Fluctuation of organic carbon isotopes of the Lower Cretaceous in coastal southeastern China: Terrestrial response to the Oceanic Anoxic Events (OAE1b)

Palaeogeography, Palaeoclimatology, Palaeoecology 399 (2014) 352–362

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Fluctuation of organic carbon isotopes of the Lower Cretaceous in coastal southeastern China: Terrestrial response to the Oceanic Anoxic Events (OAE1b) Guang Hu a,b,c, Wenxuan Hu a,⁎, Jian Cao a, Suping Yao a, Wenhui Liu b, Zuyi Zhou c a b c

State Key Laboratory for Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing, Jiangsu 210023, China Wuxi Institute of Petroleum Geology, SINOPEC Exploration and Production Research Institute, Wuxi 214151, China State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China

a r t i c l e

i n f o

Article history: Received 7 October 2013 Received in revised form 21 January 2014 Accepted 31 January 2014 Available online 9 February 2014 Keywords: Lower Cretaceous Coastal southeastern China Organic carbon isotope Terrestrial response OAE1b

a b s t r a c t Whether there are terrestrial responses to the Oceanic Anoxic Events (OAEs) remain enigmatic. In this paper, we report a case study in the Lower Cretaceous continental and transitional sequences of coastal southeastern China. It may represent the first late Aptian to Albian terrestrial carbon isotope record from China. Several organic-rich black shale and grayish black mudstone horizons were found to develop widely in the coastal southeastern China during the period from 113 ± 3 Ma to 99 ± 3 Ma, when the Cretaceous OAEs were recorded primarily in marine sequences. Organic carbon isotope (δ13Corg), total organic carbon (TOC), and total nitrogen (TN) were analyzed for the black shales from two sections, i.e., the Shipu transitional and Chong'an continental sections. Five negative δ13Corg excursions can be recognized for both the sections, termed as SI–SV and CI–CV events for the Shipu (S) and Chong'an (C) sections, respectively. Correlations between δ13Corg, TOC and N/C ratio suggest that the SI–SIII and CI–CIII events might have been caused by global factors, i.e., the OAE1b event based on analogs to typical OAE1b reported in the Tethyan and Pacific basins. They archived the terrestrial carbon isotope response to the OAE1b accompanied with a disturbance of carbon isotope in carbon reservoir of the global ocean–atmosphere system under a warm and humid climate. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The Cretaceous Oceanic Anoxic Events (OAEs) are important geological events in the Cretaceous world (Skelton et al., 2003; Jenkyns et al., 2004) and have been intensively studied in many fields, e.g., paleoceanography, paleoclimatology, sedimentology, petroleum geology and paleoecology (Arthur and Schlanger, 1979; Bottjer and Jablonski, 1988; Ross and Scotese, 1988; Bottjer et al., 2001; Huber et al., 2002; Jenkyns et al., 2004). They are commonly characterized by organic enriched black shales, important geochemical anomalies and dramatic turnovers of fossil assemblages (Browning and Watkins, 2008). However, the investigations were focused mostly on marine sequences, and the triggers for OAEs remain controversial, such as monsoonally driven changes in temperature and evaporation/precipitation (Herrle et al., 2003), the large volume of CO2 released by large igneous province (LIP) (Larson, 1991a,b; Gradstein et al., 2012), methane emission from marine hydrates (Wagner et al., 2007), extreme greenhouse conditions

⁎ Corresponding author at: State Key Laboratory for Mineral Deposits Research, School of Earth Sciences and Engineering, Nanjing University, Nanjing, Jiangsu 210023, China. E-mail address: [email protected] (W. Hu). 0031-0182/$ – see front matter © 2014 Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.palaeo.2014.01.027

(Huber et al., 2002; Jenkyns et al., 2004) and sea-level rise (Miller et al., 2005). To solve these issues, samples from different sedimentary settings are helpful besides the samples collected predominantly from marine sequences. The continental sedimentary rocks may act as a candidate because the OAEs should have a terrestrial response via global ocean– atmosphere cycling irrespective of the multiple origins of the events (Wagner et al., 2007; Trabucho et al., 2011). Thus, the works on the continental sedimentary rocks coincided with the OAEs that are significant supplements for the studies in the oceanic realms. They can give external tests for the triggers and patterns of OAEs, and they might address the response mechanism of continental system to oceanic realm. Furthermore, these black shales deposited in continental settings should be important correlation markers for continental and marine sequences, especially for the continental sequences lacking index fossils. However, this issue has not been well discussed. The Lower Cretaceous in the coastal southeastern China has several horizons of organic-rich black shale associations (including shales, mudstones and muddy siltstones) (Fujian Geology and Mineral Resources Bureau, 1985, 1997; Zhejiang Geology and Mineral Resources Bureau, 1989, 1996), which have been reported to be influenced by marine transgressions and were deposited during 113 ± 3 Ma to 99 ± 2 Ma (i.e., from the late Aptian to the early Albian interval) (Xie et al.,

G. Hu et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 399 (2014) 352–362

2010; Hu et al., 2012a,b). In addition, the Lower Cretaceous black shale associations in southeastern China are synchronous to those deposited in northeastern China which are also be influenced by transgressions (Sha, 2007; Sha et al., 2008; Sha and Spencer, 2012), likely implying a supraregional geological event that took place in eastern China. During the Aptian–Albian boundary interval, black shales were widespread in marine basins (Coccioni, 1996; Bralower et al., 1999; Jenkyns and Wilson, 1999; Erbacher et al., 2001; Hofmann et al., 2001; Herrle et al., 2004; Luciani et al., 2004; Tsikos et al., 2004; Coccioni et al., 2006; Arnaboldi and Meyers, 2007; Browning and Watkins, 2008; Hofmann et al., 2008; Trabucho et al., 2011). Four organic-rich layers deposited during the interval (i.e., the Jacob, Kilian, Paquier and Leenhardt) in the Vocontian Basin in southeastern France have been intensively studied (Bréhéret, 1988; Tribovillard and Gorin, 1991; Bréhéret, 1997; Herrle et al., 2004, 2010) and the “Ocean Anoxic Event 1b” (OAE1b) has been applied to the all four black shales or to a portion (Gradstein et al., 2012). Therefore, the Lower Cretaceous terrestrial shale associations in coastal southeastern China are likely good materials for studying terrestrial response to OAEs, in particular OAE1b. OAEs are commonly associated with organic-rich sediments with total organic carbon (TOC) contents generally N 1.0% and globally synchronous perturbations in the carbon isotope records (Jenkyns and Wilson, 1999; Herrle et al., 2004; Trabucho et al., 2011). Therefore, to determine whether there is terrestrial response to the OAEs in the Early Cretaceous of coastal southeastern China, we conduct analyses of TOC, total nitrogen (TN) and organic carbon isotopic composition (δ13Corg) for black shale associations from two Early Cretaceous sections in the transitional and lacustrine depositional settings, respectively. Integrated with necessary organic petrological and palynological studies, the fluctuations of organic carbon isotopic composition of the black shale associations in the two sections are discussed. Then, we attempt to correlate these fluctuations to OAEs in a supraregional to global scale. It may represent the first OAE1b terrestrial records from China and can make a contribution to the study of the Early Cretaceous stratigraphic evolution in southeastern China.

117°

2. Geological setting The coastal southeastern China commonly refers to the area bordered by the famous Jiangshan–Shaoxing and Ganjiang faults as the northern and western boundaries, respectively (Fig. 1). The NNEtrending Zhenghe–Dapu fault crosses the central portion of the area. Tens of Early Cretaceous NE-trending rift basins are present throughout the area; they were believed to be generally formed due to the lithospheric thinning, and extensions stimulated by the westward subduction of the Paleo-Pacific plate (Wang and Zhou, 2002; Shu et al., 2009). These basins were greatly reshaped by later tectonic activities (Shu et al., 2009). The studied Shipu (transitional/tidal) and Chong'an (continental/ lacustrine) sections are located in the Shipu town of the Zhejiang Province and the Chong'an town of the Fujian Province, respectively (Hu et al., 2012b) (Fig. 1). In the Shipu section, the sedimentary rocks are developed mainly in the middle (30–90 m) and upper (90–120 m) portions, while the lower portion (0–30 m) consists primarily of volcanic breccia that are interbedded with tuffaceous sandstones (Fig. 2). The middle portion consists predominantly of silicified tuffaceous sandstones and siltstones, black silicified shales, and limestones. The upper portion consists mainly of silicified tuffaceous sandstones and siltstones, limestones, and grayish black mudstones. The marine limestones are primarily stromatolite algal-bonded, oolitic, bioclastic, and micritic limestones, and marlstone (Xu and Zheng, 1989; Wang et al., 2012). In the Chong'an section, as shown in Fig. 3, the lower part of the section (0–50 m) consists primarily of tuffaceous conglomerates and sandstones. The middle portion (50–250 m) consists mainly of siltstones that are interbedded with black laminated shales. The siltstones dominate the upper portion of the section (250–525 m) with massive grayish black mudstones that have alternating thin layers of yellowishgreen siltstones. In contrast to the tidal marine limestones and mudstones in the Shipu section, the black shales and mudstones in the Chong'an section were formed in a lacustrine setting influenced by marine transgressions (Xie et al., 2010; Hu et al, 2012b).

119°

121°

123°

30°

30°

115°

353

Ningbo

ce

g jian

Nanchang

Shipu

one

Lishui Xiaoling Basin

s

ault Z

n ai

nt

ou

t Z on

2

2

Fa

Chong’an

bu

Present-day Early Cretaceous remained basins

117°

e

Marine or transitional zone

ro vi nc

Da Sanming

Inferred marine or transitional zone with uncertain boundary

nP

Fu jia

Zh

en

gh

e

Julan

Section location

ul

W

Legend

Lacustrine intertongued with transgression zone Terrestrial brackish or saline lacustrine zone

Fuzhou 119°

121°

Fig. 1. Schematic map showing the location of the Shipu and Chong'an sections in coastal southeast China. Modified after Hu et al., 2012b.

123°

26°

J iang

uy

e

iM

xi Pr ovinc e

28°

ang F Ganji

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26°

P

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Yong’an

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in rov

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Depth (m)

Lithological column

Samples

Total Organic Carbon and Organic Carbon isotope

Q δ 13Corg (‰) 120

99±3Ma

-28 bsp-80

-20

-16

-12

SV

bsp-70 bsp-68

110

-24

SIV

bsp-65 bsp-61 bsp-57

100

100±2Ma

90 bsp-53 bsp-48

80 109±2Mabsp-43 bsp-42 bsp-41 bsp-39 bsp-36 bsp-30 bsp-29 bsp-28 bsp-26 bsp-24 bsp-23 bsp-21

70

60

SIII

SII

bsp-18

50 bsp-16 bsp-14 bsp-12 bsp-10 bsp-08

40

SI

113±3Ma bsp-04 bsp-01

30

0

0.5

1

TOC (%)

volcanic breccia siltstone

tuffaceous sandstone mudstone and shale

tuffaceous siltstone

tuffaceous mudstone

limestone

sandstone siltstone with stromatolite

Fig. 2. Generalized stratigraphic column of the Shipu section. We present sample localities, total organic carbon (TOC) contents, organic carbon isotope (δ13Corg) values, four zircon U–Pb ages reported by Hu et al. (2012a), and five δ13Corg negative excursion events SI–SV.

Our previous chronometric studies suggested that the sequences in the two sections are comparable, younger than 113 ± 3 Ma and older than 99 ± 2 Ma (Hu et al., 2011a, 2012a). Therefore, the black shales and mudstones in the two sections were deposited during the late Aptian to Albian period and can be well correlated.

3. Sampling and methods A total of 28 samples from the Shipu section and 44 samples from the Chong'an section were analyzed, respectively, including 17 muddy siltstones, 20 mudstones, and 35 shales in total (Figs. 2 and 3; Table 1).

G. Hu et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 399 (2014) 352–362

Depth (m)

Lithological column

355

Total organic carbon and Organic Carbon isotope

Samples

δ 13Corg (‰) -29 xwy-76 xwy-75 xwy-74 xwy-72 xwy-71 xwy-70 xwy-69 xwy-59 xwy-57 xwy-56 xwy-55 xwy-52 xwy-51

500

-28

-27

-26

-25

-24

CV CIV

400

300 xwy-50 xwy-49 xwy-47 xwy-46 xwy-45 xwy-42 xwy-40 xwy-39 xwy-38 xwy-36 xwy-35 xwy-34

200

CIII

CII

xwy-32 xwy-30 xwy-27 xwy-26 xwy-24 xwy-23a xwy-23 xwy-25a

CI

xwy-20a

xwy-20 xwy-16 xwy-10 xwy-09 xwy-08 xwy-07 xwy-05 xwy-03 xwy-02 xwy-01

100

0

0.5

1.5 1 TOC (%)

2

116±2M a

0 volcanic breccia

tuffaceous sandstone

sandstone

siltstone

mudstone and shale

Fig. 3. Generalized stratigraphic column of the Chong'an section. We present sample localities, total organic carbon (TOC) contents, organic carbon isotope (δ13Corg) values, zircon U–Pb age at the base of the section reported by Hu et al. (2012a), and five δ13Corg negative excursion events CI–CV.

We perform a relatively dense sampling strategy in shales and mudstones, with an approximately 0.5 m interval or less. This aims to obtain a relatively high-resolution carbon isotopic record to characterize the carbon isotopic fluctuation in detail. In contrast, the muddy siltstone

samples are comparatively sparse with an approximately 2 m or 5 m interval in the siltstone section. The TOC and the TN analyses were performed via combustion with an Elementar® Vario MACRO CHNS elemental analyzer at the State

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Table 1 Analytical results of organic carbon isotope (δ13Corg), total organic carbon (TOC), and N/C ratio. Sample

δ13Corg (‰)

TOC (%)

N/C

Lithology

xwy-01 xwy-02 xwy-03 xwy-05 xwy-07 xwy-08 xwy-09 xwy-10 xwy-16 xwy-20 xwy-20a xwy-23 xwy-23a xwy-24 xwy-25a xwy-26 xwy-27 xwy-30 xwy-32 xwy-34 xwy-35 xwy-36 xwy-38 xwy-39 xwy-40 xwy-42 xwy-45 xwy-46 xwy-47 xwy-49 xwy-50 xwy-51 xwy-52 xwy-55 xwy-56 xwy-57 xwy-59 xwy-69 xwy-70 xwy-71 xwy-72 xwy-74 xwy-75 xwy-76 bsp-01 bsp-04 bsp-08 bsp-10 bsp-12 bsp-14 bsp-16 bsp-18 bsp-21 bsp-23 bsp-24 bsp-26 bsp-28 bsp-29 bsp-30 bsp-36 bsp-39 bsp-41 bsp-42 bsp-43 bsp-48 bsp-53 bsp-57 bsp-61 bsp-65 bsp-68 bsp-70 bsp-80

−27.2 −27.6 −27.4 −27.6 −27.2 −26.8 −26.9 −24.7 −25.8 −25.5 −27.6 −28.0 −27.6 −27.4 −27.4 −27.5 −27.5 −25.3 −24.4 −26.9 −27.9 −26.9 −27.0 −27.4 −27.5 −27.6 −27.8 −27.5 −27.0 −25.5 −25.1 −23.7 −25.9 −28.3 −27.8 −26.6 −25.3 −24.9 −26.7 −25.8 −27.0 −25.5 −27.2 −24.6 −21.5 −21.3 −18.5 −26.7 −25.9 −20.2 −25.8 −19.8 −24.9 −22.7 −27.8 −16.0 −18.6 −15.6 −21.8 −22.8 −22.6 −21.2 −20.3 −20.2 −17.8 −17.1 −18.4 −22.9 −17.4 −17.1 −23.8 −16.7

0.55 0.81 0.62 0.78 1.46 1.14 0.83 0.14 0.06 0.16 1.00 1.18 1.58 1.58 1.13 0.89 0.96 0.17 0.16 1.70 1.17 0.91 0.72 1.36 0.55 1.21 0.60 0.09 0.60 0.63 0.31 0.18 0.13 1.39 1.07 0.68 0.58 1.18 1.46 0.16 1.00 0.16 0.62 0.62 0.79 0.26 0.14 0.37 0.23 0.13 0.38 0.56 0.24 0.18 0.29 0.29 0.19 0.33 0.53 0.36 0.12 0.32 0.15 0.54 0.13 0.14 0.12 0.25 0.25 0.35 0.12 0.60

1.04 0.71 0.74 0.65 0.36 0.44 0.61 2.10 4.81 3.13 0.55 0.42 0.31 0.31 0.46 0.54 0.59 2.84 2.60 0.25 0.42 0.48 0.69 0.34 0.47 0.43 0.83 2.58 0.76 0.80 1.82 2.99 – – – 0.34 0.69 0.25 0.31 1.42 0.42 1.44 0.73 0.47 1.32 1.57 4.14 2.01 1.97 6.70 0.56 0.84 2.47 2.67 1.39 1.39 5.41 3.99 0.78 0.65 5.07 2.92 5.97 0.81 4.76 10.21 3.18 3.62 2.46 4.36 2.64 2.08

Grayish green siltstone Grayish green siltstone Grayish black mudstone Grayish black mudstone Black shale Black shale Black shale Grayish green siltstone Grayish green siltstone Grayish yellow siltstone Black shale Black shale Black shale Black shale Black shale Black shale Black shale Grayish green mudstone Grayish green mudstone Black shale Black shale Black shale Black shale Black shale Black shale Black shale Black shale Grayish yellow siltstone Black shale Black shale Black shale Grayish green mudstone Grayish green mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish green mudstone Grayish green mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish green mudstone Gray siltstone Gray siltstone Black shale Black shale Black shale Black shale Black shale Black shale Black shale Black shale Black shale Black shale Gray siltstone Gray siltstone Gray siltstone Black shale Black shale Black shale Black shale Gray siltstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Grayish black mudstone Gray siltstone Grayish black mudstone

Key Laboratory for Mineral Deposits Research of Nanjing University. Prior to analysis, the samples were treated with 1 N HCl for 24 h to remove inorganic carbon. Carbon isotopic analyses are focused on organic matter irrespective of lithology in order that analytical results can be compared under the same geological and geochemical conditions. The samples for the δ13Corg analyses were from fresh shale associations. The powered sample was treated with a hot 2 N solution of HCl for 2 h to remove carbonate minerals. These acid-processed samples were baked in an oven at 650 °C for 5 h in a vacuum tube along with a quantity of CuO to convert the organic carbon into CO2 gas. After the purification of the CO2 gas on a cryogenic vacuum line, the carbon isotopic analyses were performed with Finnigan MAT 252 mass spectrometers at the State Key Laboratory for Mineral Deposits Research of Nanjing University. The mass spectrometers were calibrated with limestone powders from the U.S. National Bureau of Standards. The results are expressed in the standard δ notation, with respect to the PDB standard, and with an instrumental standard deviation of analyses of ±0.02‰. 4. Results 4.1. Petrography Field reconnaissance and thin section observations demonstrate that the black shale associations from the Shipu section are characterized by alternating layers of grayish yellow tuffaceous siltstone and black organic-rich laminas and have unambiguous but not continuous lamellations (Fig. 4a and b). The kerogen is characterized by type II as the average abundances of the sapropelinite and vitrinite compositions in the maceral are approximately 61.4% and 31%, respectively (Hu et al., 2011b). The organic matter has a low thermal maturation in general (Hu et al., 2011b). The palynological analyses indicate that the coniferophyte pollens, represented by Podocarpidites and Pinuspollenites, account for 63%, whereas the evergreen latifoliate pollens (Celtispollenites) and herbaceous pollens (Graminidites) account for 22% and 15%, respectively (Liu and Hu, in press). This implies that the coniferous plants dominate the canopy at that time. In the Chong'an section, the black shale associations are composed of alternating grayish clay and black organic-rich laminas with numerous ambiguous and continuous lamellations (Fig. 4c). This is somewhat different from the Shipu section, likely indicative of the variation of sedimentary environment from transitional tidal to continental lacustrine. Microscopic observation on rock and acid-macerated thin sections implies that the organic matter is dominated by terrestrial plants (Fig. 4d and e). In addition, some aquatic organic matters, typically represented by red algae (Fig. 4d and f), can be found in the upper portion from the depth 400 m related to marine transgression (Xie et al., 2010; Hu et al., 2012b). According to Hu et al. (2011b), the total concentration of inertinites and vitrinites in the maceral is up to 81.4%, implying that the kerogen is predominantly type III, also different from the Shipu section. This is consistent with the difference of sedimentary environment, i.e., tidal (more type I–II marine planktons) vs. lacustrine (more type III higher plants). The organic matter is highly to over mature in general (Hu et al., 2011b). The palynological study demonstrates that up to 91.8% of the pollens are Classopollis, Pinuspollenites and Tracheid, implying that the shrubs are the major type of vegetation (Liu and Hu, in press). This is also different from the Shipu section, further indicative of the variation in sedimentary environment. 4.2. General characteristics of geochemical analyses Figs. 2, 3, and Table 1 present the analytical results of TOC, TN and δ13Corg. TN results were converted to the ratio of total nitrogen to total organic carbon (N/C) (Table 1) as commonly used (Meyers, 1994). It is showed that the TOC content, N/C ratio, and δ13Corg value of the two sections vary widely. In the Shipu section, they range in 0.12–0.79%,

G. Hu et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 399 (2014) 352–362

357

Fig. 4. Photographs showing petrological characteristics of the black shales/mudstones from the Shipu and Chong'an sections. (a) Outcrop black shale from the Shipu section, the discontinuous lamellations are unambiguous; (b) outcrop black shale from the Chong'an section, the continuous lamellations are unambiguous; (c) organic-rich lamina alternated with clay mineral lamina, rock thin section, black shale from the Shipu section, transmitted light; (d) organism detritus in organic-rich lamina, rock thin section, transmitted light; (e) terrestrial plant detritus from the Chong'an section, acid maceration thin section, transmitted light; (f) algae fragment from the Chong'an section, acid maceration thin section, transmitted light.

0.56–10.21, and − 27.8‰ to − 15.6‰, respectively. In the Chong'an section, they vary between 0.06–1.70%, 0.25–4.81, and − 28.3‰ to − 23.7‰, respectively. In general, the average TOC content of the Chong'an section (0.79%) is higher than that of the Shipu section (0.30%). This might be caused by their different sedimentary settings and preservation. On the contrary, the averages of the δ13Corg and N/C ratio of the Shipu section (−20.8‰ and 3.0, respectively) are higher than those of the Chong'an section (− 26.8‰ and 1.0, respectively) (Table 1). The difference of δ13Corg and N/C ratio might be controlled by the organic matter composition. The type II organic matter in the Shipu section has more concentration of marine phytoplankton algae than the type III organic matter in the Chong'an section. The submerged plants and algae uptake carbon 13 C to a greater degree compared with the as aqueous HCO− 3 , and enrich emerged plants, which have δ13Corg values that are similar to the terrestrial plants (Smith and Epstein, 1971; Aravena, 1992). In contrast, the marine phytoplankton algae have δ13C values ranging from − 10‰ to −16‰, which are higher than those of the terrestrial plants (Park and Epstein, 1961). Thus the δ13Corg values of the Shipu section are greater than those of the Chong'an section. Similarly, the average N/C ratios of the Shipu section are also greater than those of the Chong'an section

as phytoplankton algae commonly have more nitrogen components than higher plants (Kanasassanen and Jaakkola, 1985; Meyers, 1994). 4.3. Organic carbon isotope fluctuations For the OAE1b during the late Aptian to early Albian, they are generally characterized by negative δ13C excursions in carbonate or in whole rocks, such as in the Vocontian Basin (Herrle et al., 2004), the Mazagan Plateau (Herrle et al., 2004), the Northeast Atlantic Ocean (Erbacher et al., 2001), southern England (Gröcke et al., 1999) and Portugal (Heimhofer et al., 2003). Similarly, we examine such negative δ13C excursions in this study, and successfully recognize five such excursions in both the sections (Figs. 2 and 3). We term them as SI–SV and CI–CV in ascending order for the Shipu (S) and Chong'an (C) sections, respectively (Figs. 2 and 3). The SI event, a negative excursion with an average δ13Corg value of − 26.1‰, occurs at the 44.4–44.7 m depth in the Shipu section above the tuffaceous siltstone bed that was dated at 113 ± 3 Ma (Hu et al., 2012a) (Fig. 2). The SII event, a strongest negative excursion with a peak δ13Corg value of −27.8‰, is an approximately 30 cm bed of black shale that is located at the 60–62.4 m depth (Fig. 2). The SIII event, a

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relatively weak negative δ13Corg excursion, at the 70–70.1 m depth, and is a ten centimeter bed of black shales interbedded with lamellar limestones. The δ13Corg values of the two samples in this segment are −22.6‰ and −21.2‰ (Fig. 2). Above this segment, the tuffaceous sandstone interbed was dated at 109 ± 2 Ma (Hu et al., 2012a). The SIV event, a weak negative excursion, has a δ13Corg value of − 25.8‰ and occurs at the depth of 105 m (Fig. 2). The SV event, a weak negative excursion with a δ13Corg value of − 23.8‰ occurs at the depth of 113 m (Fig. 2). In the Chong'an section, the five negative δ13Corg excursions (CI–CV) have the average values of −27.0‰, −27.9‰, −27.6‰, −28.0‰ and − 27.0‰, respectively. These events are similar to the above SI–SV events in their trends and in their visual pattern. The only difference is their stratigraphic thickness (Figs. 2 and 3). To summarize, the two sections have a similar stratigraphic carbon isotopic tendency, and the five negative δ13Corg excursion events appear to act as correlation markers in the two sections (Figs. 2 and 3). This suggests that there might be similar driving factors in these two sections.

5. Discussion 5.1. Origin of organic carbon isotope fluctuations 5.1.1. Correlation between TOC and N/C For discussing the origin of the organic carbon isotope fluctuations (excursions), correlations between TOC with N/C and δ13Corg were studied except segments SIV and SV, because the database (only two samples) is too small (Fig. 2). Fig. 5a, c and e demonstrates that there are generally two correlation patterns between TOC and N/C, as represented by CI–CIII, and CIV–CV and SI–SIII, respectively. For the CI–CIII events in the Chong'an section, an inversely linear correlation between TOC and N/C ratios is distinct with the correlation coefficient (R2) being 0.864 (Fig. 5a). A majority of the samples are located within the dotted lines, which represent the 95% confidence bands (Fig. 5a). The t-test indicates that the linear correlation is significant for the population at a 95% confidence level (t = 4.203 N t0.025 = 2.306). According to Meyers (1994), the N/C ratio

a 1.6

b

1.75

CI - CIII

XWY-24 XWY-23a

CI - CIII

XWY-24

1.50

XWY-23a

1.4 1.25

XWY-42 XWY-23 XWY-35

1.0 XWY-20a XWY-36

0.8

TOC (%)

TOC (%)

XWY-39

1.2

XWY-35 XWY-23

1.00

XWY-20a XWY-36

0.75 0.50

0.6 XWY-40

0.30

0.35

0.40

0.45

0.50

XWY-40

0.25 -28.5

0.4 0.55

-28.0

-27.5

c

1.6

-26.5

d

1.8

CIV - CV

XWY-70

CIV - CV

1.6 1.4

1.2

1.2

1.0

TOC (%)

TOC (%)

-27.0

δ 13 Corg (‰)

N/C

1.4

XWY-39

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Fig. 5. Cross plots of TOC vs. N/C and TOC vs. δ13Corg of black shales and grayish black mudstones from the five negative δ13Corg excursion events in Shipu and Chong'an sections. Samples in plots (a) and (b) are from the CI–CIII events; samples in (c) and (d) are from the CIV and CV events; samples in (e) and (f) are from the SI and SIII events. The red lines in figures (a), (d), and (f) are linear fitting lines, while those in the (c) and (e) figures are reciprocal fitting lines. The dotted lines in all figures are confidence bands at 95%.

G. Hu et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 399 (2014) 352–362

of algae is greater than that of terrestrial organic matter. Elevated N/C ratios were commonly used to identify a high input of algal organic matter (Kanasassanen and Jaakkola, 1985). On the contrary, a decrease of N/C ratio was interpreted to indicate a high input of terrestrial organic matter (Guilizzoni et al., 1996). Thus, the robust inversely covariation of TOC and N/C ratio suggests that the organic carbon content of the black shales from the CI–CIII segments is controlled primarily by the input of terrestrial detritus. In addition, the low N/C ratios within a narrow range (from 0.30 to 0.55) of the black shales from the CI–CIII segments suggest that the organic matter may be dominated by terrestrial detritus. This is consistent with kerogen type (Hu et al., 2011b), petrographic and general geochemical results discussed above. In contrast, the fitting equations that link the TOC and N/C ratios of the black shales/mudstones from the CIV–CV and SI–SIII segments are reciprocal functions in which the correlation coefficients (R2) are 0.978 and 0.963, respectively (Fig. 5c and e). The t-test indicates that the reciprocal correlations are significant for their population at a 95% confidence level. The t-test values are 11.581 and 13.716, respectively, and both of them are above the critical points (2.571 and 2.228, respectively). The nonlinear correlations between the TOC and N/C and elevated N/C ratios (Table 1) imply that the organic matters are composed of a mixture of terrestrial plants and algae. In addition, the TOC values of CIV–CV and SI–SIII segments also increase along with the N/C ratios decreasing. This is similar to the characteristics of the CI–CIII segments, indicating that the organic carbon concentration of the black shales/mudstones in the CIV–CV and SI–SIII segments is also controlled mainly by terrestrial detritus input because a decreasing N/C ratio was interpreted to indicate a high input of terrestrial organic matter (Guilizzoni et al., 1996). 5.1.2. Correlation between δ13Corg and TOC As discussed above, the organic matter precursors from CI–CIII segments are characterized by higher plants, whereas the CIV–CV and SI–SIII segments have mixed origins. In theory, if the organic matter has a mixed origin, the δ13Corg value should display a decrease with increasing TOC value when the organic carbon concentration is controlled by terrestrial detritus input, because the δ13Corg value of terrestrial plant is lower than that of submerged plants, algae and phytoplankton (Park and Epstein, 1961; Smith and Epstein, 1971; Aravena, 1992). However, this trend between δ13Corg and TOC is only observed for the CIV–CV events, whereas not in the SI–SIII events as predicted (Fig. 5d). The linear correlation coefficient (R2) for the CIV–CV events is as high as 0.916. Therefore, the terrestrial detritus input may be the major factor for the negative δ13Corg excursion in the CIV–CV segments. In contrast, the negative linear correlation between the δ13Corg value and the TOC is faint for the SI–SIII segments (Fig. 5f), although the organic matter is also from a mixed origin like that in CIV–CV segments. However, the correlation coefficient (R2) for the SI–SIII events is only 0.300, and the t-test also indicates that the linear correlation may be not significant for the population at a 95% confidence level (t = 1.9247 b t0.025 = 2.228). Therefore, the factors controlling the negative δ 13Corg excursions in the SI–SIII segments should be different from that in the CIV and CV segments. Similar nonlinear correlation pattern between δ13Corg and TOC values to that of the CIV–CV segments was observed in the CI–CIII negative δ13Corg excursion events, in which the δ13Corg values have no linear correlation with TOC (Fig. 5b). As discussed above in our petrographic and geochemical studies, the organic matter precursors in the CI–CIII segments and neighboring beds originate primarily from terrestrial plants. If the organic matter is predominantly composed of higher plants, the δ13Corg value should have little connection with the TOC value as the organic matter composition and type is relatively stable irrespective of the variation in TOC value (Park and Epstein, 1961; Smith and Epstein, 1971). This is the reason why there is little linear correlation between δ13Corg and TOC values in the CI–CIII and SI–SIII segments. Therefore, the decrease in δ13Corg values of terrestrial detritus

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in the CI–CIII and SI–SIII segments should be responsible for the six negative δ13Corg excursions. 5.1.3. Origin of SI–SIII and CI–CIII δ13Corg excursions Based on the above discussion, it is indicated that the SI–SIII and CI– CIII δ13Corg excursions are most likely caused by the decrease in δ13Corg values of terrestrial detritus. Hayes (1993, 2001) summarized that the carbon isotope composition of organic matter mainly depends on the 13C-content of the carbon source and isotope effects associated with biosynthesis and metabolism. If the decrease in δ13Corg value of terrestrial detritus was caused by 13 C-depleted carbon source, then it implies that there were decreases in δ13C of atmospheric carbon dioxide because plants mainly uptake carbon from atmospheric carbon dioxide. Therefore, these carbon isotopic signals should be archived in continental and marine sedimentary rocks via global ocean–atmosphere cycling (Wagner et al., 2007; Trabucho et al., 2011). With respect to isotope effects associated with biosynthesis, there are two main isotope-discriminating steps during biological carbon fixation in plant, i.e. the uptake and intracellular diffusion of CO2 and the biosynthesis of cellular components (Park and Epstein, 1960). The two-step model of carbon fixation suggests that isotope fractionation is dependent on the partial pressure of CO2 in system, i.e. Pi and Pa, the pCO2 in the leaf and the atmosphere, respectively (Francey and Farquhar, 1982; Farquhar, 1983). Francey and Farquhar (1982) and Farquhar (1983) determined that the δ13C values of plants have a negative linear correlation with the Pi/Pa ratio. An increase in the atmospheric CO2 level will lead to an increase in Pi/Pa ratio (Keeling et al., 1979; Farquhar, 1983), and thus the δ13Corg values will become more negative. In addition, certain factors associated with metabolism have also been demonstrated to cause decreasing δ13Corg values in terrestrial plants, e.g., decreasing enzymatic activity and stomatal amplification (Tieszen, 1991). The former may be induced by low temperature, declining irradiance, and low nutrient supply, while the latter may be caused by less salt-stress or water-stress (Tieszen, 1991). In fact, the late Aptian–early Albian period was interpreted to be warm with widespread humidity in general (Erbacher et al., 2001; Herrle et al, 2003; Browning and Watkins, 2008; Wagner et al., 2008). The two sections in this study are located in relative low latitude, being close to the Pacific (Fig. 1). This relieves the salt-stress and water-stress and increases the enzymatic activity, resulting in the more negative δ13Corg values of the terrestrial plants. Thus, the factors influencing the δ13Corg excursions here, such as 13C-depleted atmospheric carbon dioxide, an increasing level of pCO2 in the atmosphere and/or a warm and humid climate, are all global in scale. This provides good foundation for correlation with global events. 5.2. Possible correlation with OAEs As discussed above, the SI–SIII and CI–CIII events might be triggered by global factors. Therefore, these events can be used for global correlations. Previous studies suggest that the black shales/mudstones in the Shipu and Chong'an sections were formed contemporaneously (113 ± 3 Ma to 109 ± 2 Ma; Hu et al., 2011a, 2012a). Our study here indicates that the organic carbon isotopes of the two sections have similar fluctuation tendency (Figs. 2 and 3). Therefore the shales/mudstones in the two sections are comparable. Furthermore, during the SI–SIII and CI–CIII deposited interval (from 113 ± 3 Ma to 109 ± 2 Ma) (Hu et al, 2012a), OAE1b was widespread in marine basins (Bréhéret, 1988; Tribovillard and Gorin, 1991; Bréhéret, 1997; Herrle et al., 2004, 2010; Gradstein et al., 2012). Thus, we conducted a correlation of our carbon isotopic results with two representative cases with good analytical results, including the Tethyan and Pacific basins. As shown in Fig. 6, it is demonstrated that the six negative δ13Corg excursion events in the Shipu and Chong'an

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Fig. 6. Carbon isotopic correlations of the late Aptian to early Albian Oceanic Anoxic Events (OAE1b).

sections are well correlated to the OAE1b, likely indicating that the global OAEs have terrestrial responses. The carbon isotopic records in the SI and CI events are the first negative δ13Corg excursions above the top of the positive δ13Corg excursion (Fig. 6). This is comparable with the Kilian characteristics, which is a short negative excursion after the high Jacob δ13C stage (Scholle and Arthur, 1980; Bralower et al., 1999; Jenkyns and Wilson, 1999; Heimhofer et al., 2003; Herrle et al., 2004). In addition, the age of the SI and CI events is approximately younger than 113 ± 3 Ma according to zircon U–Pb dating analyses (Hu et al., 2012a), which is generally synchronous with the Kilian levels. Above the Kilian level is the Paquier level (Grötsch et al., 1998; Herrle et al., 2004; Friedrich et al., 2005). High resolution carbon isotope records from the Vocontian Basin in SE France demonstrate that the δ13C decrease stepwisely after a short recovery that is just above the Kilian level, until a small maximum that is superimposed by a short and maximum negative excursion (Herrle et al., 2004). The short and maximum negative excursion coincides with the Paquier level (Herrle et al., 2004; Friedrich et al., 2005). The variation of carbonate carbon isotope from the Pacific Ocean is almost the same as that from the Vocontian Basin (Jenkyns and Wilson, 1999). This similar pattern was observed in our SII and CII events, where the maximum negative δ13Corg values were recorded by the samples bsp-24 and xwy-23 from the Shipu and Chong'an sections, respectively (Fig. 6; Table 1). The negative SIII and CIII excursion events that occur at a depth of approximately 70 m and 250 m in the Shipu and Chong'an sections respectively have been determined to be older than 109 ± 2 Ma (Hu et al., 2012a). They are generally synchronous with the Leenhardt level and show similar carbon isotope variation trends (Fig. 6). Therefore, the six well correlated negative δ13Corg excursions in the Shipu and Chong'an sections are comparable with the OAE1b that was reported mostly in oceanic realms. Thus, our study might provide a new terrestrial supplement and these excursions should also be used as markers for stratigraphic correlations as they are well correlated and controlled by supraregional to global factors. Similar varying trends of carbon isotopes can be observed between different sedimentary environments including marine, transitional and lacustrine. This high consistency may be indicative of the global nature of the OAE1b, and imply that the atmospheric carbon reservoir was

affected under this global impact, as the single factor influencing the sotopic composition of both terrestrial organic carbon and marine carbonate carbon is the isotopic change in the global CO2 reservoir of the ocean–atmosphere system (Hasegawa, 1997, 2003). It may be predicted that many contentious issues of OAEs can be better understood as more and new continental samples can be collected. 6. Conclusions (1) Five negative δ13Corg excursion events, i.e., the SI–SV and CI–CV events, can be recognized from both the Shipu (S) and Chong'an (C) sections, respectively. (2) Of the ten excursions, the SI–SIII and CI–CIII events are most likely caused by global factors, likely an increase in the atmospheric CO2 level under a warm and humid climate, showing a potential for global correlation. (3) The SI–SIII and CI–CIII events are terrestrial responses to the OAE1b, as not only similar to OAE1b in stratigraphic carbon isotopic variation but also roughly synchronous to the OAE1b. They might provide new and good opportunities to better understand many contentious issues in the study of OAEs. Acknowledgments We thank Prof. David J. Bottjer and two anonymous reviewers for constructive comments that help to improve the article. We are grateful to Profs. Wenlan Zhang and Mingyuan Lai, and Drs. Hai Ding and Kun Jiao (all from School of Earth Sciences and Engineering, Nanjing University) for the help with experiments. M.S. Qi Chen and Xiaomin Xie (Wuxi Research Institute of Petroleum Geology, SINOPEC) are thanked for assistance with field work. This study was jointly funded by the Major State Basic Research Development Program (973 Project, Grant No. 2012CB214803 and No. 2012CB214801), and the National Natural Science Foundation of China (Grant No. 41322017 and Grant No. 41072090). References Aravena, R., 1992. Carbon isotope composition of lake sediments in relation to lake productivity and radiocarbon dating. Quat. Res. 37, 333–345.

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