The marine C₂₀-C₂₁-C₂₃TT data were compiled from 20 basins in 11 countries (Table 1). Their C₂₁/C₂₀TT and C₂₃/C₂₁TT ratios vary from 0.98 to 2.83 and 1.20 to 3.57, respectively (Figure 2). The relative abundances of C₂₀-C₂₁-C₂₃TTs display a pattern of C₂₀<C₂₁<C₂₃TT (e.g., Figure 3A).

Saline lacustrine facies is represented by the source rocks from the Junggar Basin in Northwestern China (Bian et al., 2010; Yu et al., 2017) and the oil samples from the Bohai Bay Basin in Eastern China (Table 1). Their C₂₁/C₂₀TT and C₂₃/C₂₁TT ratios vary from 1.00 to 2.08 and 1.15 to 1.52, respectively (Figure 2), which slightly overlaps with samples classified as marine in origin (Zhu and Jin, 2003; Yu et al., 2017). The relative abundances of C₂₀-C₂₁-C₂₃TTs associated with saline lacustrine facies are similar to those of marine facies.

Brackish-freshwater lacustrine facies is represented by the samples from the Muglad Basin (Sudan) (Xiao et al., 2019b), Bohai Bay Basin (Lv and Thesis, 2019) and Ordos Basin in Central China (this study, Table 1), which were deposited in semi-deep to deep brackish-freshwater lacustrine facies (Zhang et al., 2009; Xiao et al., 2019a; Ma et al., 2019). The C₂₁/C₂₀TT and C₂₃/C₂₁TT ratios fall into the ranges of 0.95–3.03 and 0.36 to 1.34, respectively (Figure 2). The relative abundances of C₂₀-C₂₁-C₂₃TTs display two patterns: C₂₀<C₂₁<C₂₃TT and C₂₀< C₂₁>C₂₁TT (e.g., Figure 3B).

Terrigenous organic matter can be transported by rivers and eventually deposited in lacustrine or marine environments. Terrigenous organic matter transported by rivers has been reported to be deposited in marine facies in the northern South China Sea (Li et al., 2011; Deng et al., 2019) and lacustrine facies in the Junggar Basin (Gao et al., 2017). On this condition, the C₂₀-C₂₁-C₂₃TT data are strongly controlled by the source instead of depositional environment. The C₂₀-C₂₁-C₂₃TT data of terrigenous source in the Junggar Basin give C₂₁/C₂₀TT ratios of 0.50–0.91, and C₂₃/C₂₁TT ratios of 0.25–0.53 (Figure 2), with a relative abundance of C₂₀>C₂₁>C₂₃TT (e.g., Figure 3C).

TTs have drawn broad attention in literature since Anders and Robinson (1971) and Gallegos (1971) described the lower homologs (C₁₉-C₂₄) of TTs in the bitumen and oils of the Green River Shale (e.g., Aquino Neto et al., 1982; Chicarelli et al., 1988; De Grande et al., 1993; Dutta et al., 2006; Tao et al., 2015; Philp et al., 2021). Multiple compound sources/precursors for TTs have been proposed, such as prokaryotic cell membrane (e.g., Ourisson et al., 1982), diterpenes in terrigenous plants (e.g., Ekweozor et al., 1983) and the now-extinct algal-like Tasmanites (e.g., Aquino Neto et al., 1982; Chicarelli et al., 1988; De Grande et al., 1993; Dutta et al., 2006). So far, no clear precursor-product relationship has been established, which has at least partially impeded the direct application of TTs as source indicators (Philp et al., 2021).

Here, the analysis of a large amount of C₂₀-C₂₁-C₂₃TT data from various depositional environments provides some insights into the origin of the observed C₂₀-C₂₁-C₂₃TT abundance patterns. Previous studies have shown that the organic matter in typical marine and shallow-water terrestrial facies mainly originated from plankton and terrestrial plants (e.g., Tissot and Welte, 1978). Accordingly, the C₂₀<C₂₁<C₂₃TT pattern likely corresponds to a dominant plankton input, while the C₂₀>C₂₁>C₂₃TT pattern may correspond to a dominant contribution of terrigenous plants. The former correspondence can be evidenced by the Neoproterozoic oil shale (900–873 Ma; maturity: 0.6–0.7%Ro) in North China, in which the terrigenous plant input can be ignored (Zhang et al., 2007). This oil shale contains cyanobacteria and green algae, and shows a C₂₀<C₂₁<C₂₃TT abundance pattern in the rock extracts, suggesting that planktons are the biological source for the TTs with C₂₀<C₂₁<C₂₃TT abundance pattern.

The analysis also indicates that neither plankton nor terrigenous plants alone can generate the TTs with C₂₀>C₂₁<C₂₃TT or C₂₀<C₂₁>C₂₃TT patterns, which have been observed in marine-continental transitional and brackish-freshwater lacustrine facies, respectively. As these two depositional environments commonly receive blended input of plankton and terrigenous plants, mixing from different sources is likely the cause of the C₂₀>C₂₁<C₂₃TT and C₂₀<C₂₁>C₂₃TT patterns. This hypothesis is well supported by the numerical modeling of a mixing between plankton and terrigenous plants at various mixing ratios (Figure 4). Representative C₂₀-C₂₁-C₂₃TT abundance patterns in the terrigenous plants domain (i.e., Point A in Figure 4A) and the plankton domain (i.e., Points B or C in Figure 4A) determined by this study, were used as end-members for the mixing modeling.

As illustrated in Figure 4, with the change of mixing ratios between A and B, the C₂₀-C₂₁-C₂₃TT abundance pattern changes from C₂₀>C₂₁>C₂₃TT to C₂₀>C₂₁<C₂₃TT (e.g., D in Figures 4A, B) when the mixing ratio (expressed as A%/B% ratio) decreases from 100/0 to 75/25, and further to C₂₀<C₂₁<C₂₃TT when the mixing ratio is lower than 44/56 (Figure 4B). Similarly, mixing between A and C can generate the C₂₀<C₂₁>C₂₃TT pattern (e.g., E in Figures 4A, C) when the mixing ratios (expressed as A%/C% ratio) are in the range of 81/19–42/58 (Figure 4C). Notably, almost all C₂₀-C₂₁-C₂₃TT data in this study fall into a range from the terrigenous plants domain to the plankton domain (Figure 5), suggesting that C₂₀-C₂₁-C₂₃TT abundance patterns are strongly controlled by the two end-member biological sources and their mixing contributions. Thus, the input mixing of plankton and terrigenous plants at different proportions should be responsible for the formation of various C₂₀-C₂₁-C₂₃TT abundance patterns.

Depositional environments control the input proportions of plankton and terrigenous plants, which is proposed to control the C₂₀-C₂₁-C₂₃TT abundance patterns in sedimentary rocks. The analysis of available C₂₀-C₂₁-C₂₃TT data from known depositional environments clearly show that the C₂₀-C₂₁-C₂₃TT abundance patterns in different depositional environments are distinct from each other, with C₂₀<C₂₁<C₂₃TT in typical marine and saline lacustrine facies, C₂₀<C₂₁>C₂₃TT in freshwater lacustrine facies, C₂₀>C₂₁>C₂₃TT in river-lake transitional facies, and C₂₀>C₂₁<C₂₃TT in marine-continental transitional facies. As illustrated in Figure 2, the C₂₃/C₂₁TT ratio not only increases progressively from freshwater lacustrine, to saline lacustrine, to marine facies, but also increases progressively from terrigenous source, to river-lake transitional, to marine-continental transitional facies. The C₂₃/C₂₁TT ratio appears to increase gradually with elevated salinity of depositional water. The increase in salinity usually indicates a decrease in fresh-water input and therefore a decrease in terrigenous input, which results in the C₂₃/C₂₁TT ratios approaching their source signatures of halophilic plankton. Accordingly, the boundary between brackish and freshwater lacustrine facies is expected to lie around a C₂₃/C₂₁TT ratio of 1.0 (Figure 2). Furthermore, the C₂₁/C₂₀TT ratios in marine and lacustrine facies are obviously greater than those in transitional facies (Figure 2), suggesting that the C₂₁/C₂₀TT ratio seems to increase with elevated depths of depositional water. In general, the greater the depositional depth, the farther offshore, and the less terrigenous input. Thus, their C₂₁/C₂₀TT ratios approach the signatures of planktons (halophilic plankton or freshwater plankton).

Based on the good correspondences between C₂₀-C₂₁-C₂₃TT abundance patterns and typical depositional environments, the cross-plot of C₂₃/C₂₁TT vs. C₂₁/C₂₀TT in context of depositional environments (Figure 2) can be used as a discriminating diagram for environmental identification. C₂₃/C₂₁TT and C₂₁/C₂₀TT are expected to be the parameters to assess water salinity and depositional depth, respectively.

The Ordos Basin undergone (1) an early Paleozoic shallow oceanic-platform stage and 2) a late Paleozoic offshore-plain stage during the Paleozoic (Li, 2004). During the Middle Ordovician, the Majiagou Formation was deposited within a semi–closed epicontinental sea environment (Li et al., 2018). From the Late Ordovician to the Early Carboniferous, the basin was uplifted by the Caledonian orogeny and underwent 130 million years of erosion (Yang et al., 2012; Xu et al., 2018). The following Hercynian orogeny caused the Late Paleozoic Ordos Basin subsidence. Large scale of transgression occurred from the east and west of the basin during the Benxi period and the seawater connected together during the Late Taiyuan period. With the gradual regression during the Shanxi period, continental deposition began to dominate, resulting in the development of marine-continental transitional facies, delta facies and lacustrine facies (Sun, 2017).

Thirteen core samples from the Ordovician-Permian in the Ordos Basin (Figure 6; Table 2) and ten coal or carbonaceous mudstones from the Eocene in the East China Sea Basin (Table 3) were used to verify the validity of C₂₀-C₂₁-C₂₃TT biomarker method.

The C₂₀-C₂₁-C₂₃TT biomarker method only requires a small amount of rock extracts, depending on the lower limit of GC-MS analysis. At present, the rock extracts more than 10 μg is guaranteed to obtain high-quality m/z191 mass chromatogram. Take the mudstone samples from the Pingbei area for example, the TOC and S₁+S₂ values of mudstone samples vary from 0.23% to 3.68% and from 0.20 mg/g to 6.03 mg/g (Table 3), which were classified as “poor” to “fair” level source rocks. Although the chloroform asphalt ‘A’ extracted from the mudstones are as low as 0.01% (Table 3), the high-quality m/z 191 mass chromatogram with C₂₀-C₂₁-C₂₃TT signature has been obtained from every mudstone sample (e.g., Figure 1). This indicates that our method is not necessarily restricted to hydrocarbon source rocks, but can be applied to a broad spectrum of rocks.

TTs are characterized by higher thermal stability than hopanes and steroterpenes (Peters et al., 1990; Farrimond et al., 1999; Xiao et al., 2019a). Thermal evolution during maturation and high-maturation stages makes little effect on the abundance patterns of C₂₀-C₂₁-C₂₃TT (Chen et al., 2017; Xiao et al., 2019b). As shown in Table 2, although the rocks have evolved into the maturities ranging from 1.6% to 2.4%Ro, C₂₀-C₂₁-C₂₃TT biomarker method are still effective in environmental identification. Furthermore, C₂₀-C₂₁-C₂₃TT distributions are stable even in the severely biodegraded oils with 25-norhopane series, showing strong resistance to biodegradation (Xiao et al., 2019a).

The Pingbei area was a hinged margin of a rift basin during the Eocene (Soreghan and Cohen, 1996). The petroleum system in the area is within the upper-middle Eocene Pinghu Formation, which consists of alternate sandstone, mudstone and thin coal seams. From bottom to top, the Pinghu Formation is further divided into the P1, P2, P3 and P4 members (Figure 11A).

Due to the scarcity of wells and the lack of typical facies indicators in drill core, the depositional environments of the Pinghu Formation have been strongly debated among three possibilities: delta environment (e.g., Shen et al., 2016; Jiang et al., 2020), tidal flat environment (e.g., Zhao et al., 2008) and mixed delta and tidal flat environment (e.g., Zhang et al., 2017; Abbas et al., 2018).

The Ro values in the Pinghu Formation in the area vary from 0.51% to 0.87% (Table 3), indicating a low-mature to mature stage. The C₂₀-C₂₁-C₂₃ TT abundance patterns cannot be affected by this maturity range, and are thus available to identify depositional environments. The C₂₁/C₂₀TT and C₂₃/C₂₁TT ratios were plotted in the discriminating diagram to identify the depositional environment of each member in the Pinghu Formation (Figure 12). Paleogeomorphologic map (Figure 11B) and seismic inversion data (Figures 11C–F) are also reported here as additional evidence to (1) support the environmental identification by the C₂₀-C₂₁-C₂₃TT biomarker method, and (2) help characterize the spatial distribution of depositional environment (Figures 12B, D, F, H).