7B boundary [Roth, 1978; Bralower et al., 1995] that is close to the Lower/Upper

Aptian boundary [Bown et al., 1998]. In spite of the relatively more common occurrence of

calcareous nannofossils in the upper interval (units 8 through 12), the acme of Nannoconus

truitti or occurrence of R. achylostaurion, both of which indicate middle Late Aptian age, are

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not recognized. Although the occurrence of M. hoschulzii is limited, it is expected that the

Lower/Upper Aptian boundary could be close to or lower than 2622.28 mbsf in the upper part

of unit 11. This is not in conflict with the carbon- and strontium-isotope stratigraphies

established in this study.

5. Discussion

5.1. Diagenetic Evaluation

Dissolution cavities/vugs and calcite veins were very rare in the studied core. Other

diagenetic features, such as karstification and hydrothermal alteration, were not observed.

Based on their investigation on the relation between the trace element concentration in

carbonates and diagentic alteration, Denison et al. [1994] found that samples with Sr/Mn > 2

or Mn < 300ppm retained the initial 87Sr/86Sr. Furthermore, Jacobsen and Kaufman [1999]

showed that Mn/Sr was available to separate diagenetically altered (Mn/Sr > 2) and unaltered

(Mn/Sr < 2) carbonate samples for 87Sr/86Sr analysis. These criteria were used in many

studies [e.g. Suzuki et al., 2012]. Two samples from unit 5 have Mn concentration slightly

greater than 300ppm. Sr/Mn is less than 2 in 15 samples (7 samples from unit 5, 1 from unit 8,

2 from unit 11, and 5 from unit 12). All the samples, however, satisfy the criterion of Mn/Sr <

2. The δ13C profile of the studied core is characterized by smooth, systematic decreases and

increases lacking distinctly anomalous values, and it is correlated well with coeval δ13C

profiles in other areas (Figures 4 and 5). Consequently, we consider that the studied samples

retain initial 87Sr/86Sr and δ13C values of carbonates when they were deposited.

The δ18O values from the studied samples are (–7.1 to –3.7‰) lower than those at other

sites (Figure 6) and those (–2.6 to 4.6‰) of calcite precipitated in oxygen isotope equilibrium

with Cretaceous seawater that were estimated assuming a δ18O value for seawater of –0.5‰

(vs. SMOW; Standard Mean Ocean Water) and temperatures of 20 to 35°C [Steuber, 1999].

However, the δ18O profile of the studied core shows a common trend with coeval δ18O

profiles in other areas (Figure 6). These suggest that, although the initial oxygen-isotope

composition has been modified because of overprints by meteoric diagenesis and/or increased

temperatures associated with burial, the overprints likely occurred to a similar extent

throughout the studied core.

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Low Fe concentration was used to identify carbonates that are likely to retain the initial

isotope composition in many studies (e.g. <3,000ppm [Denison et al., 1994]). The studied

samples contain <~9.0wt% of non-carbonate material including pyrite except for one sample

(~17wt%). Therefore, Fe concentration cannot be used to identify diagenetically

altered/unaltered carbonate samples from the studied core.

5.2. Correlation of the δ13C and δ18O Profiles with Other Reference Curves

The δ13C profile for the Upper Barremian to the Upper Aptian can be divided into eight

segments (C1 to C8, in ascending order); this scheme was originally proposed by Menegatti

et al. [1998] for the sequences at Cismon, Italy and Roter Sattel, Switzerland (Figures 2, 4,

and 5). This chemostratigraphic framework has been widely accepted by many investigators

[Erba et al., 1999; Bellanca et al., 2002; Price 2003; Heimhofer et al., 2004; Dumitrescu and

Brassell, 2006; Ando et al., 2008; Heldt et al., 2008; Li et al., 2008; Millán et al., 2009; Erba

et al., 2010; Huck et al., 2011; Kuhnt et al., 2011; Najarro et al., 2011b].

The δ13C profile of the studied core, with the C1–C8 segmentation constrained by

strontium-isotope stratigraphy and biostratigraphy, compares well with other published δ13C

profiles from the Tethyan and Pacific regions (Figure 5) [Menegatti et al., 1998; Jenkyns and

Wilson, 1999; Föllmi et al., 2006; Vahrenkamp, 2010; Kuhnt et al., 2011; Hu et al., 2012].

Although segment C3, as proposed by Menegatti et al. [1998], is characterized by a short-

lived negative shift in the sections that they investigated, we correlate the interval containing

the abrupt decrease and the subsequent relatively low δ13C values in the studied core with

segment C3 (Figure 5). The following segments C4–C6 is a recovery stage from the

minimum δ13C value (Figure 2). Segments C3–C6 are time-equivalent to OAE 1a (Livello

Selli) [e.g. Menegatti et al., 1998; Erba et al., 1999]. The δ13C profile of the studied core

records global carbon-cycle perturbations and is not affected by facies control on the initial

δ13C excursions/shifts. Other than the excellent correlation of the δ13C profile with those from

over the globe, three reasons can be presented.

1) The segment boundaries do not necessarily correspond to lithologic boundaries (unit

boundaries).

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2) It is known microbially-induced micrite is usually enriched in 13C relative to marine

carbonates [e.g. Wu and Chafets, 2000]. Because Lithocodium and Bacinella are commonly

associated with microbilalites [Hillgärtner et al., 2003], microbially-induced micrite may be

included in unit 6. However, a sedimentary facies-related increase in δ13C values is not

recognized around unit 6.

3) Local to basin-wide heterogeneity of (= spatial variations in) δ13C values of dissolved

inorganic carbon in seawater must be taken into consideration when interpreting the

geochemical record of ancient epeiric seas [Immenhauser et al., 2003, 2008]. Such effect is

evident when we compare the δ13C profiles among proximal to distal sites in the Bab Basin

(Figure 4). However, the effect has limited impacts on the profiles as discussed later, and all

of them are correlated well with those from the Tethyan and Pacific regions (Figure 4).

Although it is still controversial, the Barremian/Aptian boundary has generally been

placed at the base of the Hawar Member (units 4 and 5 in this study) in the southern Arabian

Gulf region [e.g. van Buchem et al., 2002, 2010; Al-Husseini and Matthews, 2010; Schroeder

et al., 2010]. However, the Barremian/Aptian boundary is well-defined by bio-chemo-

magnetostratigraphy in the northern Neo-Tethys sections, and correlated with a minor

negative δ13C spike in segment C2 [Erba, 1999; Gradstein et al., 2004; Huck et al., 2011].

The correlative negative spike is identified in the lower part of unit 2 of the studied core

(2664.89 mbgs; Figure 2). We obtained strontium-isotope ages, indicating the Late

Barremian from unit 2 (125.6–126.6 Ma) and unit 4 (125.4 Ma). Taking into account the

uncertainty of the numerical age of the Barremian/Aptian boundary (125.0 ± 1.0 Ma;

Gradstein et al., [2004]), we correlate the boundary to the minor negative δ13C spike in

segment C2 in unit 2 (Figure 2).

The decrease in the δ18O values in unit 5 through the upper horizon of unit 7 and the

subsequent increase to unit 8 are correlated with segments O1 and O2 defined by Menegatti

et al. [1998], respectively (Figure 6). In the studied core, segment O1 is correlated with

segment C3 to the basal part of segment C7; segment O2 corresponds to the lower part of

segment C7. The negative δ18O spike identified immediately above the Livello Selli (= at

lowermost horizon of segment C7) by Menegatti et al. [1998] is comparable to that at the

base of segment C7 (upper part of unit 7) in the studied core. Segments O1 and O2 were

identified in other published δ18O profiles from the Tethyan and Pacific regions (Figure 6)

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[Menegatti et al., 1998; Jenkyns and Wilson, 1999; Bellanca et al., 2002; Ando et al., 2008;

Erba et al., 2010; Vahrenkamp, 2010; Kuhnt et al., 2011; Hu et al., 2012]. However, timing of

segments O1 and O2 is not completely comparable among the published δ18O profiles (Figure 6). In

this paper, they are defined as those representing decreasing and increasing trends in δ18O values

commonly during the period of segments C2–C7 intercalating OAE1a.

5.3. Long-Lasting Negative δ13C shift at the Onset of OAE1a

Previous researches presented different estimates of the duration of segment C3: 27–44

ky [Li et al., 2008], 22–47 ky [Malinverno et al., 2010], >0.1 Myr [Kuhnt et al., 2011], and

0.32 Myr [Hu et al., 2012]. Therefore, we define that the terms “short-lived” and “long-

lasting” represent <50 ky and > 0.1 Myr, respectively, in this article.

The cause of the negative δ13C shift (segment C3 of Menegatti et al. [1998]) that

generally defines the onset of OAE1a has long been discussed and interpreted as resulting

from the dissociation of methane hydrates [e.g. Jahren et al., 2001; Beerling et al., 2002]

and/or emission of volcanic CO2 [e.g. Méhay et al., 2009; Tejada et al., 2009; Kuroda et al.,

2011]. The former idea is supported by the nature of a short-lived spiky shift with significant

negative δ13C values that can be explained by a rapid catastrophic release of isotopically very

light carbon derived from methane hydrate dissociation. However, in the studied core,

segment C3 is not short-lived; rather, it is a long-lasting negative shift in δ13C values (Figure

2). Li et al. [2008] calculated the sedimentation rate of segments C3–C6 using an

astrochronological method at three sites (Cismon in Italy, Santa Rosa Canyon in Mexico, and

DSDP Site 398 in the North Atlantic Ocean) and concluded that segments C3–C6 correlated

with 1.0–1.3 Myr. They calculate the duration of segment C3 as 27–44 ky. Subsequently,

Malinverno et al. [2010] estimated the timing and duration of OAE1a by an applying orbital

tuning method to a Lower Cretaceous (Barremian–Aptian) sequence at Cismon. They

concluded that OAE1a lasted for 1.11 Myr and a sudden negative δ13C shift at the base of the

Selli level was short lived (22–47 ka). In contrast, some studies showed that segment C3 was

long lived and lasted for >0.1 Myr. Kuhnt et al. [2011] identified a long-lasting negative shift

in δ13C values for segment C3 in the δ13C profile obtained from the La Bédoule section

(southeastern France) (Figure 5), where the sedimentation rate was relatively high. They

estimated the duration of segment C3 as >0.1 Myr, based on deposit thickness of the segment

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and the time interval of segments C3–C6 calculated by Li et al. [2008]. They pointed out that

the short-lived character of the negative excursion at Cismon might indicate condensed

sedimentation or a hiatus at the base of the Livello Selli. The carbon- and strontium-isotope

stratigraphies of Barremian–Aptian shoal-water carbonates in eastern France indicated that

segment C3 corresponded to a duration of 0.285–0.333 Myr [Huck et al., 2011). A similar

estimate (0.320 Myr) was obtained from litho-, chemo-, and cyclostratigraphic analyses of

Yenicesihlar section (Turkey).

As noted above, there are two contrasting views on the duration of the negative δ13C shift

at the onset of OAE1a (segment C3): short lived or long lasting. In this study, the thicknesses

of segments C3, C4, C5, and C6 are 4.63, 9.37, 1.29, and 0.24 m, respectively. If the total

duration of 1.0–1.3 Myr for segments C3–C6 is simply divided, assuming a constant

sedimentation rate, then the duration of segment C3 is calculated as ~0.30–0.39 Myr. This

supports the latter view that the negative δ13C shift was long lasting. Segment C3 in the

studied core corresponds to the entire Hawar Member (units 4 and 5) of the Kharaib

Formation. In other locations in the southern Arabian Gulf region as well, the negative shift

in δ13C values representing segment C3 is not a short-lived spiky negative shift; rather, it is a

long-lasting negative shift that continued during deposition of the Hawar Member (Figure 4)

[van Buchem et al., 2002; Droste, 2010; Strohmenger et al., 2010; Vahrenkamp, 2010].

Although it is difficult to estimate the sedimentation rate of the Hawar Member, δ13C profiles

in this region indicate that the time interval of segment C3 was relatively long. These data

from the southern Arabian Gulf region do not support the hypothesis that a short-lived

catastrophic dissociation of methane hydrate was the main driving factor for the negative

δ13C shift at the onset of OAE1a. Recently, a temporal relationship between a massive

eruption on the Ontong Java Plateau (OJP) and OAE1a was documented based on lead- and

osmium-isotope records [Tejada et al., 2009; Kuroda et al., 2011; Bottini et al., 2012]. Their

causal relationship was also verified with a numerical simulation [Misumi et al., 2009].

Consequently, a massive emission of volcanic CO2, probably primarily from the OJP, and/or

intermittent methane dissociation over a prolonged period of time are the most likely causes

of the long-lasting negative shift in δ13C values [Kuhnt et al., 2011]. In the latter case, global

warming caused by the elevated atmospheric CO2 resulting from the volcanic eruptions may

have triggered the intermittent methane dissociation during that time. The global warming at

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the early stage of OAE1a is evidenced by the decrease in δ18O values (segment O1) during

the period of segment C3 through the base of segment C7 (Figures 2 and 6).

5.4. A Synchronous Bloom of Lithocodium–Bacinella across the Proto-Bab Basin

in the Early Stage of OAE1a

During the Early Aptian, many carbonate platforms in the northern Neo-Tethys margin

were episodically drowned, which caused environmental stresses to shallow-water carbonate

factories [Föllmi et al., 1994; Bosellini et al., 1999; Wissler et al., 2003; Burla et al., 2008;

Huck et al., 2010]. In contrast, Early Aptian carbonate platforms situated along the central to

southern Neo-Tethys margins record continuous carbonate deposition marked by an abrupt

faunal change in the platform biota [Grötsch et al., 1998; Pittet et al., 2002; van Buchem et

al., 2002; Hillgärtner et al., 2003; Immenhauser et al., 2004, 2005; Huck et al., 2010; Rameil

et al., 2010]. Rudist–coral–stromatoporoid communities, which were the most common reef

builders in the Cretaceous carbonate platforms, were stressed and episodically replaced with

Lithocodium–Bacinella [e.g. Dupraz and Strasser, 1999].

In the southern Arabian Gulf region, Lithocodium–Bacinella in the Shu’aiba Formation

formed large buildups during the Early Aptian; in Oman, these outcrops measure several tens

of kilometers across, with a thickness of several tens of meters [Immenhauser et al., 2005;

Rameil et al., 2010]. In the studied core from the distal central site in the Bab Basin, the

lowest stratigraphic horizon of the Lithocodium–Bacinella bloom (base of unit 6) coincides

with the base of segment C4, which is characterized by a positive shift in δ13C values in the

early stage of OAE1a (Figure 2). In the proximal margin of the Bab Basin, the timing of the

onset of the Lithocodium–Bacinella proliferation also seems to coincide with the beginning

of the positive shift in δ13C values corresponding to segment C4 (Figure 4) [Droste, 2010;

Strohmenger et al., 2010; Vahrenkamp, 2010]. These data indicate that the Lithocodium–

Bacinella blooms originated simultaneously at the base of segment C4 in both areas. This is

also supported by stratigraphic data. The recovery from the “nannoconid crisis” [e.g. Erba,

1994; Erba et al., 2010] began and flux of nannoplanktons increased in a Tethyan pelagic

environment during the period of segment C4 [Erba et al., 2010], which coincides with the

Lithocodium–Bacinella bloom on the carbonate platforms in the southern Arabian Gulf

region.

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The Bab Basin did not exist when the Hawar Member, which underlies the Shu’aiba

Formation, accumulated during the Early Aptian. The Hawar Member exhibits a gradual

thinning and a change in its depositional facies into more distal facies toward the area where

the Bab Basin subsequently developed [Droste, 2010; Pierson et al., 2010]. Pierson et al.

[2010] documented the thickness variation based on the well data. It is 6–9 m thick in the

area that subsequently became the distal central part of the basin (including the studied core

site) and gradually thickens toward the area that subsequently became the proximal margin,

to reach its maximum thickness of ~21 m [figure 13; Pierson et al., 2010]. This suggests that

the antecedent topography was flat and that the entire platform existed in a shallow-marine

environment, which enabled the subsequent synchronous proliferation of Lithocodium–

Bacinella over the entire area. This is confirmed by seismic data indicating that the Hawar

Member forms a flat-lying reflection package in the Bab Basin [Yose et al., 2006;

Strohmenger et al., 2010].

The Arabian Plate was exposed immediately before the deposition of the Hawar Member

(segment C3; early Early Aptian) [e.g. van Buchem et al., 2002], when lowstand-wedge

deposits, composed mainly of buildups dominated by Lithocodium–Bacinella and rudists,

formed at the southeastern plate margin facing the Neo-Tethys Ocean [Hillgärtner et al.,

2003; Hillgärtner, 2010]. As the sea level rose, the bloom of Lithocodium–Bacinella spread

over the proto-Bab Basin at the beginning of segment C4, which corresponds to the onset of

the positive δ13C excursion due to significant removal of organic carbon from the ocean–

atmosphere system (Figure 2). This bloom was likely triggered by an increase in the nutrient

level.

5.5. Basin-Water Evolution

The 87Sr/86Sr evolution of global seawater shows a decreasing trend from the Late

Barremian to the Aptian/Albian boundary [Howarth and McArthur, 1997; McArthur et al.,

2001]. Intensified submarine hydrothermal activity related to the OJP formation is thought to

be the main cause of this decrease [Bralower et al., 1997; Jones and Jenkyns, 2001].

The 87Sr/86Sr profile of the studied core shows an overall decreasing trend that is comparable

to the global 87Sr/86Sr evolution except for a gradual increase upsection in units 5 through 7

(Figure 2). The higher 87Sr/86Sr compared with those of global seawater are thought to be

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caused by the local influx of isotopically heavier strontium derived from continental crust of

the Arabian Shield and/or its associated sedimentary rocks into the (proto-)Bab Basin,

probably through rivers. The higher 87Sr/86Sr interval mostly correspond to segments C3–C6,

which suggests that the local increase in 87Sr/86Sr occurred in the (proto-)Bab Basin during

the period of OAE1a.

Osmium (Os)-isotope profile through a section recording OAE1a shows two consecutive

sharp decreases in 187Os/188Os during the OAE, which are interpreted as increases in mantle-

derived osmium from the Ontong Java Plateau [Tejada et al., 2009; Bottini et al., 2012]. They

explained the reversal in 187Os/188Os between these two pulses as a transient weathering pulse

due to higher global temperatures, an increased hydrological cycle, and subsequently

increased chemical weathering. Calcium (Ca)-isotope data support indications from osmium

isotopes for a change in weathering during OAE1a [Blättler et al., 2011]. The interval

characterized by the local influx of isotopically heavier strontium into the (proto-)Bab Basin

coincides with that delineated by the δ18O decrease (segment O1). Therefore, the local influx

is likely to be related to intensified terrestrial weathering caused by global warming.

Possible candidates to explain this anomaly (higher 87Sr/86Sr in units 5 through 7) include

the limited connection between the (proto-)Bab Basin and the outer ocean (Neo-Tethys). The

limited connection is supported by the rare occurrence of nannofossils and the absence of

planktonic foraminifers from these units (Figure 2). However, the rare occurrence of

nannofossils may be due, at least in part, to the Early Aptian nannoconid crisis. In spite of the

limited connection supposed, the δ13C profile of the studied core correlates well with that

from other regions (Figure 5). As the sea level rose, the influx of isotopically heavier

strontium decreased, and 87Sr/86Sr became approximate to that of global seawater during

deposition of units 8 through 10.

It is possible that the studied core records a radiogenic excursion that is missing in other

sections used for reconstructing the 87Sr/86Sr evolution of global seawater [Howarth and

McArthur, 1997; McArthur et al., 2001]. This is supported by relatively radiogenic Sr-isotope

values (higher 87Sr/86Sr) reported from the interval that is correlative to segments C3 and C4

in the Resolution Guyot profile [Jenkyns, 2010, fig. 3]. Further studies are needed in

stratigraphically expanded sections to determine whether the global radiogenic excursion

occurred during the OAE1a.

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The abundant occurrence of planktonic foraminifers in unit 10, characterized by organic

carbon-rich carbonates (Figure 3C and 3D), reflects a distinct ocean stratification in the Bab

Basin during deposition of this unit. The basin water was composed of oxic surface water, in

which planktonic foraminifers lived, and dysoxic to anoxic bottom water rich in organic

carbon. A possible explanation for such bottom water includes high-salinity (= high-density)

water formed as the result of active evaporation, which was probably enhanced by a

decreased influx of river water into the basin, as indicated by the 87Sr/86Sr in unit 10, which

are similar to that of the outer oceans.

5.6. Biotic Responses to the Increased Nutrient Level

The Early Aptian in the (proto-)Bab Basin is characterized by three major episodes of

biotic proliferation, each corresponding to the second-order depositional sequence (Figure

1B): orbitolinid foraminifers during the earliest transgression, Lithocodium–Bacinella during

the early to late transgression, and rudists during the early highstand (limited in its

distribution to the proximal margin of the Bab Basin) [e.g. van Buchem et al., 2002, 2010].

The Lower Aptian orbitolinid-rich beds have been reported from various areas around the

circum Neo-Tethys margin, such as France [Arnaud-Vanneau and Arnaud, 1990], Switzerland

[Funk et al., 1993], Spain [Vilas et al., 1995; Ruiz-Ortiz and Castro, 1998], and Oman [Masse

et al., 1998; Immenhauser et al., 1999; Pittet et al., 2002; van Buchem et al., 2002;

Hillgärtner et al., 2003]. These orbitolinid-rich beds generally accumulated during

transgression [e.g. Vilas et al., 1995; Pittet et al., 2002; Burla et al., 2008]. There are three

intervals rich in orbitolinid foraminifers on the Arabian platform: the Lower Kharaib (Upper

Barremian) and Hawar (Lower Aptian) members of the Kharaib Formation and the Nahr Umr

Formation (Albian) [Pittet et al., 2002]. The deposits, all of which accumulated during

periods of transgression, share argillaceous lithologies and high gamma-ray intensities. Units

4 and 5, correlative to the Hawar Member, are rich in orbitolinid foraminifers and are

characterized by their argillaceous lithology, which suggests an increased influx of

terrigenous material into, as well as elevated nutrient levels in the proto-Bab Basin during

deposition of this member. Our δ13C and δ18O records revealed that the timing of the main

phase of orbitolinid proliferation (unit 5) coincides with the period of segment C3 and

segment O1 (early stage) (Figure 2). Therefore, it is likely that the global warming intensified

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weathering, which increased nutrient levels in the oceans and enhanced the proliferation of

orbitolinid foraminifers in the circum Neo-Tethys margin at the early stage of OAE1a. It is

noteworthy that some modern larger foraminifers, such as Operculina, exhibit a preference

for a muddy substrate [Hohenegger et al., 1999] and a nutrient-rich environment [Langer and

Lipps, 2003].

The Lithocodium–Bacinella buildups initially formed only along the southeastern margin

of the Arabian Plate during the early Early Aptian [Hillgärtner et al., 2003; Hillgärtner,

2010]; this was followed by their widespread bloom into the southern Arabian Gulf region,

including the proto-Bab Basin during the period of segment C4. It was inferred that this

bloom occurred under mesotrophic/eutrophic conditions [Immenhauser et al., 2005; Rameil

et al., 2010]. It is unknown why the orbitolinid foraminifers were replaced by Lithocodium–

Bacinella under the continued high-nutrient conditions.

Subsequently, the Lithocodium–Bacinella buildups, catching up with the increase in sea-

level, continued to grow only at the proximal margin (Figure 1B). In contrast, the buildups

ceased around the studied core site. As a result of the significant difference in the

sedimentation rates between the proximal and the distal central parts of the basin, a distinct

topographic depression (Bab Basin) was formed. The platform carbonates dominated by

Lithocodium–Bacinella were overlain by rudist-dominated facies in the proximal margin of

the Bab Basin [e.g. Strohmenger et al., 2010]. This biotic change occurred around the MFS

K80 in the upper Lower Aptian (Figure 4). As rudists were suspension feeders, a trophic

mode well adapted to nutrient rich biotopes [Gili et al., 1995], the replacement was likely

caused by some factors other than nutrient levels, such as substrates, sedimentation, energy

regime, and water chemistry. However, further investigations are needed to specify the main

controlling factor for the replacement.

5.7. Variations in the δ13C Values Influenced by Local Paleoenvironmental

Settings

Although the published Lower Aptian δ13C profiles show more or less similar trends

around the globe, their δ13C values show differences between localities (Figures 4 and 5) [e.g.

Menegatti et al., 1998; Jenkyns and Wilson, 1999; van Buchem et al., 2002; Strophmenger et

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al., 2010; Vahrenkamp, 2010; Kuhnt et al., 2011]. In the proximal margin of the Bab Basin,

the lowest δ13C values are ~1–2‰ in segment C3 located in the Hawar Member [Vahrenkamp,

1996, 2010; Droste, 2010; Strohmenger et al., 2010]. The studied core from the distal central

site in the Bab Basin also has a similar lowest value of 1.6‰ in segment C3. These values are

similar to those from open-marine sections in the Neo-Tethys Ocean but greater than those

from the Resolution Guyot, Mid-Pacific Mountains [Jenkyns and Wilson, 1999] (Figure 5). In

contrast, the highest values known from segment C7 show spatial variations within the Bab

Basin. The values exceed 5‰, occasionally reaching 6‰, in the proximal margin, where

rudist-dominated facies were well developed on top of the platform interior (Figure 4)

[Vahrenkamp, 1996, 2010; Al-Ghamdi and Read, 2010; Droste, 2010; Strohmenger et al.,

2010], whereas they reach 4.2‰ in the organic carbon-rich carbonates that contain abundant

planktonic foraminifers (unit 10) in the studied core (Figure 4). This value is similar to those

(4–5‰) from open-marine sections in the Tethyan and Pacific oceans (Figure 5). The similar

highest values of ~4.3‰ and 4.4‰ were reported from periplatform carbonates in the

“Lekhwair-7 well” [van Buchem et al., 2002] and “Well E” [Strohmenger et al., 2010], both

of which are located at the near-platform site in the basin (Figure 4).

The spatial variations in the δ13C values appear to be related to the development of the

Bab Basin. During the period of segment C3, there were no distinct topographic

differentiations or paleoenvironmental variations in the proto-Bab Basin. Thus, the δ13C

values are very similar throughout the proto-basin. A platform-and-basin topography existed

during deposition of segment C7. Yose et al. [2006] reported locally deposited organic

carbon-rich sediments among the rudist-dominated facies on the top of the platform interior.

The local removal of organic carbon was probably the cause of the anomalously heavy δ13C

values for the rudist-dominated facies [Vahrenkamp, 2010]. Strohmenger et al. [2010]

demonstrated a systematic change in the δ13C values from three wells along a transect from

the platform interior to the basin (Figures 1 and 4) and reported that the decreasing trend in

the δ13C values toward the distal area reflected the decrease in aragonitic material of rudist

shells, because aragonite has higher δ13C values than co-precipitated calcite [Romanek et al.,

1992; Swart and Eberli, 2005]. Our data support this interpretation, because bioclasts that

were originally aragonitic constitute a minor component of the studied core. Acc

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It was pointed out that two mechanisms, geochemically altered platform-top water masses

and the effects of early meteoric diagenesis on carbon-isotope composition, result in the

formation of local positive isotope shifts in shallow-marine settings. [Immenhauser et al.,

2003]. However, the platform-top carbonates have heavier δ13C values than those in slope

and basin carbonates in the study site. These indicate that, although spatial variations in

δ13CDIC values are one of the important factors when interpreting the geochemical record of

ancient epeiric seas, the effects are different depending on oceanographic and

sedimentological settings.

5.8. Removal of Organic Carbon in Response to the Second-Order Sea-Level

Changes

Secular variations in δ13C values of marine carbonates have generally been interpreted as

an approximation of changes in the fraction of buried organic carbon [e.g. Holser et al., 1988;

Kump and Arthur, 1999]. In the Early Aptian, the stepwise positive excursion in segments

C4–C6 and the subsequent gradual positive excursion in segment C7 indicate that the

removal of organic carbon from the ocean–atmosphere system was most significant in the

main phase of OAE1a (segments C4–C6), and that, even after this, the removal continued, to

some extent, on a global scale. It was pointed out that the timing of OAE1a coincides with a

period of sea-level rise [Haq et al., 1988; Erbacher et al., 1996; Vahrenkamp, 1996; Heldt et

al., 2008]. In the studied core, the positive δ13C excursion begins near the boundary between

units 5 and 6 and continues through the upper part of unit 10 (Figure 2). Unit 6, bearing

Lithocodium–Bacinella, grades upward into the mudstone of unit 7, which is overlain by

mudstone/wackestone with siliciclastic fraction of units 8 and 9, and unit 10 is characterized

by abundant planktonic foraminifers. This lithologic succession clearly represents a

deepening upward sequence. Units 6 through 9 correspond to the transgressive systems tract

of the second-order depositional sequence (Figures 1B and 2) [e.g. van Buchem et al., 2010].

The maximum flooding surface (MFS K80 of Sharland et al., [2001]) is interpreted to occur

at the boundary between units 9 and 10. The strontium-isotope stratigraphy and

biostratigraphy indicate that units 6–10 range from the early Early to late Early Aptian.

Consequently, our data suggest that the global removal of organic carbon began during the

initial stage of the second-order transgression in the early Early Aptian and lasted until the

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initial stage of the highstand in the late Early Aptian (Figure 4). As Vahrenkamp [2010]

pointed out, the Livello Selli, often equated with OAE1a, is the initial stage of the global

event characterized by the continued removal of organic carbon during the Early Aptian.

6. Conclusions

(1) A high resolution carbon- and oxygen-isotope stratigraphies, constrained by strontium-

isotope stratigraphy and biostratigraphy, is established for the Lower Aptian carbonates from

a distal central site in the Bab Basin. The δ13C profile, representing global carbon-cycle

perturbations across OAE1a, correlates well with those from the proximal margin of the Bab

Basin and other open-marine sections. The δ18O profile shows characteristic decrease and

increase associated with a negative spike in between, which were identified in other

published δ18O profiles from the Tethyan and Pacific regions.

(2) A long-lasting negative shift in δ13C values (segment C3), which defines the onset of

OAE1a, in the Lower Aptian orbitolinid-rich carbonates was likely caused by massive

volcanic CO2 emission and/or intermittent methane dissociation. The subsequent stepwise

positive δ13C excursion (segments C4–C6) was caused by significant removal of organic

carbon. A synchronous bloom of Lithocodium–Bacinella across the proto-Bab Basin occurred

at the beginning of segment C4.

(3) The carboantes characterized by a proliferation of orbitolinid foraminifers (unit 5) or

Lithocodium–Bacinella (unit 6) mostly show higher 87Sr/86Sr compared with those in the

global oceans; this was likely caused by a local influx of isotopically heavier strontium, along

with nutrients, into the (proto-)Bab Basin. The interval with the higher 87Sr/86Sr coincides

with that delineated by the δ18O decrease, suggesting that the local influx is likely to be

related to intensified terrestrial weathering caused by global warming. The increased nutrient

level probably triggered these biotic proliferations.

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(4) Differences in the δ13C values between proximal and distal sites in the Bab Basin became

greater as the platform-and-basin topography formed, responding to the Early Aptian rise in

sea-level. The minimum δ13C values from segment C3 are similar among sites extending

from the platform to the basin. The values are similar to those from open-marine sections in

the Neo-Tethys and Pacific oceans, although the maximum δ13C value from segment C7 is

greater in the platform compared with the basin because of local removal of organic carbon

on the top of the platform interior.

(5) The δ13C profile and the lithostratigraphy clearly indicate that the global removal of

organic carbon of OAE1a began during the initial stage of the second-order transgression in

the early Early Aptian and lasted until the initial stage of the highstand in the late Early

Aptian. The Livello Selli was the initial stage of the long-term positive δ13C excursion, which

corresponds to the early stage of the second-order transgression.

Acknowledgments

We are most grateful to Abu Dhabi National Oil Company and Abu Dhabi Oil Company,

Ltd (Japan) for their support and permission to publish this work. We also thank A. Misaki

for identifying ammonites. Deep appreciation is expressed to T. Yamada for assistance with

the carbon- and oxygen-isotope measurements. The manuscript was significantly improved

by the comments and suggestions from J. Baker, A. Immenhauser and E. Erba.

References

Allan, J. R., and R. K. Matthews (1982), Isotope signatures associated with early meteoric

diagenesis, Sedimentology, 29(6), 797–817, doi:10.1111/j.1365-3091.1982.tb00085.x.

Al-Ghamdi, N., and F. J. Read (2010), Facies-based sequence-stratigraphic framework of the

Lower Cretaceous rudist platform, Shu’aiba Formation, Saudi Arabia, in Aptian

Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate, GeoArabia Spec. Publ.,

vol. 4, edited by F. S. P. van Buchem et al., pp. 367–410, Gulf PetroLink, Bahrain.

Acc

epte

d A

rticl

e


Al-Husseini, M. I., and R. K. Matthews (2010), Tuning Late Barremian–Aptian Arabian Plate

and global sequences with orbital periods, In Aptian Stratigraphy and Petroleum Habitat

of the Eastern Arabian Plate, GeoArabia Spec. Publ. vol. 4, edited by F. S. P. van

Buchem et al., pp. 199–228, Gulf PetroLink, Bahrain.

Ando, A., T. Kakegawa, R. Takashima, and T. Saito (2002), New perspective on Aptian

carbon isotope stratigraphy: Data from δ13C records of terrestrial organic matter, Geology,

30(3), 227–230, doi: 10.1130/0091-7613(2002)030<0227:NPOACI>2.0.CO;2.

Ando, A., K. Kaiho, H. Kawahata, and T. Kakegawa (2008), Timing and magnitude of early

Aptian extreme warming: Unraveling primary δ18O variation in indurated pelagic

carbonates at Deep Sea Drilling Project Site 463, central Pacific Ocean, Palaeogeogr.

Palaeoclimatol. Palaeoecol., 260(3–4), 463–476, doi:10.1016/j.palaeo.2007.12.007.

Arnaud-Vanneau, A., and H. Arnaud (1990), Hauterivian to Lower Aptian carbonate shelf

sedimentation and sequence stratigraphy in the Jura and northern Subalpine chains

(southeastern France and Swiss Jura), in Carbonate Platforms: Facies, Sequences and

Evolution, Spec. Publ. Intl. Ass. Sedimentol., no. 9, edited by M. E. Tucker et al., pp. 203–

234, Blackwell Sci. Publ., Oxford, doi: 10.1002/9781444303834.ch8.

Arthur, M. A., H. C. Jenkyns, H.-J. Brumsack, and S. O. Schlanger (1990), Stratigraphy,

geochemistry, and paleoceanography of organic-carbon-rich Cretaceous sequences, in

Cretaceous Resources, Events and Rhythms, NATO ASI Ser., vol. 304, edited by R. N.

Ginsburg and B. Beaudoin, pp. 75–119, Kluwer Acad. Publ., Dordrecht.

Asahara, Y., H. Ishiguro, T. Tanaka, K. Yamamoto, K. Mimura, M. Minami, and H. Yoshida

(2006), Application of Sr isotopes to geochemical mapping and provenance analysis: the

case of Aichi Prefecture, central Japan. Appl. Geochem., 21(3), 419–436, doi:

10.1016/j.apgeochem.2005.12.003.

Asahara, Y., T. Tanaka, H. Kamioka, A. Nishimura, and T. Yamazaki (1999), Provenance of

the north Pacific sediments and process of source material transport as derived from Rb–

Sr isotopic systematics. Chem. Geol., 158(3–4), 271–291, doi: 10.1016/S0009-

2541(99)00056-X.

Beerling, D. J., M. R. Lomas, and D. R. Gröcke (2002), On the nature of methane gas-

hydrate dissociation during the Toarcian and Aptian oceanic anoxic events, Am. J. Sci.,

302(1), 28–49, doi:10.2475/ajs.302.1.28.

Bellanca, A., E. Erba, R. Neri, I. P. Silva, M. Sprovieri, F. Tremolada, and D. Verga (2002),

Palaeoceanographic significance of the Tethyan ‘Livello Selli’ (Early Aptian) from the

Acc

epte

d A

rticl

e


Hybla Formation, northwestern Sicily: biostratigraphy and highresolution

chemostratigraphic records, Palaeogeogr. Palaeoclimatol. Palaeoecol., 185(1–2), 175–

196, doi:10.1016/S0031-0182(02)00299-7.

Blättler C. L., H. C. Jenkyns, L. M. Reynard, and G. M. Henderson (2011), Significant

increases in global weathering during Oceanic Anoxic Events 1a and 2 indicated by

calcium isotopes, Earth Planet. Sci. Lett., 309(1–2), 77–88, doi:

10.1016/j.epsl.2011.06.029.

Bosellini, A., M. Morsilli, and C. Neri (1999), Long-term event stratigraphy of the Apulia

Platform margin (Upper Jurassic to Eocene, Gargano, southern Italy), J. Sedim. Res.,

69(6), 1241–1252, doi: 10.2110/jsr.69.124.

Bottini, C., A. S. Cohen, E. Erba, H. C. Jenkyns, and A. L. Coe (2012), Osmium-isotope

evidence for volcanism, weathering, and ocean mixing during the early Aptian OAE 1a,

Geology, 40(7), 583–586, doi: 10.1130/G33140.1.

Bown, P. R., D. C. Rutledge, J. A. Crux, and L. T. Gallagher (1998), Lower Cretaceous, in

Calcareous nannofossil biostratigraphy, British Micropalaeontological Society

Publication Series, edited by P. R. Bown, pp. 86–131, Chapman & Hall, London.

Bralower, T. J., R. M. Leckie, W. V. Sliter, and H. R. Thierstein (1995), An integrated

Cretaceous microfossil biostratigraphy, in Geochronology, Time Scales and Global

Stratigraphic Correlation, SEPM Spec. Publ., vol. 54, edited by W. A. Berggren et al., pp.

65–79, SEPM, Tulsa, Oklahoma.

Bralower, T. J., P. D. Fullager, C. K. Paull, G. S. Dwyer, and R. M. Leckie (1997), Mid-

Cretaceous strontium-isotope stratigraphy of deep-sea sections, Geol. Soc. Amer. Bull.,

109(11), 1421–1442, doi: 10.1130/0016-7606(1997)109<1421:MCSISO>2.3.CO;2.

Bralower, T. J., E. CoBabe, B. Clement, W. V. Sliter, C. L. Osburn, and J. Longoria (1999),

The record of global change in Mid-Cretaceous (Barremian-Albian) sections from the

Sierra Madre, northeastern Mexico, J. Foram. Res., 29(4), 418–437.

Burla, S., U. Heimhofer, P. A. Hochuli, H. Weissert, and P. Skelton (2008), Changes in

sedimentary patterns of coastal and deep-sea successions from the North Atlantic

(Portugal) linked to Early Cretaceous environmental change, Palaeogeogr.

Palaeoclimatol. Palaeoecol., 257(1–2), 38–57, doi:10.1016/j.palaeo.2007.09.010.

Danelian, T., H. Tsikos, S. Gardin, F. Baudin, J.-P. Bellier, and L. Emmanuel (2004), Global

and regional palaeoceanographic changes as recorded in the Mid-Cretaceous (Aptian-

Acc

epte

d A

rticl

e


Albian) sequence of the Ionian zone (NW Greece), J. Geol. Soc. London, 161(4), 703–

709, doi:10.1144/0016-764903-088.

de Gea, G. A., J. M. Castro, R. Aguado, P. A. Ruiz-Ortiz, and M. Company (2003), Lower

Aptian carbon isotope stratigraphy from a distal carbonate shelf setting: the Cau section,

Prebetic zone, SE Spain, Palaeogeogr. Palaeoclimatol. Palaeoecol., 200(1–4), 207–

219, doi:10.1016/S0031-0182(03)00451-6.

Denison, R. E., R. B. Koepnick, A. Fletcher, M. W. Howell, and W. S. Callaway (1994),

Criteria for the retention of original seawater 87Sr/86Sr in ancient shelf limestones, Chem.

Geol., 112(1–2), 131–143, doi: 10.1016/0009-2541(94)90110-4.

Droste, H. J. (2010), Sequence-stratigraphic framework of the Aptian Shu’aiba Formation in

Oman, in Aptian Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate,

GeoArabia Spec. Publ. vol. 4, edited by F. S. P. van Buchem et al., pp. 229–283, Gulf

PetroLink, Bahrain.

Dumitrescu, M., and S. C. Brassell (2006), Compositional and isotopic characteristics of

organic matter for the early Aptian oceanic anoxic event at Shatsky Rise, ODP Leg 198,

Palaeogeogr. Palaeoclimatol. Palaeoecol., 235(1–3), 168–191,

doi:10.1016/j.palaeo.2005.09.028.

Dupraz, C., and A. Strasser (1999), Microbialites and micro-encrusters in shallow coral

bioherms (Middle to Late Oxfordian, Swiss Jura Mountains), Facies, 40(1), 101–130,

doi: 10.1007/BF02537471.

Erba, E. (1994), Nannofossils and superplumes: The Early Aptian “nannoconid crisis”,

Paleoceanography, 9(3), 483–501, doi:10.1029/94PA00258.

Erba, E. (2004), Calcareous nannofossils and Mesozoic oceanic anoxic events, Mar.

Micropaleontol., 52(1–4), 85–106, doi:10.1016/j.marmicro.2004.04.007.

Erba, E., and F. Tremolada (2004), Nannofossil carbonate fluxes during the Early Cretaceous:

Phytoplankton response to nutrification episodes, atmospheric CO2, and anoxia,

Paleoceanography, 19(3), PA1008, doi:10.1029/2003PA000884.

Erba, E., J. E. T. Channell, M. Claps, C. Jones, R. L. Larson, B. N. Opdyke, I. P. Silva, A.

Riva, G. Salvini, and S. Torricelli (1999), Integrated stratigraphy of the Cismon

APTICORE (Southern Alps, Italy): a “reference section” for the Barremian–Aptian

interval at low latitudes, J. Foram. Res., 29(4), 371–391. Acc

epte

d A

rticl

e


Erba, E., C. Bottini, H. J. Weissert, and E. Keller (2010), Calcareous nannoplankton response

to surface-water acidification around Ocean Anoxic Event 1a, Science, 329(5990), 428–

432, doi: 10.1126/science.1188886.

Erbacher, J., J. Thurow, and R. Littke (1996), Evolution patterns of radiolarian and organic

matter variations: A new approach to identify sea-level changes in mid-Cretaceous

pelagic environments, Geology, 24(6), 499–502, doi: 10.1130/0091-

7613(1996)024<0499:EPORAO>2.3.CO;2.

Föllmi, K. B., H. Weissert, M. Bisping, and H. Funk (1994), Phosphogenesis, carbon-isotope

stratigraphy, and carbonate-platform evolution along the Lower Cretaceous northern

Tethyan margin, Geol. Soc. Amer. Bull., 106(6), 729–746, doi:10.1130/0016-

7606(1994)106<0729:PCISAC>2.3.CO;2.

Föllmi, K. B., A. Godet, S. Bodin, and P. Linder (2006), Interactions between environmental

change and shallow water carbonate buildup along the northern Tethyan margin and their

impact on the Early Cretaceous carbon isotope record, Paleoceanography, 21, PA4211,

doi:10.1029/2006PA001313.

Funk, H. P., K. B. Föllmi, and H. Mohr (1993), Evolution of the Tithonian-Aptian carbonate

platform along the northern Tethyan margin, eastern Helvetic Alps, in Cretaceous

Carbonate Platforms, Amer. Assoc. Petrol. Geol. Mem, no. 56, edited by J. A. T. Simo et

al., pp. 387–408, Amer. Assoc. Petrol. Geol., Tulsa.

Gili, E., J.-P. Masse, and P. W. Skelton (1995), Rudists as gregarious sediment-dwellers, not

reef-builders, on Cretaceous carbonate platforms, Palaeogeogr. Palaeoclimatol.

Palaeoecol., 118(3–4), 245–267, doi: 10.1016/0031-0182(95)00006-X.

Gradstein, F. M., J. G. Ogg, and A. G. Smith (Eds.) (2004), A Geologic Time Scale 2004,

Cambridge University Press, Cambridge.

Grauges, A., J. A. Moreno-Bedmar, and R. Martínez (2010), Desmocerátidos (Ammonoidea)

del Aptiense Inferior (Cretácico Inferior) de la subcuenca de Oliete, Cordillera Ibérica

Oriental (Teruel, España). [Lower Aptian (Lower Cretaceous) desmoceratids

(Ammonoidea) of the Oliete sub-basin, Iberian Range (Teruel, Spain).] Revista Española

de Paleontología, 25(1), 7–18.

Graziano, R. (1999), The Early Cretaceous drowning unconformities of the Apulia carbonate

platform (Gargano Promontory, southern Italy): local fingerprints of the global

palaeoceanographic events, Terra Nova, 11(6), 245–250, doi: 10.1046/j.1365-

3121.1999.00256.x.

Acc

epte

d A

rticl

e


Grötsch, J., I. Billing, and V. Vahrenkamp (1998), Carbon-isotope stratigraphy in shallow-

water carbonates: implications for Cretaceous black-shale deposition, Sedimentology,

45(4), 623–634, doi:10.1046/j.1365-3091.1998.00158.x.

Haq, B. U., J. Hardenbol, and P. R. Vail (1988), Mesozoic and Cenozoic chronostratigraphy

and cycles of sea-level change, in Sea-Level Changes: An Integrated Approach, Soc.

Econ. Palaeontol. Mineral. Spec. Publ.,no. 42, edited by C. K. Wilgus et al., pp. 71–108,

Soc. Econ. Palaeontol. Mineral., Tulsa.

Heimhofer, U., P. A. Hochuli, J. O. Herrle, N. Andersen, and H. Weissert (2004), Absence of

major vegetation and palaeoatmospheric pCO2 changes associated with oceanic anoxic

event 1a (Early Aptian, SE France), Earth Planet. Sci. Lett., 223(3–4), 303–318,

doi:10.1016/j.epsl.2004.04.037.

Heldt, M., M. Bachmann, and J. Lehmann (2008), Microfacies, biostratigraphy, and

geochemistry of the hemipelagic Barremian–Aptian in north-central Tunisia: influence of

the OAE 1a on the southern Tethys margin, Palaeogeogr. Palaeoclimatol. Palaeoecol.,

261(3–4), 246–260, doi:10.1016/j.palaeo.2008.01.013.

Hillgärtner, H. (2010), Anatomy of a microbially constructed, high-energy, ocean-facing

carbonate platform margin (earliest Aptian, northern Oman Mountains), in Aptian

Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate, GeoArabia Spec. Publ.,

vol. 4, edited by F. S. P. van Buchem et al., pp. 285–300, Gulf PetroLink, Bahrain.

Hillgärtner, H., F. S. P. van Buchem, F. Gaumet, P. Razin, B. Pittet, J. Grötsch, and H. J.

Droste (2003), The Barremian-Aptian evolution of the eastern Arabian carbonate platform

margin (northern Oman), J. Sedim. Res., 73(5), 756–773, doi: 10.1306/030503730756.

Hohenegger J., E. Yordanova, Y. Nakano, and F. Tatzreiter (1999), Habitats of larger

foraminifera on the upper reef slope of Sesoko Island, Okinawa, Japan, Mar.

Micropaleontol., 36(2–3), 109–168, doi: org/10.1016/S0377-8398(98)00030-9.

Holser, W. T., M. Schidlowski, F. T. Mackenzie, and J. B. Maynard (1988), Biogeochemical

cycles of carbon and sulfur, in Chemical Cycles in the Evolution of the Earth, edited by C.

B. Gregor, et al., pp. 105–173, Wiley, New York.

Howarth, R. J., and J. M. McArthur (1997), Statistics for strontium isotope stratigraphy: A

robust LOWESS fit to the marine strontium isotope curve for the period 0 to 206 Ma,

with look-up table for the derivation of numerical age, J. Geol., 105(4), 441–456,

doi:10.1086/515938.

Acc

epte

d A

rticl

e


Hu, X., K. Zhao, I. O. Yilmaz, and Y. Li (2012), Stratigraphic transition and

palaeoenvironmental changes from the Aptian oceanic anoxic event 1a (OAE1a) to the

oceanic red bed 1 (ORB1) in the Yenicesihlar section, central Turkey, Cret. Res., 38, 40–

51, doi:10.1016/j.cretres.2012.01.007.

Huck, S., N. Rameil, T. Korbar, U. Heimhofer, T. D. Wieczorek, and A. Immenhauser (2010),

Latitudinally different responses of Tethyan shoal-water carbonate systems to the Early

Aptian oceanic anoxic event (OAE 1a), Sedimentology, 57(7), 1585–1614,

doi:10.1111/j.1365-3091.2010.01157.x.

Huck, S., U. Heimhofer, N. Rameil, S. Bodin, and A. Immenhauser (2011), Strontium and

carbon-isotope chronostratigraphy of Barremian–Aptian shoal-water carbonates: Northern

Tethyan platform drowning predates OAE 1a, Earth Planet. Sci. Lett., 304(3–4), 547–558,

doi:10.1016/j.epsl.2011.02.031.

Hughes, G. W. (2000), Bioecostratigraphy of the Shu’aiba Formation, Shaybah field, Saudi

Arabia, GeoArabia, 5(4), 545–578.

Immenhauser, A., W. Schlager, S. J. Burns, R. W. Scott, T. Geel, J. Lehmann, S. van der

Gaast, and L. J. A. Bolder-Schruwer (1999), Late Aptian to Late Albian sea-level

fluctuations constrained by geochemical and biological evidence (Nahr Umr Formation,

Oman), J. Sedim. Res., 69(2), 434–466, doi: 10.2110/jsr.69.434.

Immenhauser, A., G. D. Porta, J. A. M. Kenter, and J. R. Bahamonde (2003), An alternative

model for positive shifts in shallow-marine carbonate δ13C and δ18O, Sedimentology,

50(5), 953–959, doi:10.1046/j.1365-3091.2003.00590.x.

Immenhauser, A., H. Hillgärtner, U. Sattler, G. Bertotti, P. Schoepfer, P. Homewood, V.

Vahrenkamp, T. Steuber, J.-P. Masse, H. J. Droste, J. T. van Koppen, B. van der Kooij, E.

van Bentum, K. Verwer, E. H. Strating, W. Swinkels, J. Peters, I. Immenhauser-Potthast,

and S. Al Maskery (2004), Barremian-lower Aptian Qishn Formation, Haushi-Huqf area,

Oman: a new outcrop analogue for the Kharaib/Shu’aiba reservoirs, GeoArabia, 9(1),

153–194.

Immenhauser, A., H. Hillgärtner, and E. van Bentum (2005), Microbial-foraminiferal

episodes in the Early Aptian of the southern Tethyan margin: ecological significance and

possible relation to oceanic anoxic event 1a, Sedimentology, 52(1), 77–99, doi:

10.1111/j.1365-3091.2004.00683.x.

Immenhauser, A., C. Holmden, and W. P. Patterson (2008), Interpreting the carbon-isotope

record of ancient shallow epeiric seas: Lessons from the recent, in Dynamics of Epeiric

Acc

epte

d A

rticl

e


Seas, Geol. Assoc. Canada Spec. Publ., vol. 48, edited by B. R. Pratt and C. Holmden, pp.

135–174, Geol. Assoc. Canada, St. John's.

Jacobsen, S. B., and A. J. Kaufman (1999), The Sr, C and O isotopic evolution of

Neoproterozoic seawater, Chem. Geol., 161(1–3), 37–57, doi: 10.1016/S0009-

2541(99)00080-7.

Jahren, A. H., N. C. Arens, G. Sarmiento, J. Guerrero, and R. Amundson (2001), Terrestrial

record of methane hydrate dissociation in the Early Cretaceous, Geology, 29(2), 159–162,

doi:10.1130/0091-7613(2001)029<0159:TROMHD>2.0.CO;2.

Jenkyns, H. C. (1995), Carbon-isotope stratigraphy and paleoceanographic significance of the

Lower Cretaceous shallow-water carbonates of Resolution Guyot, mid-Pacific Mountains,

In Proc. ODP, Sci. Res., vol. 143, edited by E. L. Winterer et al., pp. 99–104, College

Station, Texas.

Jenkyns, H. C., and P. A. Wilson (1999), Stratigraphy, paleoceanography, and evolution of

Cretaceous Pacific guyots: Relics from a greenhouse Earth, Amer. J. Sci., 299(5), 341–

392, doi:10.2475/ajs.299.5.341.

Jones, C. E., and H. C. Jenkyns (2001), Seawater strontium isotopes, oceanic anoxic events,

and seafloor hydrothermal activity in the Jurassic and Cretaceous, Am. J. Sci., 301(2),

112–149, doi: 10.2475/ajs.301.2.112.

Kuhnt, W., A. Holbourn, and M. Moullade (2011), Transient global cooling at the onset of

early Aptian oceanic anoxic event (OAE) 1a, Geology, 39(4), 323–326,

doi:10.1130/G31554.1.

Kump, L. R., and M. A. Arthur (1999), Interpreting carbon-isotope excursions: carbonates

and organic matter, Chem. Geol., 161(1–3), 181–198, doi:10.1016/S0009-

2541(99)00086-8.

Kuroda, J., M. Tanimizu, R. S. Hori, K. Suzuki, N. O. Ogawa, M. L. G. Tejada, M. F. Coffin,

R. Coccioni, E. Erba, and N. Ohkouchi (2011), Lead isotopic record of Barremian-Aptian

marine sediments: Implications for large igneous provinces and the Aptian climatic crisis,

Earth Planet. Sci. Lett., 307(1–2), 126–134, doi:10.1016/j.epsl.2011.04.021.

Langer, M. R., and J. H., Lipps (2003), Foraminiferal distribution and diversity, Madang Reef

and Lagoon, Papua New Guinea, Coral Reefs, 22(2), 143–154, doi: 10.1007/s00338-003-

0298-1. Acc

epte

d A

rticl

e


Larson, R. L., and E. Erba (1999), Onset of the Mid-Cretaceous greenhouse in the Barremian-

Aptian: Igneous events and the biological, sedimentary, and geochemical responses,

Paleoceanography, 14(6), 663–678, doi:10.1029/1999PA900040.

Li, Y.-X., T. J. Bralower, I. P. Montañez, D. A. Osleger, M. A. Arthur, D. M. Bice, T. D.

Herbert, E. Erba, and I. Premoli Silva (2008), Toward an orbital chronology for the early

Aptian oceanic anoxic event (OAE1a, ~120 Ma), Earth Planet. Sci. Lett., 271(1–4), 88–

100, doi:10.1016/j.epsl.2008.03.055.

Malinverno, A., E. Erba, and T. D. Herbert (2010), Orbital tuning as an inverse problem:

Chronology of the early Aptian oceanic anoxic event 1a (Selli Level) in the Cismon

APTICORE, Paleoceanography, 25, PA2203, doi:10.1029/2009PA001769.

Masse, J.-P., J. Borgomano, and S. Al Maskiry (1998), A platform-to-basin transition for

lower Aptian carbonates (Shuaiba Formation) of the northeastern Jebel Akhdar (Sultanate

of Oman), Sedim. Geol., 119(3–4), 297–309, doi:10.1016/S0037-0738(98)00068-2.

Masse, J.-P. (2002), Integrated stratigraphy of the lower Aptian and applications to carbonate

platforms: a state of the art, in North African Cretaceous Carbonate Platform Systems,

NATO Sci. Ser., IV. Earth Env. Sci., vol. 28, edited by E. Gili et al., pp. 203–214, Kluwer

Acad. Publ., Dordrecht, the Netherlands.

McArthur, J. M., R. J. Howarth, and T. R. Bailey (2001), Strontium isotope stratigraphy:

LOWESS version 3: Best fit to the marine Sr-isotope curve for 0–509 Ma and

accompanying look-up table for deriving numerical age, J. Geol., 109(2), 155–170,

doi:10.1086/319243.

Méhay, S., C. E. Keller, S. M. Bernasconi, H. Weissert, E. Erba, C. Bottini, and P. A. Huchuli

(2009), A volcanic CO2 pulse triggered the Cretaceous oceanic anoxic event 1a and a

biocalcification crisis, Geology, 37(9), 819–822, doi:10.1130/G30100A.1.

Menegatti, A. P., H. Weissert, R. S. Brown, R. V. Tyson, P. Farrimond, A. Strasser, and M.

Caron (1998), High-resolution δ13C stratigraphy through the early Aptian “Livello Selli”

of the Alpine Tethys, Paleoceanography, 13(5), 530–545, doi:10.1029/98PA01793.

Millán, M. I., H. J. Weissert, P. A. Fernández-Mendiola, and J. García-Mondéjar

(2009), Impact of Early Aptian carbon cycle perturbations on evolution of a marine shelf

system in the Basque-Cantabrian Basin (Aralar, N Spain), Earth Planet. Sci. Lett., 287(3–

4), 392–401, doi:10.1016/j.epsl.2009.08.023.

Misumi, K., Y. Yamanaka, and E. Tajika (2009), Numerical simulation of atmospheric and

oceanic biogeochemical cycles to an episodic CO2 release event: Implications for the

Acc

epte

d A

rticl

e


cause of mid-Cretaceous Ocean Anoxic Event-1a, Earth Planet. Sci. Lett., 286(1–2), 316–

323, doi:10.1016/j.epsl.2009.06.045.

Moreno-Bedmar, J. A., M. Company, T. Bover-Arnal, R. Salas, G. Delanoy, F. J.-M. R.

Maurrasse, A. Grauges, and R. Martínez (2010), Lower Aptian ammonite biostratigraphy

in the Maestrat Basin (Eastern Iberian Chain, Eastern Spain). A Tethyan transgressive

record enhanced by synrift subsidence, Geol. Acta, 8(3), 281–299,

doi:10.1344/105.000001534.

Murris, R. J. (1980), Middle East: Stratigraphic evolution and oil habitat, Amer. Assoc. Petrol.

Geol. Bull., 64(5), 597–618, doi:10.1306/2F918A8B-16CE-11D7-8645000102C1865D.

Najarro, M., I. Rosales, J. Martín-Chivelet (2011a), Major palaeoenvironmental perturbation

in an Early Aptian carbonate platform: prelude of the Oceanic Anoxic Event 1a?, Sedim.

Geol., 235(1–2), 50–71, doi:10.1016/j.sedgeo.2010.03.011.

Najarro, M., I. Rosales, J. A. Moreno-Bedmar, G. A. de Gea, E. Barrón, M. Company, and G.

Delanoy (2011b), High-resolution chemo- and biostratigraphic records of the Early

Aptian oceanic anoxic event in Cantabria (N Spain): Palaeoceanographic and

palaeoclimatic implications, Palaeogeogr. Palaeoclimatol. Palaeoecol., 299(1–2), 137–

158, doi:10.1016/j.palaeo.2010.10.042.

Patterson, W. P., and L. M. Walters (1994), Depletion of 13C in seawater ΣCO2 on modern

carbonate platforms: Significance for the carbon isotopic record of carbonates, Geology,

22(10), 885–888, doi:10.1130/0091-7613(1994)022<0885:DOCISC>2.3.CO;2.

Pierson, B. J., G. P. Eberli, K. Al Mehsin, S. Al Menhali, G. Warrlich, H. J. Droste, F. Maurer,

J. Whitworth, and D. Drysdale (2010), Seismic stratigraphy and depositional history of

the Upper Shu’aiba (Late Aptian) in the UAE and Oman, in Aptian Stratigraphy and

Petroleum Habitat of the Eastern Arabian Plate, GeoArabia Spec. Publ., vol. 4, edited by

F. S. P. van Buchem et al., pp. 411–444, Gulf PetroLink, Bahrain.

Pittet, B., F. S. P. van Buchem, H. Hillgärtner, P. Razin, J. Grötsch, and H. J. Droste (2002),

Ecological succession, palaeoenvironmental change, and depositional sequences of

Barremian–Aptian shallow-water carbonates in northern Oman, Sedimentology, 49(3),

555–581, doi: 10.1046/j.1365-3091.2002.00460.x.

Price, G. D. (2003), New constraints upon isotope variation during the early Cretaceous

(Barremian–Cenomanian) from the Pacific Ocean, Geol. Mag., 140(5), 513–522, doi:

10.1017/S0016756803008100.

Acc

epte

d A

rticl

e


Rameil, N., A. Immenhauser, G. Warrlich, H. Hillgärtner, and H. J. Droste (2010),

Morphological patterns of Aptian Lithocodium–Bacinella geobodies: relation to

environment and scale, Sedimentology, 57(3), 883–911, doi:10.1111/j.1365-

3091.2009.01124.x.

Romanek, C. S., E. L. Grossman, and J. W. Morse (1992), Carbon isotopic fractionation in

synthetic aragonite and calcite: Effects of temperature and precipitation rate, Geochim.

Cosmochim. Acta, 56(1), 419–430, doi:10.1016/0016-7037(92)90142-6.

Roth, P. H. (1978), Cretaceous nannoplankton biostratigraphy and oceanography of the

northwestern Atlantic Ocean, in Initial Report of the Deep Sea Drilling Project, vol. 44,

edited by W. E. Benson, R. E. Sheridan et al., pp. 731–759, U. S. Gov. Print. Office,

Washington, D. C.

Ruiz-Ortiz, P. A., and J. M. Castro (1998), Carbonate depositional sequences in shallow to

hemipelagic platform deposits; Aptian, Prebetic of Alicante (SE Spain), Bull. Soc. Géol.

France, 169(1), 21–33.

Sato, T., S. Yuguchi, T. Takayama, and K. Kameo (2004), Drastic change in the geographical

distribution of the cold-water nannofossil Coccolithus pelagicus (Wallich) Schiller at 2.74

Ma in the late Pliocene, with special reference to glaciation in the Arctic Ocean, Mar.

Micropaleontol., 52(1–4), 181–193, doi:10.1016/j.marmicro.2004.05.003.

Schlanger, S. O., and H. C. Jenkyns (1976), Cretaceous oceanic anoxic events: causes and

consequences, Geol. Mijnb., 55(3–4), 179–184.

Schroeder, R., F. S. P. van Buchem, A. Cherchi, D. Baghbani, B. Vincent, A Immenhauser,

and B. Granier (2010), Revised orbitolinid biostratigraphic zonation for the Barremian–

Aptian of the eastern Arabian Plate and implications for regional stratigraphic correlations,

in Aptian Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate, GeoArabia

Spec. Publ., vol. 4, edited by F. S. P. van Buchem et al., pp. 49–96, Gulf PetroLink,

Bahrain.

Sharland, P. R., R. Archer, D. M. Casey, R. B. Davies, S. H. Hall, A. P. Heward, A. D.

Horbury, and M. D. Simmons (2001), Sequence Stratigraphy of the Arabian Plate,

GeoArabia Spec. Publ., vol. 2, Gulf PetroLink, Bahrain.

Skelton, P. W., R. A. Spicer, S. P. Kelley, and I. Gilmour (2003), The Cretaceous World,

Cambridge University Press, Cambridge. Acc

epte

d A

rticl

e


Steuber, T. (1999), Isotopic and chemical intra-shell variations in low-Mg calcite of rudist

bivalves (Mollusca-Hippuritacea): disequilibrium fractionations and Late Cretaceous

seasonality, Int. J. Earth Sci., 88(3), 551–570, doi:10.1007/s005310050284.

Strohmenger, C. J., T. Steuber, A. Ghani, D. G. Barwick, S. H. A. Al Mazrooei, and N. O. Al

Zaabi (2010), Sedimentology and chemostratigraphy of the Hawar and Shu’aiba

depositional sequences, Abu Dhabi, United Arab Emirates, in Aptian Stratigraphy and

Petroleum Habitat of the Eastern Arabian Plate, GeoArabia Spec. Publ., vol. 4, edited by

F. S. P. van Buchem et al., pp. 341–365, Gulf PetroLink, Bahrain.

Suzuki, K., Y. Asahara, K. Miura, and T. Tanaka (2012), Another sea area separated from the

Panthalassic Ocean in the Norian, the Late Triassic: The lowest Sr isotopic composition

of the Ishimaki limestone in central Japan, Chem. Erde, 72(1), 77–84,

doi:10.1016/j.chemer.2011.06.004.

Suzuki, Y., Y. Iryu, S. Inagaki, T. Yamada, S. Aizawa, and D. A. Budd (2006), Origin of atoll

dolomites distinguished by geochemistry and crystal chemistry: Kita-daito-jima, northern

Philippine Sea, Sedim. Geol., 183(3–4), 181–202, doi:10.1016/j.sedgeo.2005.09.016.

Swart, P. K., and G. Eberli (2005), The nature of the δ13C of periplatform sediments:

Implications for stratigraphy and the global carbon cycle, Sedim. Geol., 175(1–4), 115–

129, doi:10.1016/j.sedgeo.2004.12.029.

Tejada, M. L. G., K. Suzuki, J. Kuroda, R. Coccioni, J. J. Mahoney, N. Ohkouchi, T.

Sakamoto, and Y. Tatsumi (2009), Ontong Java Plateau eruption as a trigger for the Early

Aptian oceanic anoxic event, Geology, 37(9), 855–858, doi:10.1130/G25763A.1.

Vahrenkamp, V. C. (1996), Carbon isotope stratigraphy of the Upper Kharaib and Shu’aiba

formations: implications for the Early Cretaceous evolution of the Arabian Gulf Region.

Amer. Assoc. Petrol. Geol. Bull., 80(5), 647–662, doi:10.1306/64ED8868-1724-11D7-

8645000102C1865D.

Vahrenkamp, V. C. (2010), Chemostratigraphy of the Lower Cretaceous Shu’aiba Formation:

A δ13C reference profile for the Aptian Stage from the southern Neo-Tethys Ocean, in

Aptian Stratigraphy and Petroleum Habitat of the Eastern Arabian Plate, GeoArabia

Spec. Publ., vol. 4, edited by F. S. P. van Buchem et al., pp. 107–137, Gulf PetroLink,

Bahrain.

van Breugel, Y., S. Schouten, H. Tsikos, E. Erba, G. D. Price, and J. S. S. Damsté (2007),

Synchronous negative carbon isotope shifts in marine and terrestrial biomarkers at the

onset of the Early Aptian oceanic anoxic event 1a: Evidence for the release of 13C-

Acc

epte

d A

rticl

e


depleted carbon into the atmosphere, Paleoceanography, 22, PA1210,

doi:10.1029/2006PA001341.

van Buchem, F. S. P., B. Pittet, H. Hillgärtner, J. Grotsch, A. I. Al Mansouri, I. M. Billing, H.

J. Droste, W. H. Oterdoom, and M. van Steenwinkel (2002), High-resolution sequence

stratigraphic architecture of Barremian/Aptian carbonate systems in Northern Oman and

the United Arab Emirates (Kharaib and Shu’aiba Formations), GeoArabia, 7(3), 461–500.

van Buchem, F. S. P., M. I. Al Husseini, F. Maurer, H. J. Droste, and L. A. Yose (2010),

Sequence-stratigraphic synthesis of the Barremian–Aptian of the eastern Arabian Plate

and implications for the petroleum habitat, in Aptian Stratigraphy and Petroleum Habitat

of the Eastern Arabian Plate, GeoArabia Spec. Publ., vol. 4, edited by F. S. P. van

Buchem et al., pp. 9–48, Gulf PetroLink, Bahrain.

Vilas, L., J.-P. Masse, and C. Arias (1995), Orbitolina episodes in carbonate platform

evolution: the early Aptian model from SE Spain, Palaeogeogr. Palaeoclimatol.

Palaeoecol., 119(1–2), 35–45, doi:10.1016/0031-0182(95)00058-5.

Weissert, H., and E. Erba, (2004), Volcanism, CO2 and palaeoclimate: a Late Jurassic–Early

Cretaceous carbon and oxygen isotope record, J. Geol. Soc. London, 161(4), 695–702,

doi:10.1144/0016-764903-087.

Wissler, L., H. Funk, and H. Weissert (2003), Response of Early Cretaceous carbonate

platforms to changes in atmospheric carbon dioxide levels, Palaeogeogr. Palaeoclimatol.

Palaeoecol., 200(1–4), 187–205, doi:10.1016/S0031-0182(03)00450-4.

Wu, Y., and S. Chafets (2002), 13C-enriched carbonate in Mississippian mud mounds:

Alamogordo Member, Lake Valley Formation, Sacramento Mountains, New Mexico,

U.S.A., J. Sedim. Res., 72(1), 138–145, doi: 10.1306/040201720138.

Yamamoto, K., R. Asami, and Y. Iryu (2010), Carbon and oxygen isotopic compositions of

modern brachiopod shells from a warm-temperate shelf environment, Sagami Bay, central

Japan, Palaeogeogr. Palaeoclimatol. Palaeoecol., 291(3–4), 348–359,

doi:10.1016/j.palaeo.2010.03.006.

Yose, L. A., A. S. Ruf, C. J. Strohmenger, J. S. Schuelke, A. Gombos, I. Al Hosani, S. Al

Maskary, G. Bloch, Y. Al Mehairi, and I. G. Johnson (2006), Three-dimensional

characterization of a heterogeneous carbonate reservoir, Lower Cretaceous, Abu Dhabi

(United Arab Emirates), in Giant Hydrocarbon Reservoirs of the World: From Rocks to

Reservoir Characterization and Modeling, Amer. Assoc. Petrol. Geol. Mem, no. 88, edited

Acc

epte

d A

rticl

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by P. M. Harris and L. J. Weber, pp. 173–212, Amer. Assoc. Petrol. Geol., Tulsa.

doi:10.1306/1215877M882562.

Yose, L. A., C. J. Strohmenger, I. Al Hosani, A. S. Ruf, A. M. Gombos, G. Bloch, S. S. Al

Maskary and Y. Al Mehairi (2010), Sequence-stratigraphic evolution of an Aptian

carbonate platform (Shu'aiba Formation), southern Arabian Plate, onshore Abu Dhabi,

United Arab Emirates, in Aptian Stratigraphy and Petroleum Habitat of the Eastern

Arabian Plate, GeoArabia Spec. Publ., vol. 4, edited by F. S. P. van Buchem et al., pp.

309–340, Gulf PetroLink, Bahrain.

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Figure 1. A) Paleogeographic map of the southern Arabian Gulf region during the late Early Aptian. Sites of the Lekhwair-7 well [van Buchem et al., 2002], wells C–E [Strohmenger et al., 2010], fields A and Y [Vahrenkamp, 2010] and studied core are shown. B) Stratigraphic outline of the Upepr Barremian to the Middle Albian of the Bab Basin (modified from van Buchem et al. [2010]). Carbonate sequences studied by Strohmenger et al. [2010] and Vahrenkamp [2010] (red square), van Buchem et al. [2002] (yellow square), and this study (green square) are indicated.

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Figure 2. Lithostratigraphy, chemostratigraphy (δ13C, δ18O, 87Sr/86Sr, SrO, MnO, Fe2O3, Sr/Mn, and Mn/Sr profiles), and biostratigraphy of the studied core. C1–C8 segmentation of the δ13C profile is based on Menegatti et al. [1998]. Stratigraphic horizons from which calcareous nannofossils and ammonite were found are indicated. A green arrow indicates the negative δ18O spike comparable to that identified immediately above the uppermost horizon of the Livello Selli by Menegatti et al. [1998]. 87Sr/86Sr from units 5 through 7 is higher than those of global seawater shown by a solid line with broken lines indicating a standard deviation (2σ) [Howarth and McArthur, 1997; McArthur et al., 2001]. Note pink areas denote those for samples determined as diagenetically altered based on criteria defined by Denison et al. [1994] and Jacobsen and Kaufman [1999].

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Figure 3. Photographs of a core slab and thin sections from the studied core. A) Packstone rich in orbitolinid foraminifers (OF); 2656.67 mbsf in unit 5. P, peloid; M, mollusk. B) Lithocodium–Bacinella floatstone; 2655.61 mbsf in unit 6. Ba, Bacinella; BF, benthic foraminifer; E, echinoid. C) Alternating beds of laminated and non-laminated planktonic foraminiferal wackestones rich in organic matter; 2638.37–2638.22 mbsf in unit 10. D) Laminated wackestone rich in planktonic foraminifers (PF); 2636.85 mbsf in unit 10.

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Figure 4. Correlation of the δ13C profiles ranging from proximal to distal sites in the Bab Basin [Strohmenger et al., 2010; Vahrenkamp, 2010; this study]. C1–C8 segmentation of the δ13C profiles are performed by the present authors following Menegatti et al. [1998]. The maximum flooding surface, MFS K80 [Sharland et al., 2001], can be traced throughout the Bab Basin. Note significant lateral differences in the depositional facies and deposit thickness.

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Figure 5. Correlation of the δ13C and 87Sr/86Sr profiles across OAE1a from sites in the Neo-Tethys margin and the Pacific Ocean. (a), Roter Sattel, Switzerland [Menegatti et al., 1998]; (b), La Bédoule, France [Kuhnt et al., 2011]; (c), Bab Basin, UAE (this study); (d), Ocean Drilling Program Site 866, Resolution Guyot, Mid-Pacific Mountains [Jenkyns and Wilson, 1999]. C1–C8 segmentation of the δ13C profiles is performed by the present authors following Menegatti et al. [1998] except for those from Roter Sattel and La Bédoule. Note the long-lasting negative δ13C shift in C3 at La Bédoule and in the Bab Basin.

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Figure 6. Correlation of the δ18O profiles across OAE1a from sites in the Neo-Tethys margin and the Pacific Ocean. (a), Sicily, Italy [Bellanca et al., 2002]; (b), Cismon, Italy [Erba et al., 2010]; (c), Roter Sattel, Switzerland [Menegatti et al.,1998]; (d), La Bédoule, France [Kuhnt et al., 2011]; (e), Yenicesihlar, Turkey [Hu et al., 2012]; (f), Bab Basin, UAE [Vahrenkamp, 2010]; (g), Bab Basin, UAE [This study]; (h), ODP Site 463, Mid-Pacific Mountains [Ando et al., 2008]; (i), Resolution Guyot, Mid-Pacific Mountains [Jenkyns and Wilson, 1999]. A green line in our δ18O profiles represents a simple 5-point moving average. C1–C8 segmentations are performed except for those from Sicily, Cismon, Bab Basin [Vahrenkamp, 2010], and Resolution Guyot by the present authors. O1 and O2 segmentations are performed by the present authors.

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Table 1. Calcareous nannofossils detected from the studied core.

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Table 2. Lithologic features of a carbonate sequence in the studied core.

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Table 3. δ13C and δ18O values of bulk carbonate samples from the studied core.

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Table 4. 87Sr/86Sr and the numerical ages from the studied core.

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