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

Authors


Abstract

[1] 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 the 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

[2] 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 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].

[3] Although many studies focused on pelagic deposits, their coeval shallow-water platform carbonates were also investigated to assess the 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 (Figures 1a and 1b) [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 intrashelf basin (Bab Basin) was formed that was surrounded by giant carbonate platforms (Figures 1a and 1b) [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 lithostratigraphy 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 (Figures 1a and 1b).

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 upper 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.

[4] This article presents an integrated data set consisting of lower Aptian high resolution, bulk carbonate 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 intrashelf basin at the southern Neo-Tethys margin during the early Aptian.

Figure 2.

Lithostratigraphy, chemostratigraphy (δ13C, δ18O, 87Sr/86Sr, trace element (Sr, Mn and Fe) concentrations, 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].

2 Regional Geological Framework

[5] 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].

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 nonlaminated 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.

[6] 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 intrashelf basin surrounded by carbonate platforms (Figures 1a and 1b). 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

[7] 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 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.

[8] 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.

[9] 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 X-ray diffraction 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 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 the inductively coupled plasma-atomic emission spectrometer. Oxide concentration was calculated from the determined elemental concentration (wt%). 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 concentrations (not as oxides).

[10] 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 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 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 the 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].

Table 1. Calcareous Nannofossils Detected From the Studied Core
Lithologic UnitUnit 12Unit 11Unit 10Unit 9Unit 8Unit 7Unit 5Unit 4
Depth (mbsf)2615.422616.232617.142617.962618.702619.702620.522621.402622.282623.422624.312624.862627.572628.512629.302630.242631.292632.152633.092633.972634.732635.462636.352637.272638.212639.122639.962640.932641.902642.692643.542647.922656.552660.13
Nannofossil Zone [Roth, 1978]NC7B?NC7ANC6B?
Braarudosphaera africana + ++   + +                       
Braarudosphaera bigelowii        +                         
Broinsonia matalosa                 +                
Chiastozygus litterarius +                                
Cretarhabdus conicus                     ++           
Haqius ellipticus     +                            
Hayesites irregularis+++++   +                         
Helenea chiastia ++++   ++      ++ + ++           
Micrantholithus hoschulzii        +     +                   
Nannoconus quadriangulus        +                         
Nannoconus truitti        ++  + +    +         +    
Nannoconus vocontiensis                     +    +       
Nannoconus spp.+++++ +++++    + +   ++ +++  +    
Rhagodiscus angustus  + +                +++    +     
Rhagodiscus asper++ ++  ++++    +++++ +++ ++ ++    
Rhagodiscus gallagheri                       +          
Rotelapillus laffittei +++    +                         
Staurolithites spp. + +    + +      +                
Watznaueria barnesae+++++ + +++++++ +++ ++++++++++++++
Zeugrhabdotus elegans                   +  +           
Zeugrhabdotus embergeri ++++                             
                                  +: present

[11] 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 the Technology Research Center, Japan Oil, Gas and Metals National Corporation, 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 the Technology Research Center, Japan Oil, Gas and Metals National Corporation.

4 Results

4.1 Lithostratigraphy

[12] 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 because no erosional surface indicating a hiatus is observed. X-ray diffraction 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.

Table 2. Lithologic Features of a Carbonate Sequence in the Studied Core
UnitInterval (mbsf)Thickness (m)LithologyRemarks
Unit 122619.94–2614.885.06Argillaceous mudstone with a 1.6 m-thick dolomitized interlayerCommon bioturbation in the lower and middle parts
Argillaceous, as reflected by increased gamma-ray values
Unit 112619.94–2631.6511.71Bioturbated mudstone/wackestoneCommon molluskan shell fragments
Unit 102631.65–2638.807.15Alternating beds of laminated and nonlaminated planktonic foraminiferal wackestones (Figure 3c)Rich in organic carbon (Figure 3d)
Unit 92638.80–2639.911.11Mudstone with ammonitesCommon pressure solution seams and partially pyritized bioclasts
Highly fragmented (<3.5 cm across) ammonites occurring exclusively as molds with siliciclastic fraction
Unit 82639.91–2644.394.48Bioturbated mudstone/wackestoneCommon pressure solution seams and partially pyritized bioclasts
Unit 72644.39–2650.816.42MudstoneAbundant stylolites and pressure solution seams
Unit 62650.81–2656.465.65Lithocodium-Bacinella floatstone/framestoneFramework composed of Lithocodium-Bacinella found in the middle part of this unit (Figure 3b).
Lower and upper parts consisting of floatstone abundant in pebble-sized fragments of Lithocodium-Bacinella associated with peloids and bioclasts of mollusks, benthic foraminifers, and echinoids
Lithocodium-Bacinella fragments decreasing upward in the upper floatstone
Unit 52656.46–2659.693.23Benthic foraminiferal packstoneAbundant tests of orbitolinid foraminifers (Figure 3a) and common peloids and bioclasts of mollusks and echinoids
Unit 42659.69–2661.061.37Benthic foraminiferal packstone/grainstone
Argillaceous
Characterized by high gamma-ray values and well-developed pressure solution seams
Unit 32661.06–2662.741.68Bioclastic wackestoneCharacterized by the occurrence of stylolites and pressure solution seams
Unit 22662.74–2665.953.21Bioclastic wackestone/mudstone
Unit 12665.95–2667.911.96Peloidal-bioclastic grainstone/packstoneRich in peloids, intraclasts, and bioclasts, such as mollusks, echinoids, and benthic foraminifers

[13] 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 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

[14] 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

[15] The δ13C values, which range from 1.6 to 4.2‰, show negative and positive excursions upsection (Figure 2 and 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 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.

Table 3. δ13C and δ18O Values of Bulk Carbonate Samples From the Studied Core
Depth (mbsf)Lithostratigraphic UnitSegment Boundaryδ13C (‰ VPDB)δ18O (‰ VPDB)Depth (mbsf)Lithostratigraphic UnitSegment Boundaryδ13C (‰ VPDB)δ18O (‰ VPDB)
2614.92Unit 12 2.76–6.042645.80Unit 7 3.29–5.51
2614.92Unit 12 2.75–5.652645.82Unit 7 3.22–5.40
2617.16Unit 12 2.41–5.372645.83Unit 7 3.30–5.68
2618.78Unit 12 3.12–5.142645.86Unit 7C5/C62.99–5.24
2619.99Unit 11 2.88–5.612645.89Unit 7 3.27–5.14
2620.66Unit 11 3.09–6.172645.92Unit 7 3.29–6.91
2621.31Unit 11 3.44–4.472645.95Unit 7 3.35–6.71
2621.91Unit 11 3.17–5.532645.98Unit 7 3.44–5.98
2622.31Unit 11 3.28–6.182646.01Unit 7 3.33–6.80
2622.57Unit 11 3.33–5.942646.04Unit 7 3.32–6.64
2623.06Unit 11 3.21–5.342646.07Unit 7 3.38–6.41
2623.72Unit 11 3.64–5.752646.30Unit 7 3.31–5.80
2624.07Unit 11 3.41–5.772646.65Unit 7 2.99–5.61
2624.28Unit 11 3.47–5.822646.89Unit 7 3.18–6.83
2624.82Unit 11 3.49–6.582647.15Unit 7C4/C53.38–6.29
2625.13Unit 11 3.47–5.452647.45Unit 7 3.20–5.34
2625.54Unit 11 3.70–5.642647.74Unit 7 3.28–6.48
2626.10Unit 11 3.72–5.702648.18Unit 7 3.02–5.12
2626.54Unit 11 3.51–6.932648.58Unit 7 2.90–5.68
2626.82Unit 11 3.55–5.912648.83Unit 7 2.68–5.83
2627.45Unit 11 3.61–6.812649.18Unit 7 2.70–6.43
2627.83Unit 11 3.62–5.532649.53Unit 7 2.64–6.31
2628.19Unit 11 3.71–6.592649.82Unit 7 2.47–6.30
2628.61Unit 11 3.78–6.312650.09Unit 7 2.38–6.29
2628.96Unit 11 3.69–6.752650.38Unit 7 2.35–6.06
2629.60Unit 11 3.40–5.872650.69Unit 7 2.30–6.58
2629.89Unit 11 3.73–6.552650.93Unit 6 2.13–6.11
2630.58Unit 11 3.67–6.852651.35Unit 6 2.11–6.56
2630.86Unit 11 3.46–5.372651.78Unit 6 2.19–6.63
2631.38Unit 11 3.49–5.362652.23Unit 6 2.01–6.38
2631.63Unit 11C7/C83.86–6.182652.50Unit 6 2.10–6.60
2631.78Unit 10 3.86–6.432652.85Unit 6 2.09–6.34
2632.18Unit 10 3.79–5.462653.20Unit 6 2.01–6.56
2632.56Unit 10 4.18–4.902653.41Unit 6 1.99–6.28
2632.94Unit 10 3.92–6.262653.72Unit 6 1.98–6.32
2633.23Unit 10 3.91–6.042653.99Unit 6 2.07–5.93
2633.52Unit 10 3.31–5.892654.31Unit 6 2.02–6.13
2634.00Unit 10 3.42–6.362654.70Unit 6 1.97–6.05
2634.35Unit 10 3.78–6.222654.96Unit 6 1.65–5.55
2634.67Unit 10 3.51–6.512655.20Unit 6 1.89–5.99
2634.88Unit 10 3.66–6.502655.45Unit 6 1.76–5.98
2635.25Unit 10 3.78–6.612655.79Unit 6 1.97–6.62
2635.52Unit 10 3.63–6.592655.95Unit 6 1.88–6.32
2635.82Unit 10 3.62–6.402656.07Unit 6 1.83–5.55
2636.11Unit 10 3.37–6.182656.20Unit 6 1.78–5.40
2636.43Unit 10 3.64–6.702656.33Unit 6 2.23–6.73
2637.05Unit 10 3.36–4.742656.52Unit 5C3/C41.62–4.69
2637.58Unit 10 3.46–5.512656.73Unit 5 2.02–4.90
2638.08Unit 10 3.89–6.302657.02Unit 5 1.68–3.70
2638.37Unit 10 3.72–6.592657.31Unit 5 2.11–5.37
2638.68Unit 10 3.88–5.952657.61Unit 5 2.18–5.84
2638.84Unit 9 3.81–6.302658.08Unit 5 2.21–5.71
2639.14Unit 9 3.71–6.022658.34Unit 5 2.30–5.07
2639.42Unit 9 3.75–6.702658.63Unit 5 2.23–5.00
2639.87Unit 9 3.72–6.292659.02Unit 5 2.34–4.49
2640.00Unit 8 3.76–4.802659.37Unit 5 2.25–4.63
2640.43Unit 8 3.73–4.152659.75Unit 4 2.20–5.29
2640.84Unit 8 3.69–6.162660.09Unit 4 2.02–4.93
2641.35Unit 8 3.71–5.542660.27Unit 4 2.54–5.22
2642.01Unit 8 3.55–6.392660.54Unit 4 2.78–5.09
2642.60Unit 8 3.52–5.592660.81Unit 4 3.02–5.25
2643.03Unit 8 3.45–6.522661.04Unit 4 3.26–5.90
2643.75Unit 8 3.46–5.612661.15Unit 3C2/C33.42–5.51
2644.25Unit 8 3.33–6.252661.45Unit 3 3.47–5.67
2644.74Unit 7 3.39–5.522661.71Unit 3 3.47–5.40
2645.27Unit 7 3.42–5.122662.26Unit 3 3.54–5.71
2645.42Unit 7 3.21–7.102662.51Unit 3 3.62–5.64
2645.45Unit 7 3.25–6.982662.86Unit 2 3.65–5.50
2645.49Unit 7 3.33–6.302663.22Unit 2 3.65–5.77
2645.52Unit 7 3.35–6.482664.04Unit 2 3.43–5.55
2645.55Unit 7 3.36–6.452664.31Unit 2 3.47–5.33
2645.58Unit 7 3.32–6.352664.89Unit 2 3.37–5.82
2645.62Unit 7 3.34–6.532665.23Unit 2 3.49–5.99
2645.62Unit 7C6/C73.45–5.572665.52Unit 2C1/C23.67–6.16
2645.65Unit 7 3.32–6.622665.86Unit 2 3.78–5.15
2645.68Unit 7 3.41–6.172665.98Unit 1 3.66–5.45
2645.71Unit 7 3.37–6.402666.34Unit 1 3.68–5.17
2645.74Unit 7 3.40–6.102667.21Unit 1 3.72–5.88
2645.77Unit 7 3.29–4.652667.83Unit 1 3.67–6.40

[16] The δ18O values range from –7.1 to –3.7‰ in the studied section and exhibit many fluctuations, with amplitudes <~2.0‰ (Figure 2 and 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

[17] 87Sr/86Sr ranges from 0.70752 to 0.70736 in the studied core (Figure 2 and 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 0.70749 in the upper Barremian to 0.70722 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).

Table 4. 87Sr/86Sr and the Numerical Ages From the Studied Core
Depth (mbsf)Lithological Unit87Sr/86Sr2SE (×10–6)Age (Ma)
2614.92Unit 120.70738514123.4+0.7–0.8
2617.16Unit 120.70738616123.4+0.7–0.8
2618.78Unit 120.70739314123.7+0.7–0.7
2620.53Unit 110.70740516124.2+0.7–0.7
2622.22Unit 110.70737316122.9+0.7–0.9
2624.07Unit 110.70738016123.2+0.8–0.9
2626.01Unit 110.70738416123.4+0.7–0.7
2627.59Unit 110.70735617122.0+0.8–1.1
2629.30Unit 110.70740316124.1+0.9–0.9
2631.21Unit 110.70737321122.9+0.9–1.0
2633.06Unit 100.70736816122.6+0.8–1.0
2634.88Unit 100.70737216122.8+0.7–0.8
2636.61Unit 100.70735714122.1+1.0–1.2
2638.20Unit 100.70738814123.5+0.7–0.9
2639.28Unit 90.70738014123.2+0.9–1.0
2640.13Unit 80.70740814124.3+0.8–0.9
2641.91Unit 80.70741118124.4+0.8–0.9
2643.51Unit 80.70741418124.5+0.7–0.7
2645.19Unit 70.70744618125.6+0.7–0.6
2646.89Unit 70.70748816127.5n/a–1.1
2648.58Unit 70.70750717n/an/an/a
2650.51Unit 70.70747813126.8n/a–0.9
2652.23Unit 60.70749816n/an/an/a
2654.23Unit 60.70747017126.4n/a–0.9
2655.79Unit 60.70747320126.5n/a–0.8
2656.73Unit 50.70747114126.4n/a–0.8
2658.63Unit 50.70745213125.8+0.8–0.6
2660.28Unit 40.70743816125.4+0.7–0.7
2661.86Unit 30.70751914n/an/an/a
2663.92Unit 20.70744514125.6+0.7–0.6
2665.68Unit 20.70747414126.6n/a–0.8
2666.27Unit 10.70747114126.4+1.1–0.8
2667.39Unit 10.70746014126.0+0.7–0.6

[18] 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.

4.5 Biostratigraphy

[19] 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 and Table 1).

[20] 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.

[21] 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.

[22] 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.

[23] 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], which 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 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

[24] 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.

[25] 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 < 300 ppm 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.

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.

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); and (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.

[26] 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.

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]; and (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.

[27] Low Fe concentration was used to identify carbonates that are likely to retain the initial isotope composition in many studies (e.g., <3000 ppm [Denison et al., 1994]). The studied samples contain <~9wt % of noncarbonate 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

[28] 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].

[29] 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).
  2. It is known that microbially induced micrite is usually enriched in 13C relative to marine carbonates [e.g., Wu and Chafets, 2002]. 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).

[30] 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 et al., 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).

[31] 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) [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

[32] 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.

[33] 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 applying an 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 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).

[34] 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 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

[35] 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].

[36] 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.

[37] 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 [Pierson et al., 2010, Figure 13]. 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].

[38] 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

[39] 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 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.

[40] The 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]. These studies 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.

[41] 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 those 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.

[42] 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, Figure 3]. Further studies are needed in stratigraphically expanded sections to determine whether the global radiogenic excursion occurred during the OAE1a.

[43] 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

[44] 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].

[45] 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 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].

[46] 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.

[47] 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

[48] 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; Strohmenger et 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).

[49] 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 coprecipitated 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.

[50] 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

[51] 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 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. 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 carbonates 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.
  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

[52] 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, E. Erba, and H. Jenkyns.

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