Marine to brackish depositional environments of the Jurassic–Cretaceous Suowa Formation, Qiangtang Basin (Tibet), China

Palaeogeography, Palaeoclimatology, Palaeoecology 473 (2017) 41–56

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Marine to brackish depositional environments of the Jurassic–Cretaceous Suowa Formation, Qiangtang Basin (Tibet), China Ruofei Yang a, Jian Cao a,⁎, Guang Hu a,b, Lizeng Bian a, Kai Hu a, Xiugen Fu c a b c

State Key Laboratory for Mineral Deposits Research, Department of Earth Sciences, Nanjing University, Nanjing, Jiangsu 210023, China School of Geoscience and Technology, Southwest Petroleum University, Chengdu, Sichuan 610500, China Chengdu Institute of Geology and Mineral Resources, Chengdu, Sichuan 610081, China

a r t i c l e

i n f o

Article history: Received 29 November 2016 Received in revised form 21 February 2017 Accepted 21 February 2017 Available online 24 February 2017 Keywords: Black shales Organic matter Tethys Marine–continental transition Transgression

a b s t r a c t The Late Jurassic to Early Cretaceous depositional environments of the Qiangtang Basin in Tibet have the potential to provide significant insight into mechanisms of black shale deposition, organic matter accumulation, and the timing of closure of the Mesotethys Ocean. However, the depositional setting has not been well constrained. Here we apply multiple geochemical proxies and petrologic analyses to representative samples of black shale, marl, and micrite collected from a section in the region. These indicate a transitional marine–continental environment, with brackish to saline water. Redox conditions were weakly oxic to suboxic. Environments represented by the studied section varied over time. The lowermost marls were deposited in a low-salinity, weakly oxic shore to shallow lake environment, under a warm and humid climate regime. The micrites in the middle part of the section were deposited in a lagoonal environment, with intense evaporation, high salinity, and water column stratification. The uppermost shales were deposited in a reducing, semi-enclosed lagoon environment during a marine transgression. These results suggest that the Late Jurassic–Early Cretaceous succession was deposited in a tidal-flat or lagoonal environment. © 2017 Elsevier B.V. All rights reserved.

1. Introduction The Qiangtang Basin, located in the north–central Qinghai–Tibet Plateau, preserves a sedimentary record of the eastern Tethys and has been the subject of research for decades (Kapp et al., 2000, 2007; Zhang et al., 2006a, 2006b, 2006c, 2011; Pullen et al., 2008; Zhang and Tang, 2009; Zhu et al., 2011a). The Late Jurassic to Early Cretaceous was an important interval in the geologic history of this region, because this is when the critical shift from marine to continental settings (i.e., the closure of the Mesotethys Ocean) is thought to have occurred (Zhang, 2000, 2004; Zhang et al., 2002, 2004a, 2007a, 2012, 2014; Skelton et al., 2003). However, the timing of this transition is debated, with proposed closure ages ranging from the Late Jurassic to the Early Cretaceous (Zhang, 2000, 2004; Zhang et al., 2002, 2004a; Lu et al., 2003; Mo and Pan, 2006; Shi, 2007), and even into the Late Cretaceous (Zhang et al., 2012, 2014). Therefore, whether the Late Jurassic–Early Cretaceous environment of the Qiangtang Basin was marine or continental is of critical importance for constraining the timing of closure of the Mesotethys. As tectonic events are typically reflected in sedimentological responses, sedimentary analysis can be used to improve our

⁎ Corresponding author. E-mail address: [email protected] (J. Cao).

http://dx.doi.org/10.1016/j.palaeo.2017.02.031 0031-0182/© 2017 Elsevier B.V. All rights reserved.

understanding of this issue (Blair and Bilodeau, 1988; Zhang, 2000, 2004; Zhang et al., 2002, 2004a, 2006c; Herrle et al., 2003; Brumsack, 2006; Sha et al., 2008; Yang, 2013). Sedimentological studies of the Bangong Mesotethys, with implications for the Jurassic–Cretaceous tectonic evolution of the Mesotethys Ocean, have been conducted since the 1980s (Yin et al., 1988; Leeder et al., 1988; XZBGM, 1993; Zhang, 2000, 2004; Zhang et al., 2002, 2004a, 2007a, 2012; Yi et al., 2003; Baxter et al., 2009). However, the Jurassic–Cretaceous marine sedimentary succession has not been fully described, and requires additional research. Recently, a widespread and continuous succession of Upper Jurassic or Lower Cretaceous black shales and oil shales was described cropping out along the Shengli River, in the northern part of the Qiangtang Basin (Fu et al., 2009; Yang et al., 2015). Based on the regional geology, these shales were deposited in a marine embayment (Fu et al., 2010, 2011). These strata provide a new opportunity to study the Late Jurassic– Early Cretaceous depositional environment of the basin, further constraining the timing of closure of the Mesotethys. Previous studies have not investigated these issues, instead focusing primarily on the depositional age, provenance, and hydrocarbon potential of the shales (Wang et al., 2007, 2009; Fu et al., 2010, 2011, 2012; Yang et al., 2015). Environmental interpretations may have been avoided in the past due to the poor preservation of the section, and particularly sedimentary structures (Fu et al., 2010, 2011), caused by high maturity (Yang et al., 2015).

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Using samples from a representative section of these shales, we conducted a preliminary study to characterize the Late Jurassic–Early Cretaceous paleoenvironments of the Qiantang Basin. Our interpretations are based on a combination of multiple petrologic and geochemical approaches, and the results of this study have broader implications for mechanisms of black shale deposition, organic matter accumulation, and hydrocarbon resource potential.

2. Geologic setting 2.1. Tectonic setting The Qiangtang Basin, which has an area of approximately 180,000 km2, is located in the northern–central part of the Qinghai– Tibet Plateau (32–35°N, 83–93°E). The basin is situated between the Hoh Xil–Jinshajiang and Bangonghu–Nujiang suture zones (Fig. 1A), and consists of three secondary structural units: the Northern

Qiangtang Depression, the Central Uplift, and the Southern Qiangtang Depression (Fig. 1A). The basin is part of the Qiangtang Block, which is believed to represent a large-scale anticlinorium (XZBGM, 1993). The central part of the block is an anticline composed of pre-Jurassic sedimentary strata and metamorphic rocks, while the northern and southern depressions are synclines composed mainly of Mesozoic sedimentary rocks (XZBGM, 1993; Zhang et al., 2002, 2006c). The central Qiangtang metamorphic belt is ~500 km long and up to 100 km wide, and is made up of blueschist- (Kapp et al., 2000, 2007; Zhang et al., 2006a) and ecoligite-bearing rocks (Zhang et al., 2006b; Pullen et al., 2008; Zhang and Tang, 2009) which represent the in situ expression of the paleo-Tethyan suture (Shuanghu suture; Zhang et al., 2006a, 2006b, 2006c). During the Early and Middle Triassic, this metamorphic belt was underthrust by the Hoh Xil–Jinsha suture, leading to the exhumation of the interior of the Qiangtang Block (Kapp et al., 2000, 2007; Pullen et al., 2008). This belt separates the Qiangtang block into northern and southern sections. During the Late Triassic to

Fig. 1. Geological map of the study area. (A) Structural units of the Qiangtang Basin. (B) Simplified geological map of the sampling area. (Modified after Fu et al., 2010)

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Jurassic, the amalgamation of the Qiangtang and Kunlun blocks formed a foreland basin in the study area (Leeder et al., 1988; Zhang et al., 2006c). The basin was mostly submerged, and accumulated a succession of carbonates interbedded with clastic rocks, thousands of meters thick (Zhao et al., 2000). From the Early to Middle Cretaceous, the collision of the Lhasa Block with the Qiangtang Block resulted in the subduction of the Bangonghu– Nujiang Ocean (the Mesotethys). This closure process appears to have continued into the Late Cretaceous, as evidenced by the presence of ophiolites rich in Middle Cretaceous radiolarians in central Tibet, dated to 132–104 Ma (Baxter et al., 2009; Zhang et al., 2012, 2014). This subduction resulted in the uplift of the central Qiangtang Block, forming volcanic arcs and causing a marked marine regression. Finally, following the flat-slab subduction of the Yalung–Zangpo Plate (i.e., Neotethys Ocean) during the Late Cretaceous, (as recorded by adakites in the southernmost Lhasa Block), marine sedimentation ceased as the Qiantang Basin transitioned to a continental remnant basin (Zhang et al., 2012). Thus, the change from marine to continental environments is thought to have taken place during the Cretaceous (Zhao et al., 2000; Hu et al., 2001; Fu et al., 2011; Zhang et al., 2012, 2014). 2.2. Stratigraphy Five sedimentary units were deposited in the Northern Qiangtang Depression during the Jurassic and Cretaceous: the Lower Jurassic Quse Formation, the Middle Jurassic Quemocui, Buqu and Xiali formations, the Upper Jurassic to Lower Cretaceous Suowa Formation, and the Lower to Upper Cretaceous Xueshan Formation, in ascending order (Wang et al., 2007). The black shales examined in this study occur in the upper part of the Suowa Formation (Fig. 2; Wang et al., 2007). The Suowa Formation was deposited in a continental to marine transitional environment, and consists of two lithologically distinct members. The Lower Member is composed of Jurassic marine carbonates, and is widely distributed in the Southern and Northern Qiangtang depressions (Zhao et al., 2000). The Upper Member consists of finegrained Upper Jurassic to Lower Cretaceous clastic strata, which formed during the transition from continental to marine deposition. This part of the formation occurs mainly in the central and western parts of the

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Northern Qiangtang Depression. Deltaic mudstones, siltstones, and sandstones were deposited to the east of the Quemocuo–Dazhuoma area, while sandstones, micrites, marls, limestones, and shales were deposited in tidal-flat or lagoon facies in the central part of the Northern Qiangtang Depression, and are overlain by thick layers of gypsum (Zeng et al., 2012). The Upper Jurassic–Lower Cretaceous black shales that form the focus of this study are widely distributed, and crop out along the Shengli River, on Changshe Mountain, and in the Nadge, Kangri, and Tuonamu areas. The stratigraphic thickness of the shales generally decreases toward the east (Zeng et al., 2012; Fig. 3). Due to an absence of index fossils or volcanic ash layers suitable for precise radiometric dating (e.g., zircon U\\Pb dating), it presently cannot be determined whether the black shales and oil shales are Late Jurassic or Early Cretaceous in age. Using the Re\\Os dating method, Wang et al. (2007) proposed that the black shales in the lower part of the Upper Member of the Suowa Formation were deposited around 113 ± 29 Ma. Fang et al. (2002), using electron spin resonance dating (ESR), reported that the overlying Xueshan Formation was deposited from 93 to 103.5 Ma. Numerous sporopollen taxa have been identified in black shales from the northern Qiangtang Depression, including Apiculatisporites, Cyathidites minor, Cicatricosisporites, Jiaohepollis, Cerebropollenites sp., Chasmatosporite, Ephedripites, Cycadopites, and Classopollis at the Shengli River section (Wang et al., 2007); Classopollis sp., Vitreisporites sp., Cyathidites australis, Lygodiumsporites subsimplex, Klukisporites sp., Pinuspollenites sp., and Abietineaepollenites at the Nagde Kangri section (Zhu et al., 2012); and Dicheiropollis, Cyathidites sp., and Classopollis sp. at the Tuonamu section. Classopollis commonly occurs in Jurassic to Early Cretaceous strata, and Cicatricosisporites, Jiaohepollis, Cerebropollenites sp., Cycadopites and Lygodiumsporites are all typical Cretaceous sporopollens. Combining the sporopollen biostratigraphy with the dating methods described above, we can conclude that the black shales and oil shales in the upper part of Suowa Formation were deposited during the Late Jurassic or Early Cretaceous. 2.3. Lithology The shales included in this study are fine-grained, clastic sedimentary rocks composed primarily of mud. The mud is a mixture of

Fig. 2. Middle Jurassic to Lower Cretaceous strata in the northern Qiangtang Depression, with black shales and oil shales from the Shengli River section. Biostratigraphic data are from Zhao et al. (2000), Fang et al. (2002), Zeng et al. (2012) and Zhu et al. (2012).

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Fig. 3. Stratigraphic correlation of Upper Jurassic–Lower Cretaceous black shales and oil shales in the Northern Qiangtang Depression. Based on Zeng et al. (2012) and Zhu et al. (2012).

phylosilicates (clays) and silt-sized particles of other minerals, primarily quartz and calcite; the ratio of clay to other minerals is variable (Blatt et al., 2005). The shales are fissile, typically breaking along parallel laminae less than one centimeter in thickness. In addition to shales, two other rocks types are present in the studied section: micrite and marl (Fig. 2). Although the CaCO3 content of all samples is N50%, our samples have been divided lithologically into shale, micrite, and marl based on their chemical composition and fissility (see discussion in Section 4.4.1 for detail). 3. Samples and methods A total of 28 samples, consisting of black shale, micrite, and gray to dark-gray marl, were collected from a section located on the west bank of the Shengli River (33° 34′ N, 87° 30′ E), in the northern Qiangtang Basin (Fig. 1B). Details of the sample numbers and lithologies are presented in Fig. 2. To minimize the potential effects of surface weathering and contamination, material on weathered surfaces was totally removed, and the samples were then cleaned thoroughly using distilled water. Systematic analyses of organic petrology, and organic and inorganic geochemistry were conducted for all samples. 3.1. Organic petrological analyses Organic petrological observations were made at the State Key Laboratory for Mineral Deposit Research, at Nanjing University. Techniques included conventional thin section microscopy and scanning electron microscopy (SEM). All 28 samples were sectioned perpendicular to

bedding, then embedded in Buehler's epoxy resin with hardener (in a 5:1 ratio), allowed to dry, and polished. Procedures are described in Taylor et al. (1998) and Amijaya and Littke (2006). The thin sections were examined at various magnifications (from ×5 to ×100) with incident white light and blue light excitation, using a Nikon LHS2H100C21 microscope. For SEM observation, air-dried samples were mounted on stubs using double-sided tape, and coated with Pt–Pd in a Polaron E5000 sputter coater for 2 × 2 min at 1.2 kV. Samples were then studied using a JSM-6490 scanning electron microscope, operating at an accelerating voltage of 15.0 kV with a beam current of 1.00–2.00 × 10−9 A. 3.2. Organic geochemical analyses Organic geochemical measurements conducted for this study include total organic carbon (TOC), total sulfur (TS), kerogen stable carbon isotopes (δ13Ckerogen), chloroform-extractable bitumen and corresponding carbon isotope ratios (δ13Cbitumen, δ13Csaturates and δ13 Caromatics), and biomarkers. All analyses were performed at the Wuxi Research Institute of Petroleum Geology, SINOPEC. TOC and TS values were measured using a LECO-CS-200 carbon/sulfur analyzer. Prior to analysis, samples were powdered to b 100 mesh and treated with HCl at 60 °C to remove carbonate, then rinsed with distilled water to remove the HCl. Carbon isotope (δ13C) values of kerogen were determined using the residues of powdered samples that were reacted with concentrated HCl, then heated in an HCl–HF solution, to remove carbonate and silicate minerals respectively. The residues then underwent heavy liquid separation and subsequent washing, before being extracted with CHCl3 to

R. Yang et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 473 (2017) 41–56

eliminate soluble organic matter (OM), leaving kerogen. δ13C values were determined using a MAT 253 isotope mass spectrometer, with a precision better than 0.1‰. Isotopic ratios are reported in standard δnotation relative to the Vienna Peedee Belemnite (VPDB) standard. Isotopic measurements were calibrated to Chinese national standard charcoal sample GBW 04407 (δ13 CVPDB = −22.43‰ ± 0.07‰). To extract chloroform bitumen, and separate the various bitumen components, powdered samples (b 100 mesh) were extracted for 72 h using a Soxhlet apparatus, with a solvent mixture of dichloromethane and methanol (93:7). The resulting extracts (after evaporation and concentration) were termed chloroform bitumen. The concentrated extracts were mixed with n-hexane and allowed to stand for 12 h to yield asphaltene. The residual extracts were then fractionated using open silica gel column chromatography, with a sequence of solvents: n-hexane, followed by a mixture of n-hexane and dichloromethane (2:1), and finally methanol, yielding saturated hydrocarbons, aromatic hydrocarbons, and nonhydrocarbons respectively. Stable carbon isotope compositions of extracted bitumen components were determined using the static combustion method, following the analytical procedure of Engel and Maynard (1989). The saturated and aromatic hydrocarbons were combusted in quartz tubes at 800 °C for 10 min. After cooling, the CO2 produced from the combustion was introduced directly into the inlet system of a Finnigan MAT 253 mass spectrometer. The results are reported in standard δ-notation relative to Vienna Peedee Belemnite (VPDB); the precision of the measurements was better than 0.1‰. The saturated hydrocarbon fraction was analyzed via gas chromatography (GC) and gas chromatography–mass spectrometry (GC–MS). The GC analysis was conducted using an HP 6890 gas chromatograph, fitted with a 30 m × 0.32 mm i.d. HP-5 column with a film thickness of 0.25 μm, using N2 as a carrier gas. The GC oven temperature was held at 80 °C for the initial 5 min, then was increased to 290 °C at 4 °C/min, and held at 290 °C for 30 min. The GC–MS analysis was conducted using an Agilent 5973I mass spectrometer interfaced with an HP6890 gas chromatograph, configured as described above, but employing He as a carrier gas. The GC oven temperature during GC– MS analysis was held at 60 °C for the initial 5 min, then increased to 120 °C at 8 °C/min, further increased to 290 °C at 2 °C/min, then held at 290 °C for 30 min.

3.3. Inorganic geochemical analyses Inorganic geochemical analyses of major, trace and rare earth elements were carried out for this study. A total of 19 samples were analyzed, at the Institute of Geochemistry, Chinese Academy of Sciences. For major element analysis, samples were dried at 70 °C, crushed using an Alceramic shatter box, then sieved using a nylon sieve to b75 μm. The resulting powder was baked at 105 °C to remove adsorbed water prior to analysis. Approximately 500 mg of powder and 4 g of Li2B4O7 was fused into a glass disc at 1200 °C; major elements were then measured using X-ray fluorescence spectroscopy (XRF). Analytical precision for major elements was within 1%, and the detection limits were generally lower than 3%. Trace and rare earth elements were measured using inductively coupled plasma mass spectrometry (ICP-MS), following the procedure described by Qi et al. (2000). Powdered samples (b200 mesh) were treated with 0.5 ml of HF and 1 ml of HNO3, in a Teflon bomb. The sealed bombs were then placed in an electric oven, and heated to 185 °C for 12 h. After cooling, the bombs were opened and evaporated to dryness on a hot plate, followed by a second treatment with HNO3 and evaporation to dryness. An internal standard solution composed of 2 ml of HNO3 and 1 ml of 500 ng/ml Rh was then added to the Teflon bomb, which was sealed and placed in an electric oven at 140 °C for 5 h to dissolve the residue. After cooling, 0.4 ml of the final solution was decanted into a 15 ml centrifuge tube and diluted to 8–10 ml with distilled de-

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ionized water for ICP-MS analysis. The analytical precision was better than 5%.

4. Results 4.1. Organic petrology Using organic petrological analyses, we characterized the morphology, quantity, size, and fluorescence of biological materials and fossil remains in the examined samples. Results indicate that benthic algae were the main source of organic matter in the shales, marls, and micrites. The samples contain calcified skeletal fragments of benthic fauna and microorganisms, however small amounts of vascular plant pollen were also detected. Benthic algae are the most common biological remains in the studied black shales. Identification of the benthic algae was based on comparisons to previously described specimens of various ages collected from China. The leafy and strip-shaped benthic algae and their sporangial debris (Fig. 4A–D) are similar to the benthic red algae described from the Precambrian Xiamaling oil shales in Hebei, in northern China (Bian et al., 2005), early Cambrian sediments of the Tarim Basin of northwestern China (Zhang et al., 2004b), the Upper Triassic Yanchang Formation in the southwestern part of the Ordos Basin, in central China (Ji and Xu, 2007), Jurassic sediments in the northern Qaidam Basin of northwestern China (Cao et al., 2009), Eocene sediments in the Boxing Sag, within the Dongying Depression of the Bohai Bay Basin in eastern China (Jiang et al., 2011), and the Eocene Huadian Formation oil shales, in Huadian, northeastern China (Xie et al., 2014). These similarities suggest that the benthic algae described here (Fig. 4A–D) are likely red algae. The benthic algae have been divided into three types based on their morphology and occurrence. The first type morphologically resembles large leaves or porous strips, and occurs in samples SLR-04 to SLR-28. These remains appear to be highly mature and carbonized, as they are red-brown to black under plane-polarized light and show little fluorescence (Fig. 4A and B). The second type is lattice-like in shape, and is present in all samples. The lattice is formed by brown filaments, and the mesh is typically filled with clastic material. Some of the specimens are located inside the shells of benthic animals, and are filled with carbonate. The lattices show brown–yellow fluorescence, mixed with the yellow–green fluorescence of the carbonate bitumen matrix (Fig. 4C). The third type consists of clusters of small spheres (10–50 μm in diameter), representing algal sporangial debris. They are dark brown to black in color under plane-polarized light, and show no fluorescence. This type of benthic algae is found mainly in marl samples SLR-01 to SLR-03 (Fig. 4D). Overall, these algae are widely distributed in all of the sampled marls, micrites, and calcareous shales. Calcified skeletal fragments belonging to benthic organisms are the second most abundant type of biological remains in the black shales. They occur mainly in samples SLR-04 to SLR-19 and are abundant (~ 50% of total fragments) in samples SLR-05 to SLR-15. The calcified skeletal fragments are derived primarily from bivalves, gastropods, and echinoderms, however small numbers of bryozoans and foraminifera are also observed. The skeletal remains are filled with carbonate cement, and show a yellow–green fluorescence. The brachiopods and gastropods are filled with coarse carbonate crystals (Fig. 4E), while the echinoderms are mostly filled with fine-grained calcite (Fig. 4F). Higher plant pollen was only detected in small amounts in marl samples SLR-01 to SLR-04. Pollen is present in samples as organic debris with a diameter of 10–15 μm. The grains are barely visible under planepolarized light, and show golden-yellow fluorescence (Fig. 4G). Remains of microorganisms were detected in all samples, and are abundant in the shales. The remains appear as tiny disseminated green spots. When examined under SEM, these microorganisms are usually spherical (~2 μm in diameter), and mostly occur on the surfaces of mineral grains such as calcite and apatite (Fig. 4H).

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Fig. 4. Photomicrographs of biological remains identified in Upper Jurassic–Lower Cretaceous black shales, marls, and micrites: (A) leafy benthic alga showing fluorescence, sample SLR-22, scale bar equals 100 μm; (B) carbonized strip-shaped benthic algae under plane-polarized light, sample SLR-24, scale bar equals 100 μm; (C) lattice-shaped benthic alga showing fluorescence, sample SLR-04, scale bar equals 100 μm; (D) sporangium of benthic alga showing fluorescence, sample SLR-02, scale bar equals 100 μm; (E) calcite-filled gastropod shell showing fluorescence, sample SLR-07, scale bar equals 200 μm; (F) calcified echinoderm skeletal debris in plane-polarized light, sample SLR-04, scale bar equals 400 μm; (G) vascular plant pollen showing fluorescence, sample SLR-01, scale bar equals 40 μm; (H) microorganism on a crystal surface under SEM, sample SLR-10, scale bar equals 10 μm. See Fig. 2 for sample locations and lithology.

4.2. Bulk organic geochemistry The bulk organic geochemical characteristics of the samples from the Shengli River section, including TOC, TS, δ13Ckerogen and δ13Cbitumen values, are reported in Supplementary Table 1. As shown in Fig. 5 and

Supplementary Table 1, TOC values range from 1.74% and 7.71%, while TS values range from 0.12% to 0.56%. In general, they show a linear correlation. TOC and TS initially increase up-section, then begin to decrease, with peak values in samples SLR-05 ~ SLR-11. The TOC and TS contents differ based on lithology, with marl samples having significantly lower

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4.3.3. Terpanes The terpanes detected in the studied samples are mainly tricyclic, tetracyclic, and pentacyclic (hopanes) terpanes, with a small amount of bicyclic sesquiterpanes. For the tricyclic terpanes, the carbon number ranges between C19 and C29, peaking at C23 in all samples. C24 tetracyclic terpane was identified in all samples, with the ratio of C24 tetracyclic terpane/(C24 tetracyclic terpane + C26 tricyclic terpane) (i.e., C24Tet/ (C24Tet + C26TT) ranging from 0.79 to 0.84. Pentacyclic triterpanes (hopanes) were detected in all samples, with carbon numbers ranging between C27 and C35, with a peak at C30. 4.3.4. Steranes Steranes were detected in all analyzed samples, including C21–C22 and C27–C29 regular steranes and C27–C29 rearranged steranes. The regular C27–C29 steranes in all samples show an asymmetrical V-shaped distribution (C27 N C29 N C28), with the abundance of C27, C28 and C29 being 44–51%, 14–17% and 32–40%, respectively. 4.4. Inorganic geochemistry

Fig. 5. TOC vs. TS values in Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. The boundary threshold between oxic and suboxic samples (TS/TOC = 0.15) is based on Berner and Raiswell (1983) and Algeo and Maynard (2004).

TOC and TS values (2.80% for TOC and 0.19% for TS) than micrites (6.03% for TOC and 0.40% for TS) or black shales (5.51% for TOC and 0.38% for TS). The δ13Ckerogen values of the samples in this study range from −22.0‰ to −20.3‰, with a mean of −21.1‰, and no significant difference between different lithologies. Measured δ13Cbitumen values range from − 22.6‰ to − 21.7‰, with an average value of − 22.0‰. Values are slightly lower than those of δ13Ckerogen, but the overall stratigraphic trend is similar (Fig. 2). This result is consistent with the typical pattern of stable carbon isotope partitioning between kerogen and derived bitumen (Peters et al., 2005). The carbon isotopes of the saturated (δ13Csat) and aromatic (δ13Caro) fractions of extracted bitumen range from − 25.0‰ to − 22.9‰ (mean = − 23.7‰) and from − 22.1‰ to −21.5‰ (mean = −21.7), respectively. 4.3. Biomarkers Several types of biomarkers were examined in the black shales, marls, and micrites, including alkanes, acyclic isoprenoids, terpanes, and steranes (Supplementary Table 1, Fig. 7). 4.3.1. n-Alkanes Samples in this study yielded n-alkanes distributed between n-C12 and n-C35 forming a single peak, with the exception of sample SLR-01. The major peak is from n-C17 to n-C19. The odd–even predominance (OEP, defined as [(Ci + 6Ci + 2 + Ci + 4) / (4Ci + 1 + 4Ci + 3)](−1)i + 1, of which the “i + 2” denotes the n-alkane number with the main peak) ratio values are 0.97–1.04; i.e., with no obvious odd–even pre+ dominance. The ratio of C− 21/C22 is over 1.5 except for sample SLR-01, which has a value of 1.12. 4.3.2. Acyclic isoprenoids Acyclic isoprenoids were detected in all studied samples. The pristane/phytane ratio (Pr/Ph) is the most commonly used acyclic isoprenoid proxy for environmental conditions (Didyk et al., 1978). In the Shengli River samples, the Pr/Ph ratio ranges from 0.71 to 0.88, with little difference between marl (0.73–0.88), micrite (0.87) and shale (0.71– 0.83) samples (Supplementary Table 1). The ratios of isoprenoids to their adjacent n-alkanes are 0.39–0.61 for Pr/n-C17 and 0.48–0.71 for Ph/n-C18; they show little difference between different sample lithologies, similar to the Pr/Ph values.

The inorganic (elemental) geochemistry of a sedimentary rock, including the major, trace and rare earth element composition, is a good indicator of depositional environment (Taylor and McLennan, 1985; Morford and Emerson, 1999). 4.4.1. Major elements The concentrations of major elements, presented as oxides, in the samples from the Shengli River section are shown in Fig. 8 and Supplementary Table 2. In general, the most abundant major element oxides are CaO (28.6–38.5%) and SiO2 (4.22–25.2%). In contrast, Al2O3 (0.58– 6.67%), MgO (0.83–13.0%), Fe2O3 (0.77–3.63%) and K2O (0.17–1.66%) have relatively low concentrations (1–10%). The concentrations of other major element oxides, including P2O5 (0.04–0.87%), TiO2 (0.03– 0.30%), Na2O (0.03–0.09%) and MnO (0.03–0.06%), are all b1%. When compared to samples with different lithologies, the marl samples (SLR-01–SLR-04) are rich in Si, Al and Ti, indicating more terrigenous input during the deposition of these beds. In contrast, the micrite sample SLR-12 is rich in Ca, and especially Mg (13.0%). The shale samples display major element compositions between those of the marl and micrite samples. The strong correlations between SiO2, Al2O3, and TiO2 suggest that all three of these elements reflect the relative amount of terrigenous input (Kennedy et al., 2002), with Si and Al derived mainly from terrigenous debris. In contrast, there is no significant correlation between Al2O3 and Fe, Ca, P, Mn, Mg, or other oxides, indicating that these elements are more likely to have been controlled by conditions in the depositional environment. The high abundance of Si, Al and Ti in marl samples indicates more intense weathering and terrigenous input during the deposition of marls than during the deposition of micrites or shales, which is consistent with the petrographic observations (Section 4.1). The high Mg content of sample SLR-12 (MgO = 13.0%, 3 times greater than observed in any other sample) indicates that the sample was deposited during an arid period, with strong evaporation (Folk and Land, 1975). The chemical composition of marine sediments is controlled by the three main sediment sources: terrigenous, biogenic and hydrothermal. Boström et al. (1973) proposed that sediments with a Fe/Ti ratio of N20, or an Al/(Al + Fe) ratio of b 0.4, reflect a hydrothermal influence, whereas a Fe/Ti ratio of b 20 or an Al/(Al + Fe) ratio of N 0.4 reflect a predominantly terrigenous source. As shown in Fig. 8, all marls and most shale samples in this study have a Fe/Ti ratio of b20, and an Al/ (Al + Fe) ratio of N0.4, suggesting little hydrothermal influence during shale deposition. The micritic sample SLR-12, and three shale samples with TOC N 6.0% (SLR-05, SLR-07 and SLR-08), show slight hydrothermal influence.

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4.4.2. Trace elements The concentrations of trace elements, including U, V, Zn, Pb, Cu, Ni, Cr and Co, are listed in Supplementary Table 3. In general, the samples contain a greater amount of V (27.3–69 ppm), Ni (20.7–64 ppm), Zn (20.4–104 ppm), Cr (8.9–67 ppm), and Cu (5.0–39 ppm) than they do of elements such as Pb (2.33–19.6 ppm), Co (1.77–10.6 ppm), and U (0.35–2.93 ppm). 4.4.3. Rare earth elements The concentrations of rare earth elements (REE) in samples from the Shengli River section are reported in Supplementary Table 4. The North American shale composite (NASC)-normalized REE distribution patterns are presented in Fig. 9, and the stratigraphic distribution of REE ratios is illustrated in Fig. 10. As shown in Supplementary Table 4, the total REE concentration in the analyzed samples is between 9.36 and 54.5 ppm, much lower than that of NASC (173 ppm), which reflects the effects of lithology and provenance on REE content (Abanda and Hannigan, 2006; Santos et al., 2009). Total REE concentrations are highest in marls, intermediate in shales, and lowest in micrites, indicating that terrigenous input may have been the major source of REEs (Murray et al., 1991). This view is also supported by the strong positive correlation between the total REE concentration and Al, Si and Ti (Figs. 8 and 10). All the samples show similar NASC-normalized REE patterns, with enrichment in light REEs (Fig. 9). This indicates that all three sample types, although exhibiting different total REE content, had a similar source of terrigenous sediment (Johannesson and Hendry, 2000). Specifically, the ratio of light REEs to heavy REEs is 7.00–9.32, similar to that of NASC (7.50). The (La/Yb)n ratio is 1.24–1.56, the (La/Sm)n ratio is 1.10–1.39, and the (Tb/Yb)n ratio is 1.10–1.30. The Ce and Eu anomalies are commonly used proxies in REE analysis, and are typically defined by the equations δCe = Cen / (Lan × Prn)1/2 and δEu = Eun / (Smn × Gdn)1/2. Values b0.95 represent a negative anomaly, while values N 1.0 represent a positive anomaly (Wang et al., 1989). In our section, samples SLR-01 to SLR-05 have δCe values between 0.85 and 0.87, while all other samples have values b 0.81. Thus, all samples show negative Ce anomalies (average = 0.81). The δEu values of all samples are between 0.84 and 1.06 (average = 0.96), showing no apparent Eu anomalies or a slightly negative anomaly. 5. Discussion Based on our petrologic and geochemical results, we interpreted the depositional environments represented by the strata in the Shengli River section. This analysis will help to enhance our understanding of Late Jurassic–Early Cretaceous depositional environments in the Qiangtang Basin, and by extention the timing of closure of the Mesotethys. 5.1. Depositional environments Based on optical microscope observations, red algae was identified as the predominant biological precursor in the studied samples. There are approximately 2500–6000 species of red algae in the world (Woelkerling, 1990), with N97% of these species living in marine environments (intertidal or sublittoral areas), and only 36 genera and fewer than 200 species found in freshwater (Seckbach and Chapman, 2010; Xie et al., 2014; Zhang et al., 2007b). Freshwater red algae live primarily in fast-flowing streams and rivers (Sheath and Hambrock, 1990; Xie et al., 2014). However, the black shales in this study exhibit laminations, implying a low-energy environment and ruling out a freshwater origin for the red algae. Thus, we can infer that the Upper Jurassic– Lower Cretaceous black shales in the studied section were deposited in a marine environment. In addition to benthic algae, the calcified shells of benthic organisms were identified in most micrite and shale samples, which also implies a supralittoral or facies, in a transitional to

marine environment (Bian et al., 2005; Han et al., 2015). Vascular plant pollen, indicative of terrigenous input, were only identified in marl samples. The depositional environment of the marl samples was likely more continental in nature than the depositional environment of the other samples. Overall, a marine to continental transitional environment can be inferred. The relationship between TOC and TS values can be used to interpret the depositional environment of sediments (Berner and Raiswell, 1983; Leventhal, 1987; Rimmer et al., 2004). In normal marine sediments, TOC and TS values are positively correlated, and the intercept is close to 0; in contrast, TOC and TS values show no correlation in strongly reducing and anoxic environments, and the intercept is N0 (Leventhal, 1987). The dissolved oxygen content of water in the depositional environment can also be assessed using the TS/TOC ratio; samples with a TS/TOC value b0.15 represent a suboxic to anoxic environment while TS/TOC values N 0.15 represent an oxic environment (according to Berner and Raiswell (1983), TOC/S(py) ≈ 5 in normal Early Cretaceous marine environments; here we use total sulfur (TS) as a proxy for sulfide S, which composes 60–90% of total sulfur in similar facies (Algeo and Maynard, 2004)). TOC and TS values in the studied section show a relatively strong positive correlation (Fig. 5), with an intercept close to 0. This suggests that the samples from the Shengli River section were mainly deposited in a suboxic to anoxic marine environment. Previous studies have reported that the correlation between carbon isotopes in the saturated (δ13Csat) and aromatic fractions (δ13Caro) of extracted bitumen can be used to distinguish between marine and continental depositional environments; the boundary line between the marine and continental fields is defined by the equation: δ13Caro = 1.14 × δ13Csat + 5.46 (Sofer, 1984). Most samples from the Shengli River section plot on the marine side, clustered around the boundary (Fig. 6A), suggesting a transitional to marine depositional environment. In addition, the δ13Ckerogen (− 22.0‰ to − 20.3‰) and δ13Cbitumen (− 22.6‰ and − 21.7‰) values indicate deposition of marine organic matter in a brackish to saline environment (Degens, 1969; Schidlowski et al., 1994; Schouten et al., 2001), which is consistent with the results obtained using other geochemical proxies. Biomarker characteristics can reflect the composition of precursor organic material, the sedimentary environment, and the thermal evolution of organic matter in sediments (Peters et al., 2005). The samples included in this study show major peaks in light n-alkanes (n-C17 to n+ C19), no obvious odd-even predominance, and low C− 21/C22 values. These trends can potentially be explained in two ways. One is that the aquatic (marine or lacustrine) algae were the major source of the organic matter. The other interpretation is high maturity; regardless of the type of biological precursor, the maturation process will involve the break-up of high-molecular-weight hydrocarbons into low-molecularweight hydrocarbons, leading to a preponderance of light n-alkanes (Peters et al., 2005). Given that benthic algae (rather than lacustrine aquatic algae) constitute the most common type of biological remains in the studied samples (Section 4.1), the predominance of light-molecular-weight alkanes here is likely to have resulted from high maturity, which is also supported by the low fluorescence intensity (Fig. 4). Furthermore, organic matter in the studied samples has VRo values of ~1.3%, suggesting high maturity (Qin, 2006a; Yang et al., 2015). Plotting the ratios of acyclic isoprenoids to their adjacent n-alkanes (i.e. Pr/n-C17 and Ph/n-C18), all samples fall into the field indicating reducing conditions in a saline depositional environment (Fig. 6B; Shanmugam, 1985; Fu et al., 1991), which is consistent with other analyses. Tricyclic terpanes are believed to be derived mainly from prokaryotic membranes or algae (Ourisson et al., 1982; Volkman et al., 1986), but there are also some tricyclic terpanes derived from vascular plants, such as C19 tricyclic terpane (Yang et al., 2008). The concentrations of C20, C21, and C23 tricyclic terpanes in all samples follow the order C20 b C21 b C23, indicating a similar composition of algal bio-precursors, which is consistent with our organic petrologic results (Section 4.1). Peters et al.

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Fig. 6. Cross plots of various organic geochemical parameters, reflecting the depositional facies of Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. (A) δ13Caro vs. δ13Csat (dashed line refers to the equation δ13Caro = 1.14 × δ13Csat + 5.46; Sofer, 1984); (B) Pr/n-C17 vs. Ph/n-C18 (Shanmugam, 1985); (C) C22/C21 vs. C24/C23 in tricyclic terpanes (Peters et al., 2005); (D) C26/C25 in tricyclic terpanes vs. C31R/C30 in hopanes (Peters et al., 2005).

(2005) proposed that the C22/C21, C24/C23, and C26/C25 ratios in tricyclic terpanes can be used to differentiate rock extracts and oils derived from different depositional environments; e.g., marine, lacustrine, transitional, and coaliferous. In this study, the C22/C21, C24/C23, and C26/C25 ratios range from 0.32 to 0.36, 0.50 to 0.57, and 0.66 to 0.75, respectively. There is little difference between the various lithologies, implying similar depositional environments for all samples. In crossplots of the C22/ C21 ratio vs. the C24/C23 ratio in tricyclic terpanes (Fig. 6C), and the C26/C25 ratio in tricyclic terpanes vs. the C31R/C30 ratio in hopanes (Fig. 6D), all samples from the Shengli River section plot in the field representing transitional–marine depositional environments. This interpretation is in good agreement with the petrological observations and other geochemical analyses discussed above. A high C24 tetracyclic terpane content indicates a highly saline depositional environment (Clark and Philp, 1989). The enrichment of C24 tetracyclic terpane (C24Tet / (C24Tet + C26TT) = 0.79–0.84) in the samples included in this study indicates deposition in a saline marine environment (Peters et al., 2005). Pentacyclic triterpanes (hopanes) are usually derived from prokaryotes or vascular (higher) plants, rather than eukaryotic algae; bacteria are also an important source of hopanes (Ourisson et al., 1979; Peters et al., 2005). Some hopane ratios can be used to indicate redox conditions. For example, a high abundance of C35 homohopane generally indicates marine carbonates or evaporites (Fu et al., 1986; Clark and Philp, 1989). The ratio of C35/(C31–C35) homohopanes (HHI) has been successfully applied to determine depositional environments. HHI is typically N0.1 in salt water, and b0.06 in freshwater (Peters and

Moldowan, 1991). The HHI value of the samples from the Shengli River section is between 0.08 and 0.1, indicating a marine brackish environment. Another hopane indicative of depositional environment is gammacerane, although the biological source of this compound is not fully understood; it is generally considered to be a diagenetic product of tetrahymanol (Peters et al., 2005). Gammacerane usually occurs in saline environments and is typically more abundant than C30 hopane in environments with extremely high salinity (Sinninghe Damsté et al., 1995). High levels of gammacerane are typically interpreted as a sign of water column stratification. The values of the gammacerane index (gammacerane/C30 hopane) of rock extracts and oils from freshwater, marine, and saline lake environments are usually b 0.2, 0.2–0.4, and N 0.4 respectively (Fu et al., 1991). The gammacerane index of the samples included in this study is generally between 0.20 and 0.35, indicating a marine brackish environment. The ratio of C3122R/C30 hopane (C31R/C30H) can also be used to distinguish between marine and lacustrine sediments, as rock extracts and oils from marine shales, marls, and carbonates usually having a relatively high ratio (N 0.25, Peters et al., 2005). The C31R/C30H ratios of the samples in this study range from 0.30 to 0.41, indicative of marine sediments. In summary, based on the results of our organic petrology and organic geochemical analyses above, we conclude that the samples from the Shengli River section were deposited in a transitional marine environment, under saline to brackish water conditions. The depositional environment of the marl samples, especially sample SLR-01, was more

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Fig. 7. Mass chromatograms of Upper Jurassic–Lower Cretaceous black shales, marls, and micrites.

influenced by terrigenous input than that of the other samples. The micrite sample was likely deposited during an arid period, in an environment with strong evaporation (i.e., a lagoon). 5.2. Depositional redox conditions The acyclic isoprenoids pristane and phytane preserved in ancient rock extracts and oils are largely, though not exclusively, produced diagenetically via the catagenesis of phytol. Phytol is derived from phytyl, the major chlorophyll phytyl side-chain of photoautotrophs (Powell and McKirdy, 1973). It has long been known that phytol diagenetic reaction pathways are different under oxidizing and reducing conditions in the sediment. Reducing conditions promote the conversion of phytol to phytane, while oxidizing conditions promote the conversion of phytol to pristane (Didyk et al., 1978; Peters et al., 2005). Thus, a Pr/ Ph ratio of b 1.0 is generally indicative of sediments deposited under reducing conditions. In the studied Shengli River samples, the Pr/Ph value ranges from 0.71 to 0.88, indicating a reducing early diagenetic environment. Volkman et al. (1983) argued that C30 diahopane (C30*) originates from bacteria, via clay mineral catalysis under oxic to semi-oxic conditions. Therefore, rock extracts and oils derived from source rocks deposited in oxic environments normally have a high 17α(H)-diahopanes/ 18-α-30-norneohopane ratio (C30*/C29Ts). The C30*/C29Ts values of samples from the Shengli River section are between 0.09 and 0.19. When sample SLR-01 is excluded, the C30*/C29Ts value ranges from

0.09–0.13, indicating a reducing depositional environment. Sample SLR-01 may have been deposited in a more oxic environment, relative to the other samples. Oxic conditions and high Eh values often promote the conversion of sterols to regular sterenes, which are eventually reduced to rearranged steranes. Accordingly, organic extracts or oils with low values of rearranged/regular steranes are indicative of reducing environments (Moldowan et al., 1986). In this study, the C29 rearranged/regular sterane ratio ranges from 0.32 to 0.54. Except for sample SLR-01, C29 rearranged/regular sterane ratios are consistently b0.39, indicating a reducing environment. In contrast, the value for sample SLR-01 indicates a more oxic depositional environment. This is consistent with our C30 diahopane results discussed above. The trace element composition of sediments has been widely applied as a tool to constrain redox conditions (Morford and Emerson, 1999; Algeo and Lyons, 2006; Algeo and Tribovillard, 2009; Algeo and Rowe, 2012; Tribovillard et al., 2012). The underlying principle is that elements exhibit different geochemical behaviors under different conditions. As a result of changes in hydrodynamic and/or redox conditions, elements may migrate, aggregate, or precipitate. For example, redoxsensitive elements such as U, Mo, V and Ce will have high oxidation numbers (U (VI), Mo (VI), V (V), and Cu (II)), and remain mobile in an oxidizing environment; in contrast, they will be present with low oxidation numbers (U (IV), Mo (IV), V (III), and Cu (I)), and readily precipitate in a reducing environment (Jones and manning, 1994; Algeo and Maynard, 2004). According to Algeo and Maynard (2004), the behavior

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Fig. 8. Major element composition of Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. Data are reported in Supplementary Table 2.

of most redox-sensitive trace elements follows one of two patterns: (1) Some elements (such as Mo, U, V, Pb and Zn) exhibit moderate enrichment factors (EFs) and strong covariance with TOC in oxic or suboxic conditions, but exhibit high EFs and weak covariance with TOC in euxinic environment; these elements have a “strong euxinic affinity”; (2) other redox-sensitive elements (such as Cu, Cr, Ni and Co) exhibit relatively low EFs and moderate to strong covariance with TOC in all conditions, which are called “weak euxinic affinity”. The EF values of redox-sensitive trace elements (U, V, Pb, Zn, Cu, Cr, Ni and Co) in samples from the Shengli River section are reported in Supplementary Table 3. EFs were calculated following the formula: EFX = (X/Al)sample / (X/Al)PAAS (Taylor and McLennan, 1985). EF values

Fig. 9. NASC-normalized REE patterns of Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. NASC data are from Gromet et al. (1984).

of U, Pb, Zn, Ni and Cu in most samples are higher than 2.0, while EF values of V, Cr and Co are in the range of 1.0–2.0. The EFs of all redoxsensitive elements are higher in shale and micrite samples than in marl samples, and the EFs of V, Cr, Co and Cu in marl samples are sometimes lower than 1.0. This suggests that redox conditions during deposition of the marl units were oxic to dysoxic, while micrite and shale samples were probably deposited in suboxic to anoxic conditions. Because EFs of redox-sensitive trace elements in all samples are relatively low, the depositional environment was likely predominantly suboxic. Plotting Al-normalized elements against TOC is another technique for assessing depositional redox conditions (Algeo and Maynard, 2004; Fig. 11). Excluding those samples which may have been influenced by hydrothermal fluids (Section 4.4.1), Al-normalized element concentrations in most samples exhibit a strong (U, V, Pb and Cr) or moderate (Cu, Ni and Co) covariance with TOC, while Al-normalized Zn exhibit poor covariance with TOC. This suggests that the depositional environment was likely predominantly suboxic and some individual samples were deposited in anoxic environments (Algeo and Maynard, 2004). The anomalous EF values of V and Ni in some micrite and shale samples can be attributed to hydrothermal influence, which can be corroborated by the relatively high abundance of Mn and Fe (Boström et al., 1973; Liu et al., 2015). When interpreting rare earth elements (REEs), the Ce and Eu anomalies are commonly used to determine redox conditions in the depositional environment. Previous studies have shown that Ce tends to precipitate in the form of Ce4+ in oxidizing environments, leading to a decrease in the Ce content in seawater (Tenger et al., 2006). In contrast, Eu3+ will convert to Eu2+ and become concentrated in seawater in reducing environments. Therefore, positive Ce and Eu anomalies, or a lack of such anomalies in sediments, are indicative of a reducing

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Fig. 10. Stratigraphic distribution of REE parameters in Upper Jurassic–Lower Cretaceous black shales, marls, and micrites.

environment, while negative Ce and Eu anomalies indicate an oxic environment (Wright et al., 1987; Wilde et al., 1996; Tenger et al., 2006). In this study, all samples show a negative Ce anomaly (average = 0.81) and virtually no Eu anomaly (average = 0.96). The Ce anomaly indicates an oxidizing environment, but the Eu anomaly indicates an oxidizing to reducing transitional environment. This seemingly contradictory implies that proximity to a continental margin may have influenced the δCe value, resulting in noise in the data (Murray et al., 1990). Ceanom is an alternative definition of the Ce anomaly, calculated as: Ceanom = log [3 × Cen / (2 × Lan + Ndn)]. It has been observed that this ratio is generally greater than − 0.1 in sediments deposited in reducing environments, and less than − 0.1 in oxic environments (Wright et al., 1987). The Ceanom values of samples in this study range from − 0.11 to − 0.06, with an average value of − 0.10, indicating a transitional, oxidizing to reducing environment. This is consistent with the interpretation of δCe values discussed above. In summary, samples from the Shengli River section were deposited primarily in an oxygen-poor environment, with a dissolved oxygen content between 0.2 and 2.0 ml/l (i.e., suboxic). Only a few samples have trace-element ratios indicative of oxic or anoxic environments. However, because the ratios in these samples are close to the threshold values of oxygen-depleted environments, we interpret these samples as also having been deposited in weakly oxidizing to reducing environments. These results are generally consistent with our findings from organic petrology and other geochemical proxies presented above. 5.3. Evolution of the depositional environment Based on our comprehensive study of depositional environments and redox conditions, we conclude that during the Upper Jurassic–

Early Cretaceous, the Qiangtang Basin (Tibet) was generally marine– transitional (though typically to marine). During this period, several cycles of marine transgression and regression coincide with the uplift of the central Qiangtang Basin; as a consequence, the black shales and oil shales were generally deposited in transitional facies such as intertidal and lagoonal environments (Fu et al., 2009; Wang et al., 2010; Zeng et al., 2012). However, few detailed studies characterizing the depositional environment of the black shales and oil shales have been conducted, and the origin of the various lithologies remains enigmatic. Based on our analyses of organic petrology, organic/inorganic geochemistry, and isotope geochemistry, it is inferred that the black shales, marls, and micrites in the Shengli River section were deposited in a saline to brackish environment, under weakly oxic to suboxic conditions. The presence of small amounts of vascular plant pollen in the samples indicates that the rocks were deposited in a marine–continental transitional environment. The depositional facies varied over time, and is described in detail in Table 1. The majority of the Qiangtang Basin, with the exception of the central uplift, was submerged during the Late Jurassic (Fig. 12A). As a consequence, sediments were mainly deposited in deep basinal and platform facies, while marine–transitional deltaic facies were developed locally along the northern margin of the Northern Qiangtang Depression (Hu et al., 2001; Qin, 2006b; Zeng et al., 2012). Later, during the Early Cretaceous, the closure of the MesoTethys and associated collision of the Qiangtang and Lhasa blocks caused uplift of the central Qiangtang Basin, and most of the Northern Qiantang Depression was subaerially exposed. Only some areas in the Northern Qiangtang Depression remained submerged, forming a semi-enclosed bay open to the west and south (Zhao et al., 2000; Wang et al., 2010; Zhang et al., 2012).

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Fig. 11. Correlation of Al-normalized major elements with TOC (%) in Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. Red square indicates samples with possible hydrothermal influence. Data are reported in Supplementary Table 2. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Carbonate and clastic sediments developed throughout this area, in deltaic, tidal, and lagoonal facies. The areas that contain black shales and oil shales, such as the Shengli River, Changshe Mountain, Nadge Kangri, and Tuonamu (Li et al., 2005; Fu et al., 2010; Zhu et al., 2011b; Fig. 3) were mainly tidal-flat to lagoonal environments during the

Early Cretaceous (Fig. 12B). This finding not only has general implications for marine–continental transitional process in the Qiantang Basin during late Mesozoic, but also may provide evidence to constrain the timing of the Mesotethys closure (i.e., no earlier than the Late Jurassic–Early Cretaceous).

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Fig. 12. Schematic map showing the paleogeography of the Qiangtang Basin during the Late Jurassic and Early Cretaceous. (A) Late Jurassic (modified from Hu et al., 2001); (B) Early Cretaceous (modified from Zhang et al., 2012).

6. Conclusions The Upper Jurassic–Early Cretaceous black shales, oil shales, marls, and micrites seen in the Shengli River section in the northern Qiangtang Basin (Tibet) were deposited in primarily saline to brackish

environments, in marine to transitional facies. The environment ranged from weakly oxic to suboxic. The depositional environment of the rocks changed gradually over time. The marls at the base of the succession have relatively low organic matter content, and were influenced by terrigenous input. The climate

Table 1 Characteristics of depositional environments and associated characteristics inferred from Upper Jurassic–Lower Cretaceous black shales, marls, and micrites. Lithology

OM abundance TOC (%)

Depositional environment

Redox conditions

Other characteristics

Micritic limestone

High (5.61–6.46)

Lagoon

Suboxic – anoxic

Shale

High (4.12–7.71) Low (1.74–3.77)

Semi-closed lagoon

Suboxic

Shore to shallow lake

Weakly oxic

Abundant benthic algae and oriented calcified skeletal debris. Very low terrigenous input. Strong evaporation and high salinity. Abundant benthic algae and calcified skeletal debris. Marine transgression. Relatively more terrigenous plants and clastic input. Warm and humid climate. Low salinity.

Marl

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during deposition of the shales was warm and humid, and the shales were deposited in a weakly oxidizing shallow-lake or shore environment with low salinity. The middle and upper shale units have a high organic matter content, and were probably deposited in a reducing, semi-enclosed lagoon environment. The micrites interbedded with the shales also contain abundant organic matter, and were likely deposited in a lagoon environment with intense evaporation, high salinity, and water column stratification. Acknowledgments We sincerely thank Editor-in-Chief Thomas Algeo for his constructive and detailed comments and patience in handling this manuscript. We would also like to thank two anonymous reviewers for their constructive feedback, which helped us to greatly improve this study. This work was jointly funded by a Chinese National Science and Technology Major Project Grant (No. 2016ZX05002-006-005) and the National Natural Science Foundation of China (Grant Nos. 41322017 and 41472100). Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.palaeo.2017.02.031. References Abanda, P.A., Hannigan, R.E., 2006. Effect of diagenesis on trace element partitioning in shales. Chem. Geol. 230, 42–59. Algeo, T.J., Lyons, T.W., 2006. Mo–total organic carbon covariation in modern anoxic marine environments: implications for analysis of paleoredox and paleohydrographic conditions. Paleoceanography 21, PA1016. http://dx.doi.org/10.1029/2004PA001112. Algeo, T.J., Maynard, J.B., 2004. Trace-element behavior and redox facies in core shales of upper Pennsylvanian Kansas-type cyclothems. Chem. Geol. 206, 289–318. Algeo, T.J., Rowe, H., 2012. Paleoceanographic applications of trace-metal concentration data. Chem. Geol. 324, 6–18. Algeo, T.J., Tribovillard, N., 2009. Environmental analysis of paleoceanographic systems based on molybdenum–uranium covariation. Chem. Geol. 268, 211–225. Amijaya, H., Littke, R., 2006. Properties of thermally metamorphosed coal from Tanjung Enim Area, South Sumatra Basin, Indonesia with special reference to the coalification path of macerals. Int. J. Coal Geol. 66, 271–295. Baxter, A.T., Aitchison, J.C., Zyabrev, S.V., 2009. Radiolarian age constraints on Mesotethyan ocean evolution, and their implications for development of the Bangong–Nujiang suture, Tibet. J. Geol. Soc. 166, 689–694. Berner, R.A., Raiswell, R., 1983. Burial of organic carbon and pyrite sulfur in sediments over Phanerozoic time: a new theory. Geochim. Cosmochim. Acta 47, 855–862. Bian, L.Z., Zhang, S.C., Zhang, B.M., Wang, D.R., 2005. Red algal fossils discovered from the Neoproterozoic Xiamaling oil shales, Xiahuayuan town of Hebei Province. Acta Microbiol Sin. 22, 209–216 (in Chinese with English abstract). Blair, T.C., Bilodeau, W.L., 1988. Development of tectonic cyclothems in rift, pull-apart, and foreland basins: sedimentary response to episodic tectonism. Geology 16, 517–520. Blatt, H., Tracy, R.J., Owens, B.E., 2005. Petrology: Igneous, Sedimentary and Metamorphic. third ed. W.H. Freeman. Boström, K., Kraemer, T., Gartner, S., 1973. Provenance and accumulation rates of opaline silica, Al, Ti, Fe, Mn, Cu, Ni and Co in Pacific pelagic sediments. Chem. Geol. 11, 123–148. Brumsack, H.J., 2006. The trace metal content of recent organic carbon-rich sediments: implications for Cretaceous black shale formation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 232, 344–361. Cao, J., Bian, L.Z., Hu, K., Liu, Y.T., Wang, L.Q., Yang, S.Y., Chen, Y., Peng, X.Q., 2009. Benthic macro red alga: a new possible bio-precursor of Jurassic mudstone source rocks in the northern Qaidam Basin, northwestern China. Sci. China Earth Sci. 52, 647–654. Clark, J.P., Philp, R.P., 1989. Geochemical characterization of evaporite and carbonate depositional environments and correlation of associated crude oils in the Black Creek Basin, Alberta. Bull. Can. Petrol. Geol. 37, 401–416. Degens, E.T., 1969. Biogeochemistry of stable carbon isotopes. In: Eglinton, G.E., Murphy, M.T.J. (Eds.), Organic Geochemistry: Methods and Results. Springer, Berlin Heidelberg, pp. 304–329. Didyk, B.M., Simoneit, B.R.T., Brassell, S.C., Eglinton, G., 1978. Organic geochemical indicators of palaeoenvironmental conditions of sedimentation. Nature 272, 216–222. Engel, M.H., Maynard, R.J., 1989. Preparation of organic matter for stable carbon isotope determination by sealed tube combustion: a cautionary note. Anal. Chem. 61, 1996–1998. Fang, D.Q., Yun, J.B., Li, C., 2002. Discussion of the Xueshan Formation in the north of Qiangtang Basin, Qinghai–Tibet Plateau. J. Stratigr. 26, 68–72 (in Chinese with English abstract). Folk, R.L., Land, L.S., 1975. Mg/Ca ratio and salinity: two controls over crystallization of dolomite. AAPG Bull. 59, 60–68.

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