The studied samples came from the same wells discussed in Qiao et al. (2020). 80 samples were used to perform hydrocarbon extraction and fractionation as well as quantification of each compound and gas chromatography-mass spectrometry (GC-MS) analysis. 37 saturated hydrocarbon fractions were used to perform gas chromatography-flame ionization detector (GC-FID), and stable carbon isotope composition (δ¹³C) for all fractions was performed in 25 samples. Moreover, clay minerals compositions from 232 samples as well as porosity and oil saturation data from 83 samples that came from EF 2 well were provided by the Xinjiang Oil Field.

Before the test, all analyzed samples were pulverized with a size of <80 mesh to obtain extractable OM (EOM) (i.e., saturated hydrocarbons, aromatic hydrocarbons and resins) by using a Soxhlet extractor system (SEM) with a solvent of dichloromethane (DCM). After the extraction, the addition of activated copper powder and anhydrous sodium sulfate, in turn, to the EOM achieved the removal of water and elemental sulfur, respectively.

The solvents were evaporated to near dryness at below 40°C and then removed into triangular bottles. Adding n-hexane with constant shaking, then performing ultrasonic dissolution. The asphaltene was entirely precipitated after a static settlement for more than 12 h. Then the material was removed to a short-necked funnel stuffed with skimmed cotton and washed with n-hexane until the filtrate was colorless to obtain the EOM which was distilled to 2–3 mL. The asphaltene remained in the short-necked funnel, cotton and triangular bottle was dissolved and washed with DCM until the filtrate was colorless. The EOM composition was eluted with n-pentane for aliphatic hydrocarbon compounds, a mixture solvent of n-pentane and DCM (2:3; V:V) for aromatic hydrocarbon compounds, and methanol for resin compounds, in turn, by using a silica gel (activation in an electric drying oven at 200°C for at least 4 h)-Al₂O₃ (activation in an electric drying oven at 400°C–450°C for at least 4 h) chromatography. The separated saturated and aromatic hydrocarbon components were volatilized to dry at the condition of <40°C, and the resin and asphaltene components were volatilized to dry at the condition of <60°C.

After the solvent in remnants was evaporated to dryness, the weight difference between the two weighing intervals of 30 min of <0.3 mg can be regarded as constant weight.

n-alkanes and isoprenoids in the saturated hydrocarbons were analyzed by GC-FID, while terpanes (m/z = 191; Figures 3A,C) and steranes (m/z = 217; Figures 3B,D) in saturated hydrocarbons, as well as aromatic hydrocarbons including methylphenanthrenes (MPs, m/z = 194; Figure 3E), trimethylphenanthrenes (TMPs, m/z = 220; Figure 3F), chrysene (m/z = 228; Figure 3G), and triaromatic steroid (TMPs, m/z = 231; Figure 3H), were performed by GC-MS.

The instrument of GC-FID was Shimadzu GC-2010 (Japan) equipped with HP-5MS chromatographic column (30 m × 0.25 mm × 0.25 µm). The carrier gas was 99.999% helium at a velocity of 1.0 mL/min. The inlet temperature was 300°C and the split ratio was 50:1. During the analysis, the initial temperature of the GC oven was 100°C for 1 min, then programmed to raise to 300°C at 4°C/min and hold for 25 min. The temperature of the FID was 300°C with hydrogen of 40 mL/min and air of 400 mL/min.

The instrument of GC-MS was Agilent 7890GC/5977MS (America) equipped with HP-5MS chromatographic column (the stationary phase consists of 95% poly (methylsilicone) and 5% poly (phenylmethylsilicone); 30 m × 0.25 mm × 0.25 µm). The carrier gas was 99.999% helium at a velocity of 1.0 mL/min. During the analysis, the initial temperature of the GC oven was 100°C for 1 min, then programmed to raise to 300°C at 4°C/min and hold for 20 min. The MS was performed in electron ionization (EI⁺) mode with an ion source temperature of 230°C and ionization energy of 70 eV.

The relative abundance of each compound was calculated from the peak area, and the identification of compounds was based on their relative retention time compared with literature data (e.g., Wu et al., 2022b; Zheng et al., 2022).

The measurement of δ¹³C_org was performed using a gas isotope ratio mass spectrometer (Finnigan MAT-252) equipped with a 30 m Porapak Q chromatographic column with 20 μm film thickness. The carrier gas was 99.999% helium at a velocity of 1.2 mL/min. During the analysis, the initial temperature was 40°C for 1 min, then programmed to raise to 160°C at 15°C/min. The δ¹³C_org value was calculated by comparing it to the intentional PDB (Pee Dee Belemnite) standard, and the analytical error was ±0.1‰.

The kerogen type for the examined samples has been interpreted based on the results from maceral composition, Rock-Eval pyrolysis, and atomic ratios of isolated kerogen (i.e., H/C, O/C and C/N) (Qiao et al., 2020). The kerogen type was generally interpreted as type III kerogen, although there are large differences based on the indicators mentioned above (cf. Figure 7 in Qiao et al., 2020). So, it is necessary to use more parameters to interpret the OM origins.

The stable carbon isotope (δ¹³C) compositions of the saturated hydrocarbon fraction (δ¹³C_Sat.) in the analyzed J₁b, J₁s and J₂x OM-rich shales vary from –28.95 to −29.97‰, −27.35‰ to −31.49‰, and −27.26‰ to −27.60‰, respectively; and of the aromatic fraction (δ¹³C_Aro.) from –25.47 to −28.46‰, −25.52‰ to −29.67‰, and −25.38‰ to −25.72‰, respectively. The δ¹³C_Sat. and δ¹³C_Aro. compositions of the investigated shales plotted on the Sofer (1984) diagram showing that all these extracts originated from shales that received predominantly terrestrial land plants (Figure 4). Another parameter, i.e., Canonical Variable (CV = − 2.53 × δ¹³C_Sat. + 2.22 × δ¹³C_Aro. − 11.65) (Sofer, 1984), can be used to discriminate OM from terrigenous and marine settings. Values of >0.47 indicate predominantly non-waxy terrigenous organic sources, whereas of <0.47 indicate more contribution from marine OM (Sofer, 1984; El Diasty et al., 2016). The calculated CV values for the studied samples ranging from −0.47 to −11.25 with 80% of >0.47 indicate that the OM was derived from terrestrial settings, and the samples below the threshold are likely to be affected by maturity (Collister and Wavrek, 1996) and caused by higher salinity mentioned later. Moreover, the CV values indicate the stratifications in water columns as well (Collister and Wavrek, 1996). The terrestrial OM can be further supported by the depleted δ¹³C values which are gradually increasing from the saturated hydrocarbon to the aromatic hydrocarbon, then to the resins ranging from −28.94‰ to −25.41‰ (avg. −26.68‰) and finally to the asphaltene ranging from −28.82‰ to −24.61‰ (avg. −26.31‰) (Galimov, 2006). The type III kerogen can be further supported by the relative concentrations of the asphaltenes and resins (from 3.2% to 92.4% and 39.4% on average, of EOM) compared to the concentrations of aliphatic and aromatic hydrocarbons (Figure 5).

Biomarkers, as indicators of parent sources of OM, have been widely used (Luo et al., 2016; Wu et al., 2022b; Zheng et al., 2022; Zhu et al., 2022). The distribution pattern of saturated hydrocarbons contains a lot of information on biological sources. n-alkanes with low carbon numbers (i.e., n-C₁₂–C₁₈) mainly originate from phytoplankton and algae (Meyers, 1997; Luo et al., 2016). Short-chain n-alkanes with relatively high concentrations of C₁₇ or C₁₈ typically originate from marine algae and cyanobacteria (Riboulleau et al., 2007; Sachse and Sachs, 2008). Middle-chain n-alkanes (n-C_21–25) are sourced from aquatic higher plants in common (Ficken et al., 2000), especially with peak carbon of n-C₂₃ and n-C₂₅ alkanes (Huang et al., 1999) being inductive of aquatic pollen taxa, Nymphaea (Coetzee, 1967), Sphagnum moss species (Nichols et al., 2006; Bingham et al., 2010), fresh-water algae (Riboulleau et al., 2007). Furthermore, long-chain n-alkanes with obvious odd carbon predominance are generally input of higher plants (Eglinton and Calvin, 1967) and the high concentrations of C₂₇, C₂₉, and C₃₁ n-alkanes are indicative of the land plant epicuticular waxes (Meyers, 1997). Based on the above indicative information, the parent material source of OM can be interpreted. The Paq values ranging from 0.36 to 0.90 (avg. 0.68) indicate aquatic macrophytes make an important contribution to OM (Ficken et al., 2000). The moderate to high WI values of 0.70–4.50 (avg. 1.91) in most samples indicate a substantial contribution of terrestrial OM. The CPI values of the investigated samples here are consistent with terrestrial OM input based on the consideration of the maturity stage. Additionally, the occurrence of branched alkanes indicates a contribution from bacteria to OM (Shiea et al., 1990). For example, the appearance of farnesane (i-C₁₅) indicates the input of green sulfur bacteria (Summons and Powell, 1987). The concentrations of branched alkanes (e.g., i-C₁₅/n-C₁₃ are 0–2.57 with an average value of 0.33; i-C₁₆/n-C₁₄ are 0–2.52 with an average value of 0.52; i-C₁₇/n-C₁₅ are 0–2.89 with an average value of 0.24) indicate the contribution from bacterial was abundant.

However, it must be noticed that these results drawn from the distribution of n-alkanes distributions should be used with caution, because the longer the carbon chain, the easier it is to convert to shorter chains with increasing maturity and thus change its distribution pattern (Peters et al., 2005) and the biological source signal based on n-alkanes is not exclusive (Oros et al., 1999; Gonzales et al., 2020).

Steranes are mainly derived from higher plants (e.g., a high concentration of C₂₉ steranes) and algae (e.g., a predominance of C₂₇ steranes) (Huang and Meinschein 1979; Moldowan et al., 1986; Volkman 2003), and hopane sources are mainly from aerobic bacterial membranes (Ourisson and Rohmer, 1992; Peters et al., 2005). The low steranes/hopanes ratio ranging between 0.07 and 2.98 with an average value of 0.37 indicates a substantial input of bacterial biomass.

The ternary diagram of C₂₇, C₂₈ and C₂₉ ααα 20R steranes distributions (Figure 6; Huang and Meinschein 1979) indicates that the OM of the J₂x Formation sourced mainly from higher plants was deposited in a terrestrial environment, while the OM of the J₁ formations (i.e., J₁b and J₁s formations) was a mixed OM derived from higher plants, bacteria and algae deposited under bay/estuary environments alternated with a terrestrial environment (Figure 6). Moreover, the presences of abundant TMPs and chrysene (Figure 3F), etc. are typically indicative of terrigenous higher plant input (Chaffee and Johns 1983; Chaffee and Fookes 1988).

Based on the lithology assemblage (Figure 2), a typical lacustrine sedimental environment can be recognized. The lithology enriched in clay minerals was deposited in a non-restricted setting, which can be demonstrated by the high C₂₇ diasteranes abundances compared to regular steranes, but low C_21-22 steranes abundances compared to C_27-29 steranes (Figure 7) (Wang et al., 2015; Qiao et al., 2021a; Zheng et al., 2022). Clay-rich lithology is further supported by the higher C₃₀ hopane abundance compared to the lower C₂₉ hopane abundance (Figure 3) with the C₂₉/C₃₀ hopane of 0.41–1.50 (0.68 on average) (Gürgey, 1999). This is consistent with the distribution of gradually regular decrease from C₃₁ to C₃₅ hopanes indicating clastic facies (Figure 3; Waseda and Nishita, 1998), the C₂₆/C₂₅ Tri (tricyclic terpane) (0.81–3.30 with an average value of 1.89) and C₃₄/C₃₅ hopane (1.02–4.35) ratios, low C₂₂/C₂₁ (0–1.08 with an average value of 0.25) and C₂₄/C₂₃ Tri ratios (0–6.21 with an average value of 0.63), as well as low HHI values (0–0.10) (Peters and Moldowan, 1991; Sinninghe Damsté et al., 1995; Peters et al., 2005) (Figure 8). Non-marine setting can be supported by the high ratios of C₃₀ hopane/(20R steranes + C₃₀ hopane) ranging between 0.18 and 0.83 (avg. 0.66) (Holba et al., 2003). The lacustrine depositional environment can be also supported by the low steranes abundances compared to hopanes in all formations (0–0.95 with an average value of 0.06) (Peters et al., 2005). Considering that steranes abundances would also be high in lacustrine settings which developed plenty of biomass inputs, the low steranes abundances in the studied area are consistent with the OM compositions discussed above.

However, the parameter of C₃₁ 22R/C₃₀ hopane ratios (0–0.52 with an average value of 0.25) (Figure 8B), which is inductive of marine situations at the situation of ratio >0.25 and lacustrine environments at of <0.25 (Peters et al., 2005), is inconsistent with the above discussion for the analyzed samples, which might be due to the paleosalinity conditions.

Thermal maturity is an essential parameter for shale, which determines whether shale resources are available and what type of shale resources they are (e.g., shale-oil and shale-gas), and shows impacts on the diagenesis and porosity (Liu et al., 2022).

Some molecular parameters are widely used to evaluate thermal maturity (Holba et al., 2003; Volk et al., 2005; Zhu et al., 2012). The R configuration on the side chain of some compounds can be converted to S configuration with the increase of thermal maturity, and finally reached an equilibrium mixture of S and R configurations, thereby being regarded as maturity-dependent ratios, e.g., C₃₁ and C₃₂ homohopanes 22S/(22S + 22R) ratios as well as C₂₉ ααα 20S/(20S + 20R) steranes. Moreover, the stability of a flat three-dimensional configuration (i.e., ααα configuration) is poor, which will transit to a more thermodynamically stable αββ configuration until ααα and αββ configurations reach an equilibrium ratio with the increase of maturity. So, C₂₉ ββ/(ββ + αα) steranes can be regarded as a maturity-dependent parameter. Molecular maturity-related parameters are easily affected by conditions of sedimentary water columns. For example, ratios of 22S/(22S + 22R) hopanes and ββ/(ββ + αα) steranes show a higher maturity stage in immature extracts from hypersaline rocks (ten Haven et al., 1986; Rullkötter et al., 1994). This discrepancy can be observed in the Pleistocene hypersaline sediments in the Qaidam Basin, i.e., the obviously too high hopane and steranes maturity-dependent ratios in the definitely immature sediments (Qiao et al., 2021b). Moreover, sterane isomerization ratios are also affected by biodegradation and weathering (Peters et al., 2005). For the analyzed Jurassic shales, there is no clear weathering and biodegradation occurred. Moreover, the salinity of the water columns showed large variation, but it did not develop hypersaline conditions, so the molecular parameters are effective in the study area. A comparison of C₃₁ and C₃₂ homohopanes 22S/(22S + 22R) ratios being 0.45–0.63 and 0.29–0.59, respectively, indicate a thermal maturity between the early and main oil window (Figure 10A; Peters et al., 2005). The C₂₉ ααα 20S/(20S + 20R) (0.23–0.50) vs ββ/(ββ + αα) (0.08–0.64) steranes plot shows similar maturity conditions (Figure 10B; Peters et al., 2005). It should be noted that two samples from J₂x samples are immature based on Figure 10B, which might be because interference by co-eluting peaks can make the homohopane parameters invalid (Peterset al., 2005). The plots of C₃₁ homohopanes 22S/(22S + 22R) ratio, C₂₉ ααα 20S/(20S + 20R) steranes and C₂₉ ββ/(ββ + αα) steranes vs depth (Figure 11) indicate that all the samples’ thermal maturity has entered oil windows and the molecular parameters appear to have reached the equilibrium value at around 3,500 m. Furthermore, the thermal maturity is further supported by the CPI (carbon preference index; Bray and Evans 1961) and OEP (odd-to-even predominance; Peters and Moldowan 1993) values. These two parameters are easily affected by the origin of OM because there is no clear odd-to-even predominance in the middle and long chains in the marine algae/saltwater algae. Considering the composition of the OM in the studied samples, the CPI of 1.01–1.43 and OEP of 0.83–1.38 are valid maturity parameters, indicating that most samples have entered into the oil window. The maturity that has entered the oil window is further confirmed by the high saturate/aromatic ratios (0.11–21.84) (Figure 5; Peters et al., 2005) and low to middle percentages of the resins compared to the saturated and aromatic hydrocarbons (from 2.3% to 46.5%, of EOM) (Ayinla et al., 2017). A similar maturity is interpreted by the MPs distribution (Radke, 1983) and the relative abundance of C₂₆ and C₂₈ 20S triaromatic steroid (Thompson-Butler et al., 2019) (Figures 3E,H). The above results are consistent with the thermal maturity ranging from low mature and mature supported by Rock-Eval T_max (426°C–468°C) and vitrinite reflectance (0.47–1.59%VR_r) values released by Qiao et al. (2020).

Generally, the high PI values are caused by the early generation of hetero-compounds at the lower maturity (Curiale 1986), which are decreasing with the increase of thermal maturity (Lewan 1994). For the studied samples, a small number of samples have high PI values of > 0.4 with relatively lower T_max values (Qiao et al., 2020), which are located in the area of the poorly-drained system where hydrocarbons can be impregnated in the coarser particle grained laminae in the shale (Jarvie 2012; Li et al., 2018). The relationship between PI and T_max values (cf. Figure 4B in Qiao et al., 2020) indicates most analyzed samples are well-drained source rocks which expelled generated hydrocarbons efficiently to the various types of coarser grained siltstone closed intervals (Figure 2) (Raji et al., 2015). This assemblage of shale and siltstone is more typical of the hydrocarbon resources in the Chang 7 Member where coarser intervals (i.e., siltstone/muddy-siltstone) occurred in the direct vicinity within the shaly intervals act as tight-oil reservoirs and the shaly intervals are important shale-oil reservoirs (Qiao et al., 2021a). The economic hydrocarbon potential for the tight-oil reservoirs and shale-oil reservoirs can be further supported by the high oil saturation and the wide range of porosity from the EF 2 well (Figure 2).

Kaolinite formed at temperatures higher than 15°C is commonly found in well-drained soils with high precipitation under humid tropical climates (Egger et al., 2002; John et al., 2012), whereas illite is generally formed in climates with very low weathering rates, such as cold regions and desert areas. Although sea level change and rainfall may also affect kaolinite deposition (e.g., Gibson et al., 2000), climate change is considered to be the primary control for clay minerals composition due to the lack of correlation between sea level change and kaolinite content (e.g., Robert and Chamley, 1991; Bolle and Adatte, 2001). Besides climate conditions, burial diagenesis affects clay mineral composition. Conversion occurs between different clay minerals due to the increase in formation pressure and geotemperature as well as the release of water between clay mineral layers and the migration of interlayer cations. Under shallow burial conditions, kaolinite and smectite exist in clay minerals, but these minerals disappear and are converted to illite and chlorite under deep burial conditions. Considering that the thermal maturity of the examined samples corresponds to the early stage of the intermediate diagenetic stage, the content of kaolinite is much greater than that of illite (Figure 12A), which is sufficient to indicate that the paleoclimate conditions of the study area were warm and humid, although the transition to illite has happened and the transition to chlorite has just begun at this stage as well as how much of these transitions have occurred is uncertain. Moreover, the kaolinite/illite ratio is significantly lower in the J₂x samples compared to that in the J₁ samples (Figure 12B), implying that the paleoclimate conditions were colder and drier during the J₂x period, which is consistent with the results based on elemental Sr/Cu ratios (cf. Figure 6 in Qiao et al., 2020). This paleoclimatic change is also supported by the lithologic assemblages, i.e., the coal seam is mainly distributed in the J₁b Formation and the lower part of the J₂x Formation (Figure 2). The above inferences are consistent with the interglacial period that occurred during the Jurassic in NW China (Sellwood and Valdes, 1997). The same results were supported by sedimentologic, paleocurrent, and subsidence analyses (Hendrix et al., 1992), lithology characteristics and palynology (Ashraf et al., 2010), as well as sequence stratigraphy characteristics and fossil vertebrates (Eberth et al., 2001; Li et al., 2014).

The change of paleoclimate in the formation of shale is mainly reflected in the water environment and biological composition. The colder and drier paleoclimate conditions may result in a smaller scale of the lake basin and shallower water depth during the J₂x period. This resulted in the oxic benthic water column condition (Figure 9B and discussion in Section 4.2.2) and the proportion of terrestrial higher plants in the OM composition increased (Figure 6 and discussion in Section 4.1) during the J₂x period in comparison to those during the J₁ period. Another important indicator showing the J₂x period was cooler than the J₁ period is β-Carotane which occurs only in the J₂x Formation (Jiang and Fowler, 1986; Fu et al., 1990; Koopmans et al., 1997) although β-Carotane is not very common.

The humid and warm climate conditions during the J₁b, J₁s and J₂x periods of the study area were conducive to the development of the terrestrial higher plant but hard to form brackish water conditions. Even though the studied area became colder and drier during the J₂x period, the change was not significant enough to cause the change of salinity in the water column and the vanish of terrestrial higher plants during the corresponding period (see Section 4.4 and cf; Qiao et al., 2020). Similar paleodepositional conditions can be observed in the Triassic Ordos Basin which developed a mixed OM composition and oxic to suboxic but freshwater conditions (Qiao et al., 2021a). Moreover, it is noteworthy that these salinities did not vary regularly in space and time, showing that they were not affected by climate change, so this variation was more likely due to seawater intrusion as proposed by Sha et al. (2011).

Generally, the scale and enrichment degree of shales are controlled by the input and preservation condition of OM, which includes multiple factors, such as climate, paleontology growth, subsidence rate, as well as level and maturity of lake (Shen et al., 2015; Liu et al., 2021). All of these can be categorized into two groups, i.e., paleoproductivity and depositional conditions (Katz, 2001).

Based on the elemental evaluation proposed by Qiao et al. (2020), the investigated study area developed middle to high productivity. Moreover, the paleoproductivity during the J₁ period was higher than that during the J₂x period based on Ba concentrations (cf. Figure 6 in Qiao et al., 2020), but the TOC contents show similar characteristics vertically in all formations (Figure 2). This is because the Ba concentration only indicates the level of organic carbon fluxes in water columns, while the OM of the J₂x is sourced mainly from terrestrial OM and the redox conditions were oxygen-rich conditions which were not prone to the OM preservation. Based on the relationship between the δ¹³C_org and TOC content, the OM accumulation was controlled by paleaoproductivity in the EF 2 well which is located at the margin of the sag, whereas the redox conditions exerted a primary control on the TOC contents in the samples from the EF 7 well located at the more central area of the sag (cf. Figure 13 in Qiao et al., 2020). This is because the basin margin was more prone to obtain OM from terrestrial higher plants, but the water was shallower and more oxidized due to a reduction in the size of the lake basin caused by cooler and drier paleoclimate conditions, which experienced strong oxygenolysis and were not in favor of the OM preservation. Although the substantial OM input from terrestrial inputs can promote the palaeoproductivity in the sag margin, this input was discontinuous resulting in the large variation of TOC content in the EF 2 well. Therefore, the large variations in the biotic palaeoproductivity from the terrestrial organisms were the material basis controlling the development of the organic-rich shales. By contrast, the redox conditions were more important for controlling OM presentation in the central area of the sag which was far from the terrestrial OM source with a small variation in TOC content (cf. Figure 13 in Qiao et al., 2020). In the central region of the basin, the paleoredox conditions of the benthic water column were deeper and more reductive and were less affected by the reduction of lake scale caused by the paleoclimate change. Considering the accumulation model deduced by Qiao et al. (2020), the paleoredox and salinity conditions as well as OM composition obtained in this study are more reasonable and supply more evidence.