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© 2013 American Geophysical Union. All Rights Reserved.
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
Department of Earth and Planetary Sciences, Graduate School of Environmental Studies, Nagoya University, Furo-cho, Chikusa-ku, Nagoya 464-8601, Japan
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
Department of Earth and Planetary Sciences, Graduate School of Environmental Studies, Nagoya University, Furo-cho, Chikusa-ku, Nagoya 464-0814, Japan
Present address: Institute of Geology and Paleontology, Graduate School of Science, Tohoku University, Aobayama, Sendai 980-0856, Japan (iryu@m.tohoku.ac.jp)
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.
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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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