Sage Journals: Discover world-class research

Abstract

The formation environment and preservation conditions of sedimentary organic matter (OM) play an important role in the accumulation of shale gas. In the present study, inorganic and organic geochemical data were analyzed to determine the origin and preservation environment of sedimentary OM in the Wc-1 well of the Wufeng–Longmaxi (WF–LMX) Formation in northeastern Chongqing, China. In a biomarkers analysis, the numerical characteristics of n-alkanes (n-C₁₇/n-C₃₁>4.0), tricyclic terpenes (C₂₃TT/C₃₀H>1.0), and steranes (C₂₇/C₂₉St>1.0) suggested that the main origin of OM in the black shale was planktonic algae. High values of P/Ti and Ba_XS in the paleoproductivity indices suggested that primary productivity in the WF–LMX Formation was relatively high, peaking in the lower LMX Formation. Relative enrichment in U, V, and Mo, and the changing trends in V/(V+Ni) and Ni/Co suggested that the redox conditions for the bottom water, which changed from the WF Formation to the lower and upper LMX Formation, were oxic/dysoxic to anoxic and dysoxic, respectively. The relationship between total organic carbon and the above indexes indicates that different key factors controlled OM enrichment in the WF–LMX Formation. In the WF Formation, oxic bottom water was not conducive to the preservation of sedimentary OM. In the lower LMX Formation, the highest paleoproductivity and anoxic bottom water conditions promoted the enrichment and preservation of sedimentary OM. In the upper LMX Formation, excessive terrigenous inputs and relatively low paleoproductivity limited the enrichment of sedimentary OM.

Keywords

Biomarkers paleoproductivity redox conditions organic matter Wufeng-Longmaxi formation

Introduction

As one of the largest shale gas basins in China, the Sichuan Basin has undergone substantial exploration and development in recent years (Bao et al., 2020; Chen, 2016; Chen et al., 2014; Gao et al., 2012; Guo and Zeng, 2015; Liu et al., 2019). Northeastern Chongqing lies within the Dabashan fault zone at the northeastern margin of the Sichuan Basin, where the geological structure is complicated by a multistage geological history (Li et al., 2007). Some exploratory wells in this area have produced shale gas, such as the Wuxi-2 well in the Ordovician–Silurian Wufeng–Longmaxi (WF–LMX) Formation and the Chatang-1 well in the lower Cambrian Shuijingtuo Formation. These formations have shale field desorption gas contents of ∼8 m³/t, which indicate considerable shale gas exploration potential in this area (Liang et al., 2016; Wang et al., 2015). At present, research on the black shale of the WF-LMX Formation in northeastern Chongqing has focused on determining its sedimentary structure, pore structure characteristics, and shale gas resource (Chen et al., 2016; Dong et al., 2011; Wang et al., 2015). Therefore, the origin of sedimentary organic matter (OM), controlling factors, and the characteristics of the sedimentary environment in this area require further study.

The enrichment and preservation of sedimentary OM is a complex physical and chemical process, which is influenced by the primary productivity, redox state of the water column, OM deposition rate, and other sedimentary processes (Chen et al., 2016; Demaison and Moore, 1980; Hatch and Leventhal, 1992; Lash and Blood, 2014; Murphy et al., 2000). Given that these conditions cannot be explained by a single factor, we must consider multiple factors to reproduce the sedimentary environment and key controlling factors (Rimmer, 2004; Tribovillard et al., 2006).

In the present study, the geochemical and biomarker characteristics of the WF–LMX Formation samples obtained from the Wc-1 well were comprehensively analyzed and used, along with data on tectonic development, to determine the origin and controlling factors of sedimentary OM. The findings of this study will provide guidance for subsequent shale gas exploration.

Geological setting

The study area is located in Chengkou County, Chongqing City, which borders Sichuan, Shaanxi and Hubei provinces in China. The geotectonic setting is the junction of the Sichuan Mesozoic Foreland basin and the southern passive margin fold-thrust belt of the upper Yangtze plate. The main area is located to the north of the Shashi buried fault and south of the Chengba fault, including the Dabashan platform margin depression and part of the Dabashan fold area, as shown in Figure 1.

Figure 1.

Geodetic location and stratigraphic column of the Wc-1 well.

The Wc-1 well is located in the western part of the study area. The well experienced a series of considerable environmental and biological disturbances during the transition from the Ordovician to Silurian, including the long-term cooling trend in the Katian stage, the glaciation of the Hirnantian, and a biological extinction event (Algeo et al., 2016; Brenchley et al., 2003; Fan et al., 2009; Finnegan et al., 2011; Harper et al., 2014). According to field investigations and total organic matter (TOC) content, this transition period can be divided into three units. During the Upper Ordovician period, black shale and siliceous rock of the WF Formation were deposited, with a thickness of ∼10 m. As the climate warmed, biological recovery, and uplift of ancient land occurred in the Early Silurian period. The sedimentary environment of the LMX Formation was characterized by gentle continental shelf facies deposition (Chen et al., 2014), in which shale and sandstone with a thickness of ∼100 m were deposited according to the different organic carbon contents and lithologies. The lower unit in the LMX Formation is black graphite shale with developed horizontal bedding, which has a high organic carbon content and thickness of ∼ 45 m and is characterized as deep-water shelf deposition. The upper unit in the LMX Formation is a light–yellow–green shale, gray silty mudstone and argillaceous siltstone with a thickness of ∼55 m, which is characterized as shallow–water shelf deposition (Figure 1).

Samples and methods

Samples

Forty samples were selected from the WF–LMX Formation from the Wc-1 well for analysis of TOC and major and trace elements. Among the 40 samples, 13 were from the upper LMX Formation, 20 were from the lower LMX Formation, and 7 were from the WF Formation. Furthermore, 18 of the 40 samples were tested for biomarkers: 6 from the upper LMX Formation, 10 from the lower LMX Formation, and 2 from the WF Formation. Before testing, all samples were washed, dried, stored separately in paper bags, and wrapped in polythene bags to ensure minimal contamination.

Analytical methods

The TOC contents of the 40 samples, which were ground to 200 mesh, were analyzed using a CS-330 Carbon-Sulfur Analyzer (LECO Corporation, USA) at the Jiangsu Geological and Mineral Resources Design and Research Institute, with an accuracy of 95%. Major elements (e.g., SiO₂, Al₂O₃, and MgO) were analyzed by BRUKER S8 TIGER X-ray fluorescence spectrometry (XRF, AXS Corporation, Germany) at the China University of Mining and Technology Advanced Analysis and Computation Center. The analytical procedures for TOC and major elements were performed according to the Chinese national standards GB/T19145-2003(2003) and GB/T14506.28–2010, respectively.

Trace elements (e.g., U, V, and Mo) were analyzed using a PE Elan6000 standard inductively coupled plasma mass spectrometer (ICP-MS) at the Jiangsu Geological and Mineral Resources Design and Research Institute, with an accuracy of 95%. The analytical procedures for trace elements were performed according to the Chinese national standards GB/T 14506.30–2010.

The saturated hydrocarbons of the 18 samples, which were ground to 100 mesh, were analyzed by gas chromatography-mass spectrometry (GC-MS, Shimadzu, Japan) at the Key Laboratory of CBM Resources and Dynamic Accumulation Process, China University of Mining and Technology to identify biomarkers, including normal alkanes (n-alkanes), isoprenoids, terpanes, and steranes. The analytical procedure for the biomarkers was performed according to the Chinese national standard GB/T 18606–2017.

Data presentation

The degree of trace element enrichment in the sediments is controlled by the number of their authigenic components (Liu et al., 2019), and the enrichment factor (EF) is often used for evaluation (Tribovillard et al., 2006). The calculation formula is as follows:

X_{EF} = {(X / Al)}_{sample} / {(X / Al)}_{AUCC}

(1)

Data for the average upper continental crust (AUCC) were proposed by McLennan (McLennan, 2001). When X_EF > 1, X (trace element) is enriched relative to AUCC, and vice versa.

The chemical index of alteration (CIA) of the source region is widely used in the reconstruction of paleoclimatic changes for the quantitative characterization of the degree of chemical weathering (Nesbitt and Young, 1982). The calculation formula is as follows:

CIA = {Al}_{2} O_{3} / ({Al}_{2} O_{3} + Ca O^{*} + {Na}_{2} O + K_{2} O) \times 100

(2)

where the unit of each element is the mole fraction and CaO* only represents the silicate mineral content (Nesbitt and Young, 1982). According to the method proposed by Bock et al. (1998), CaO* is calibrated as follows: when CaO > Na₂O, CaO* = Na₂O; when CaO ≤ Na₂O, CaO* = CaO.

Barium (Ba) is often used as an indicator of paleoproductivity (Algeo and Maynard, 2004). To eliminate the errors caused by the input of terrigenous debris, Ba is modified to Ba_XS and Ba_XS, which is used to assess the paleoproductivity level (Algeo and Maynard, 2004). The calculation formula is as follows:

{Ba}_{XS} = {Ba}_{sample} - {Al}_{sample} \times {(Ba / Al)}_{PASS}

(3)

where

{(Ba / Al)}_{PASS} = 0.0077

is the post-Archean Australian shale content (Taylor and McClcnnan, 1985).

The source of siliceous matter can be divided into terrigenous clastic, hydrothermal, and biogenic origins (Bostrom et al., 1973; Liu et al., 2017). The excess siliceous mineral content (Si_ex) refers to siliceous minerals other than normal continental clastic deposits, and its calculation formula is as follows:

{Si}_{ex} = {Si}_{sample} - {(Si / Al)}_{background} \times {Al}_{sample}

(4)

where (Si/Al)_background = 3.11, which is based on the relationship between major elements (Wedepohl, 1971).

Results

TOC abundance

Among the Wc-1 well samples, shales from the WF Formation have low TOC content ranging from 0.94% to 2.18% (mean = 1.54%), shales from the lower LMX Formation have relatively high TOC contents ranging from 3.26% to 6.89% (mean = 5.54%), and shales from the upper LMX Formation have TOC contents ranging from 1.15% to 4.45% (mean = 3.24%). This change in TOC content is shown in Figure 1 and Table 1.

Table 1.

The major elements of Wc-1 well samples.

Formation	Depth m	TOC %	Major elements (%)							CIA	Si_ex %
Formation	Depth m	TOC %	SiO₂	Al₂O₃	CaO	Na₂O	K₂O	MgO	TiO₂	CIA	Si_ex %
upper LMX	1105.8	1.15	66.65	12.50	0.64	1.20	3.41	2.29	0.44	70	10.52
	1116.1	2.10	74.82	11.50	0.39	1.00	2.12	1.18	0.35	77	15.98
	1125.8	3.41	62.16	12.20	2.51	1.20	2.94	1.89	0.4	70	8.92
	1135.0	3.62	61.09	10.10	1.02	1.00	3.38	1.97	0.41	65	11.88
	1135.8	4.18	63.12	13.95	2.62	0.80	3.59	2.32	0.47	73	6.49
	1138.9	4.28	68.69	12.68	0.41	0.90	3.30	1.57	0.46	73	11.18
	1141.2	3.31	63.02	10.30	0.81	1.10	3.38	1.93	0.41	66	12.45
	1142.1	4.12	71.16	10.70	0.59	0.80	2.73	1.15	0.34	72	15.59
	1145.6	4.45	69.61	8.50	3.12	0.60	2.13	1.65	0.29	72	18.49
	1147.9	3.17	68.72	9.10	1.25	1.10	2.64	1.58	0.34	65	17.09
	1152.6	3.19	62.93	15.50	0.95	1.20	3.72	1.63	0.63	73	3.85
	1156.2	2.28	56.23	11.20	6.31	0.80	2.46	4.43	0.31	73	7.80
	1157.8	2.87	71.78	10.60	0.65	0.90	2.74	1.08	0.36	71	16.04
lower LMX	1160.2	6.23	72.15	9.80	0.62	1.00	2.64	0.93	0.36	70	17.53
	1161.5	6.03	78.08	6.70	0.82	1.00	1.64	0.58	0.27	66	25.41
	1164.6	6.59	76.39	7.40	1.30	0.70	1.91	0.66	0.3	69	23.46
	1165.8	6.89	72.79	9.30	0.76	0.90	2.33	0.76	0.39	70	18.66
	1165.9	4.64	78.51	8.60	1.38	0.50	1.37	0.76	0.19	78	22.48
	1167.7	6.11	78.07	6.70	1.08	0.60	1.83	0.59	0.27	69	25.40
	1170.3	6.11	80.38	5.60	1.68	0.50	1.61	0.55	0.24	68	28.29
	1172.9	6.26	79.40	6.30	1.42	0.60	1.76	0.58	0.26	68	26.68
	1173.8	6.11	69.79	5.70	1.09	1.00	2.00	0.91	0.27	59	23.18
	1175.2	5.97	79.32	5.80	1.66	0.50	1.63	0.49	0.25	69	27.47
	1178.3	5.82	81.68	4.80	1.57	0.40	1.38	0.44	0.21	69	30.21
	1181.5	5.11	75.45	8.80	0.94	0.80	2.42	0.86	0.33	69	20.72
	1182.4	3.26	41.87	10.11	12.17	0.70	1.51	9.71	0.20	78	2.89
	1183.5	4.92	78.78	7.30	1.26	0.70	1.86	0.69	0.31	69	24.74
	1187.4	5.39	81.06	6.30	1.03	0.80	1.20	0.67	0.17	69	27.46
	1190.5	4.14	82.63	5.30	1.61	0.50	1.59	0.48	0.23	67	29.83
	1191.8	6.89	79.54	6.30	1.03	0.70	1.80	0.47	0.28	66	26.75
	1196.1	4.36	77.99	6.40	0.98	0.60	1.32	0.79	0.18	72	25.86
	1198.4	6.09	79.09	7.05	0.68	0.70	2.08	0.63	0.28	67	25.30
	1208.0	3.91	79.06	6.04	1.43	0.40	1.71	0.81	0.21	71	26.95
WF	1209.2	1.30	88.67	6.22	1.10	0.40	0.72	0.78	0.1	80	31.14
	1212.4	1.54	87.19	5.30	1.09	1.00	1.16	0.64	0.17	63	31.96
	1214.3	2.18	82.69	7.50	1.08	0.40	1.62	0.68	0.22	76	26.24
	1215.1	1.72	85.32	5.21	1.49	1.10	1.43	0.61	0.26	59	31.24
	1215.9	1.22	87.65	6.03	0.83	0.30	0.62	0.59	0.10	83	30.97
	1217.3	1.85	84.28	5.75	1.27	0.90	0.95	0.59	0.15	68	29.86
	1219.1	0.94	85.34	7.30	0.37	0.40	1.79	0.58	0.23	74	27.81

WF: Wufeng formation; CIA: chemical index of alteration.

Major and trace elements

In the present study, the ternary plot of the three endmembers of mudstone (SiO₂-Al₂O₃-CaO) shows that all three units contain more SiO₂ than Al₂O₃ and CaO (Figure 2(a)). The average SiO₂ content in the WF, lower LMX, and upper LMX Formations is 85.88%, 76.10%, and 66.15%, respectively (Table 1). The contents of other major elements, such as K₂O, P₂O₅, and MgO, increase gradually from the WF Formation to the upper LMX Formation (Table 1). The positive relationship of Al₂O₃ with K₂O and TiO₂ in the lower and upper LMX Formation suggests that these major elements have a terrigenous clastic source (r = 0.51, 0.49, and 0.65, 0.86, respectively, p < 0.05), as shown in Figure 2(b) and (c). However, this correlation is not significant in the WF Formation and may be influenced by tectonic movement and sea-level change (see ‘Terrestrial input flux and climate change’ section). The distributions of K/Al and Rb/Al in the study area exhibit similar characteristics (Figure 3). K/Al values are 0.16–0.43 (mean = 0.30) in the WF Formation, 0.23–0.55 (mean = 0.41) in the lower LMX Formation, and 0.29–0.52 (mean = 0.41) in the upper LMX Formation. Rb/Al values are 5.14–14.99 μg/(wt,%) (mean = 9.64 μg/(wt, %)) in the WF Formation, 6.41–21.52 μg/(wt, %) (mean = 16.46 μg/(wt, %)) in the lower LMX Formation, and 10.82–20.55 μg/(wt, %) (mean = 16.93 μg/(wt, %)) in the upper LMX Formation.

Figure 2.

Major element characteristics for the Wc-1 well. (a) Ternary diagram of SiO₂-Al₂O₃-CaO. The value for average shale was taken from Wedepohl (1971). (b) Covariance diagram of Al₂O₃ and K₂O. (c) Covariance diagram of Al₂O₃ and TiO₂.

Figure 3.

Stratigraphic profiles of the Wc-1 well for total organic carbon (TOC), the chemical index of alteration (CIA), and trace elements. The threshold values of V/(V+Ni) and Ni/Co were obtained from Jones and Manning (1994).

Moreover, the CIA values are 59–83 (mean = 72) in the WF Formation, 59–78 (mean =69) in the lower LMX Formation, and 65–77 (mean = 71) in the upper LMX Formation, Figure 3 and Table 1.

Trace elements, such as U, V, Ni, Mo, Cr, and Co, are all enriched in the three units (Figure 3 and Table 2). The values of Ni/Co, V/(V+Ni), P/Ti, and Ba_XS, which represent paleo-hydrographic proxies, are higher in the lower LMX Formation (mean = 12.19, 0.69, 0.25 and 4322.86 ppm, respectively), and lower in the WF and upper LMX Formations (mean = 5.80, 0.49, 0.18, and 1659.90 ppm, and 6.51, 0.60, 0.15, and 3938.73 ppm, respectively), as shown in Figure 3 and Table 2.

Table 2.

Trace elements and paleoenvironmental proxies of Wc-1 well.

Formation	Depth	Trace element						Paleoenvironmental proxies
Formation	Depth	V_EF	Cr_EF	Ni_EF	Mo_EF	U_EF	Co_EF	V/V+Ni	Ni/Co	Ba_XS (ppm)	P/Ti
upper LMX	1105.8	1.73	1.58	1.98	4.81	2.57	0.84	0.68	6.11	1833.88	0.14
	1116.1	1.72	0.89	3.14	11.33	4.12	1.04	0.57	7.83	2375.69	0.14
	1125.8	2.68	1.24	3.61	18.05	4.51	1.18	0.64	7.93	2801.12	0.16
	1135.0	2.11	1.62	2.91	19.49	6.52	1.95	0.64	3.87	4644.82	0.14
	1135.8	2.24	1.31	2.90	11.53	4.73	1.16	0.65	6.48	3361.70	0.16
	1138.9	2.61	0.43	3.32	12.95	4.91	1.22	0.66	7.05	5093.53	0.18
	1141.2	2.77	1.58	4.33	17.23	5.71	1.67	0.61	6.72	5453.66	0.14
	1142.1	2.08	1.82	4.13	14.44	4.54	1.48	0.55	7.23	7049.33	0.14
	1145.6	2.27	2.22	3.90	19.53	5.63	2.08	0.59	4.86	3154.12	0.20
	1147.9	1.96	1.64	3.77	17.12	5.43	1.59	0.56	6.15	3629.63	0.16
	1152.6	1.87	1.35	2.98	11.01	3.14	1.10	0.60	7.02	5524.69	0.15
	1156.2	1.78	1.21	3.15	13.12	4.15	1.18	0.58	6.92	2660.93	0.15
	1157.8	1.26	1.81	3.19	15.79	5.17	1.27	0.49	6.51	3620.41	0.12
lower LMX	1160.2	3.70	1.96	4.17	29.33	8.33	1.23	0.68	8.79	6003.06	0.17
	1161.5	3.55	3.71	5.69	46.79	11.49	1.75	0.60	8.43	2876.57	0.26
	1164.6	4.22	3.49	6.07	49.11	10.11	1.55	0.63	10.15	5122.01	0.25
	1165.8	9.35	2.44	4.65	55.32	11.63	1.74	0.83	6.93	3607.47	0.20
	1165.9	3.08	1.09	3.92	48.49	10.11	0.75	0.66	13.55	5486.11	0.25
	1167.7	5.32	4.07	7.15	56.68	13.44	1.52	0.64	12.19	4799.57	0.23
	1170.3	6.98	4.15	8.44	58.53	15.79	1.83	0.67	11.96	4032.46	0.24
	1172.9	7.81	3.44	8.30	63.35	15.54	1.61	0.70	13.36	3513.90	0.23
	1173.8	9.08	2.11	8.69	68.11	17.96	1.89	0.72	11.92	3608.38	0.27
	1175.2	8.07	5.01	8.13	74.12	15.72	1.61	0.71	13.09	7657.67	0.22
	1178.3	7.70	7.31	9.07	71.21	17.91	1.86	0.67	12.64	7125.33	0.27
	1181.5	5.43	5.33	5.73	44.33	10.12	1.21	0.70	12.28	4428.87	0.19
	1182.4	2.44	0.74	2.27	29.21	6.32	0.35	0.72	16.81	2338.61	0.36
	1183.5	5.67	5.07	6.09	46.33	11.09	1.34	0.69	11.78	4228.09	0.21
	1187.4	6.27	1.31	5.17	69.37	13.33	1.29	0.75	10.39	5608.90	0.29
	1190.5	7.15	3.13	7.49	57.47	11.27	1.36	0.70	14.28	3473.71	0.25
	1191.8	9.55	2.46	8.44	64.78	13.71	1.49	0.73	14.68	5921.90	0.24
	1196.1	7.59	2.74	6.07	60.05	10.93	1.42	0.75	11.08	2430.88	0.29
	1198.4	8.97	2.24	10.11	77.47	12.24	1.18	0.68	22.21	2794.29	0.23
	1208.0	4.69	1.89	9.93	9.84	5.44	3.54	0.53	7.27	1399.36	0.26
WF	1209.2	1.76	1.15	3.17	3.17	3.06	1.31	0.57	6.27	1372.46	0.27
	1212.4	1.95	7.33	5.83	5.83	4.17	3.01	0.45	5.02	1587.63	0.20
	1214.3	3.01	9.55	9.58	10.58	5.53	3.08	0.43	8.06	1362.27	0.13
	1215.1	2.47	5.27	5.87	6.87	4.23	3.15	0.51	4.83	1646.74	0.15
	1215.9	1.89	1.39	2.68	2.68	4.01	1.39	0.63	5.00	1879.92	0.23
	1217.3	2.45	5.98	7.35	7.35	4.68	3.16	0.45	6.03	1530.11	0.14
	1219.1	1.32	9.17	5.02	6.02	4.03	2.41	0.39	5.40	2240.19	0.17

WF: Wufeng formation.

The values of Si_ex are 26.24–31.96% (mean = 29.89%) in the WF Formation, 2.89–30.21% (mean = 23.96%) in the lower LMX Formation, and 3.85–18.49% (mean = 12.02%) in the upper LMX Formation (Table 1). To use Si_ex as a substitute for biological silica, it is necessary to prove that the non-clay silica in the sample may be biogenic rather than clastic. Although clastic SiO₂ may be negatively correlated with clay content, these sedimentary components are more likely to be co-deposited. Therefore, there is a strong negative correlation between SiO₂ and Al₂O₃ in the Wc-1 well (r = −0.54, p < 0.01, n = 40, not shown), suggesting the dominant biological properties of non-clastic SiO₂ in the samples, which is consistent with the results of previous studies (Michalopoulos and Aller, 2004). In addition, the latest petrological study of the WF-LMX Formation records the existence of sponge and radiolarian fossils, which further supports this finding (Khan et al., 2019). Therefore, it is reasonable to use excess silica as a substitute index for biological silica in the present study.

Biomarkers

GC-MS traces of n-alkanes, isoprenoids, terpanes, and steranes are shown in Figure 4. Their distribution and content do not differ significantly among the three units.

n-alkanes: the n-alkanes are mainly distributed in n-C₁₆ to n-C₃₁ in the Wc-1 well samples, with the highest values observed at n-C₁₈ (Figure 4). The values of n-C₁₇/n-C₃₁ in all samples range from 5.03 to 17.89, with an average of 10.57, indicating that light hydrocarbons are dominant (Table 1). However, it should be noted that the light hydrocarbon content in the WF Formation is lower than that in the LMX Formation (Figure 4), suggesting that OM in these shales may be influenced by biodegradation.

Figure 4.

Saturated hydrocarbon mass chromatogram of the Wc-1 well samples from the (a) upper and (b) lower Longmaxi, and (c) Wufeng formations.

Isoprenoids: All samples contain relatively high concentrations of the isoprenoids pristane (Pr) and phytane (Ph), with Pr/Ph values of 0.21–0.65 (average 0.43) (Table 1).

Terpanes: Hopanoid (H) series, tricyclic terpene (TT) series, and a small amount of tetracyclic terpene (Te) compounds were detected in the samples by mass chromatography. The C₂₃TT/C₃₀H values are 0.75–4.03 (Table 1).

Steranes: The complete series of steranes were detected in the samples that are not biodegraded (Figure 4), with ααα(20 R)C₂₇ sterane/ααα(20R)C₂₉ sterane (C₂₇/C₂₉St) values of 1.13–1.57 (mean = 1.35) (Table 3).

Table 3.

Proxies of biomarker s in WF–LMX formation of Wc-1 well.

Formation	Depth(m)	TOC	n-C₁₇/n-C₃₁	Pr/Ph	C₂₇/C₂₉St	C₂₃TT/C₃₀H
upper LMX	1105.8	1.15	7.17	0.44	1.35	2.68
	1116.1	2.10	7.02	0.35	1.23	0.75
	1125.8	3.41	12.79	0.58	1.57	4.03
	1135.8	4.18	12.21	0.47	1.31	1.17
	1147.9	3.17	10.11	0.44	1.13	1.21
	1157.8	2.87	14.76	0.47	1.27	1.45
lower LMX	1160.2	6.23	12.17	0.45	1.45	2.76
	1165.8	6.89	13.15	0.42	1.46	4.03
	1172.9	6.26	11.41	0.44	1.21	1.72
	1175.2	5.97	10.43	0.41	1.49	3.06
	1181.5	5.11	10.14	0.29	1.35	1.42
	1183.5	4.92	6.17	0.21	1.37	1.47
	1187.4	5.39	13.35	0.33	1.32	0.78
	1190.5	4.14	10.11	0.46	1.33	1.32
	1198.4	6.09	5.03	0.37	1.43	0.87
	1208.0	3.91	7.89	0.65	1.32	1.54
WF	1209.2	1.30	8.43	0.42	1.22	1.75
	1215.9	1.22	17.89	0.54	1.41	1.80

Pr: pristine; Ph: phytane; TT: tricyclic terpane; St: regular sterane; H: Hopanoid.

Discussion

Origin of sedimentary OM

A previous study on soluble OM and δ¹³C values in Lower Paleozoic marine shales in the Sichuan Basin confirmed that, due to the existence of mineral skeleton, OM formed from predominantly calcareous-walled phytoplankton (Chlorella and Dinoflagellates) in the Upper Ordovician–Lower Silurian deposition period, which would have been largely protected during pyrolysis (Tuo et al., 2016). In that study, biomarkers indicated the origin of OM relatively clearly. Thus, the content and distribution of specific components in biomarkers play an important role in identifying the OM origin (Clark and Blumer, 1967; Tang et al., 2019). In the present study, the n-alkanes of all samples were bimodal in distribution, the values of n-C₁₇/n-C₃₁ were all greater than 4.0, and the content of light hydrocarbons was relatively high (Table 1 and Figure 4). These results suggest that the OM in these black shales was predominantly derived from phytoplankton, with a small amount of OM input from terrestrial organic materials, which was also demonstrated by Clark and Blumer (1967) and Han and Calvin (1969).

Some studies have reported that tricyclic terpanes originate from lipids of bacteria or algae (Azevedo et al., 1992; Ourisson et al., 1982) and that hopanoid is mainly derived from prokaryotes such as bacteria (Ourisson et al., 1982). Tricyclic terpanes are more stable in the thermal maturity stage than hopanoid (Peters et al., 2005). In addition, based on an analysis of nine samples from the WF–LMX Formation, Li (2019) determined that the degree of OM thermal maturity in this area was in the early stage of high maturity (Ro = 1.31–1.47, mean 1.39), whereas the effect of pyrolysis on hopanoid was small (Peters et al., 2005). Therefore, the value of C₂₃TT/C₃₀H can reflect the relative contribution of lipids (tricyclic terpanes) from bacteria or algae and markers from different prokaryotes (hopanoid) in the shales of the WF–LMX Formation. The average value of C₂₃TT/C₃₀H in the black shale of the WF, upper LMX and lower LMX Formations was more than one in the Wc-1 well. In addition, there was a positive correlation with TOC (Figure 5 and Table 1), suggesting that planktonic algae were the main origin of OM in the WF–LMX Formation. Furthermore, ααα(20 R)C₂₇ steranes are mainly derived from phytoplankton, whereas ααα(20 R)C₂₉ steranes are mainly derived from higher plants (Peters et al., 2005). Therefore, the C₂₇/C₂₉St ratio of greater than 1.0 obtained for all samples (Table 1) suggests that OM in the WF–LMX Formation was mainly derived from lower marine plankton.

Figure 5.

Relationship between C₂₃TT/C₃₀H and total organic carbon (TOC) in the Wc-1 well samples.

Paleoproductivity

The relative stability of the valence states of P and Ba and their characteristics in the process of biological growth and diagenesis are often used as indicators of paleoproductivity (Algeo et al., 2011; Dymond and Collier, 1996; Latimer and Filippelli, 2002). Furthermore, the redox conditions of the marine environment also control their recycling, i.e., oxic conditions conducive to the enrichment of P and Ba in sedimentary OM, whereas anoxic (or euxinic) conditions recycle P and Ba from OM into the overlying water columns. This relationship is conducive to maintaining high primary productivity in surface water (Algeo and Maynard, 2004; Latimer and Filippelli, 2002; Liu et al., 2019). To eliminate errors caused by the input of terrigenous debris, the ratios of P to Ti (P/Ti) and Ba_XS were selected to discuss the paleoproductivity level of the study area instead of the absolute P and Ba contents (Algeo and Maynard, 2004; Latimer and Filippelli, 2002). In the present study, the values of P/Ti and Ba_XS from all three units were greater than 0.13 (PASS) and 1000 ppm, indicating high productivity (Schoepfer et al., 2015). The paleoproductivity of the lower LMX Formation was the highest (P/Ti = 0.25, Ba_XS = 4322.86 ppm) (Figure 3 and Table 2).

However, the correlations between TOC and P/Ti and Ba_XS differed for the three units in the Wc-1 well (Figure 6). For the WF Formation, the correlation between TOC and Ba_XS was negative (r = −0.76, n = 7, p < 0.05); however, the negative correlation between TOC and P/Ti was not significant (r = −0.59, n = 7, p > 0.1) (Figure 6), which may be related to the redox conditions. For the upper LMX Formation, both relationships were positively correlated (r = 0.56 and 0.58, n = 13, p < 0.05, respectively). For the lower LMX Formation, the TOC content was negatively correlated with P/Ti (r = −0.55, n = 20, p < 0.05) but positively correlated with Ba_XS (r = 0.45, n = 20, p < 0.05). As mentioned above, this result may indicate that the paleoproductivity level and organic carbon content do not always correspond to each other, as this depends on the redox conditions of the water at that time. More OM can be preserved only when both high productivity and strong reduction water conditions occur. Higher preservation is driven by anoxic conditions in the water column and sediment (see ‘Main factors controlling OM enrichment’ section).

Figure 6.

Relationship between total organic carbon (TOC) and P/Ti and Ba_XS in the Wc-1 well samples. LMX, Longmaxi Formation; WF, Wufeng Formation.

Basin restriction

Previous studies have shown that the degree of water closure, which is caused by basin restriction, often leads to changes in the redox environment and paleoproductivity (Algeo and Lyons, 2006). The degree of water closure is often evaluated by the correlation between Mo and TOC (Algeo and Lyons, 2006). This is because Mo typically exists in seawater as stable MoO₄²⁻ under oxic conditions, where it is slowly removed from the water column by adsorption on ferromanganese crusts. However, in anoxic or even euxinic water columns, Mo can be reduced to particle-reactive MoO_xS_4-x²⁻ or MoS₄²⁻ and rapidly fixed in sediments through capture by pyrite or OM (Erickson and Helz, 2000; Helz et al., 1996; Tribovillard et al., 2004). It should be noted that the accumulation of authigenic Mo, which occurred in the TOC-poor OM formed under oxic conditions, may be affected by oxic conditions due to the lack of H₂S required to convert MoO₄²⁻ to MoO_xS_4-x²⁻ or MoS₄²⁻ (Helz et al., 1996; Algeo and Rowe, 2012). Under such conditions, the low Mo concentration and Mo/TOC ratio reflect the redox conditions, not the degree of water column restriction (Westermann et al., 2013). Therefore, the Mo/TOC ratio of the TOC-poor OM section cannot be used to evaluate the degree of basin restriction. In contrast, the changes in Mo/TOC in the TOC-rich OM accumulated under anoxic conditions may better reflect restriction of the water column.

The Mo/TOC values for the lower and upper LMX Formation samples fall within the high-basin-restriction area (Figure 7), suggesting that the sedimentary OM was deposited in a deep water shelf environment, which is restricted by ancient land and has poor connectivity with the outside ocean. For the WF Formation, the Mo/TOC values of the seven samples were relatively low (mean = 2.56), falling in the high-basin-restriction area. However, as mentioned above, owing to the low TOC and Mo values of this formation (mean = 1.50% and 4.0 ppm, respectively), these Mo/TOC values may reflect the redox conditions more than the degree of water column restriction. According to the geological structure evolution during this period, it is speculated that contact between the water column and the outside marine environment was limited but not completely closed. In addition, oxygen-rich and nutrient-rich upwelling from high-latitude seawater during the ice age not only promoted the paleoproductivity level, but also changed the redox conditions of the water column (Lüning et al., 2000). This interpretation is consistent with the relatively high Si_ex and low TOC content in the WF Formation. The latest petrological study of the WF–LMX Formation in the northern margin of the Upper Yangtze Platform further supports this conclusion (Xiao et al., 2020).

Figure 7.

Relationship between Mo and total organic carbon (TOC) in the Wufeng (WF) and lower and upper Longmaxi (LMX) Formations. Modified from Algeo and Lyons (2006) and Li et al. (2017).

Redox conditions

The relative content and ratio of alkenes such as Pr/Ph are often used to measure the redox conditions and OM type of the sedimentary environment. For example, Pr/Ph < 0.8, 1.0 < Pr/Ph < 3.0, and Pr/Ph > 3.0, indicate strongly reducing (anoxic), dysoxic, and oxic depositional environments, respectively (Peters and Moldowan, 1991; Zhang et al., 2020). The Pr/Ph values from the samples were all less than 0.7 (Table 1), indicating a strongly anoxic marine sedimentary environment.

The enrichment of U, V, and Mo is often used in the analysis of redox conditions (Hatch and Leventhal, 1992; Jones and Manning, 1994; Ross and Bustin, 2009; Wilkin et al., 1997). This is because changes in the redox conditions of the sedimentary environment affect the abundance of these redox-sensitive trace elements that are abnormally enriched in a highly reducing sedimentary environment. U_EF, V_EF, and Mo_EF were all greater than one in the three units, indicating that these elements may have been deposited in anoxic or dysoxic bottom water (Lézin et al., 2013).

It should be noted that there were some differences in the relationship between TOC and U_EF, V_EF, and Mo_EF in the three units (Figure 8). Compared with the lower and upper LMX Formation, the TOC of the WF Formation has a stronger correlation with U_EF, V_EF, and Mo_EF (r = 0.78, 0.96, and 0.84, p < 0.05, respectively), suggesting that redox conditions were the main factors influencing the preservation of sedimentary OM during this period.

Figure 8.

Relationship between total organic carbon (TOC) and U_EF, V_EF, and Mo_EF in the Wc-1 well samples.

Furthermore, the redox properties of the paleo-sedimentary environment can be further evaluated using V/(V+Ni) and Ni/Co abundance ratios (Jones and Manning, 1994). The average V/(V+Ni) and Ni/Co ratios for the WF and lower and upper LMX Formations were 0.58 and 5.19, 0.69 and 12.49, and 0.66 and 5.70, respectively (Table 2). These are located in the oxic/dysoxic, anoxic, and dysoxic regions, respectively, indicating that the formation environment of the study area changed from oxic/dysoxic to anoxic then dysoxic sedimentary environments (Hatch and Leventhal, 1992; Yan et al., 2009) (Figure 3).

Moreover, except under basin restriction conditions, a high TOC flux from the surface water leads to anoxia in the water column, which leads to enrichment of trace elements (McLennan, 2001; Ross and Bustin, 2009). Importantly, biomarker data, such as Pr/Ph, can distinguish the major factor controlling sedimentary OM formation (Peters and Moldowan, 1991). In the present study, the Pr/Ph values of < 0.7, combined with the positive relationship between TOC with U_EF, V_EF, and Mo_EF (Figure 8), demonstrate that the redox conditions caused by basin restriction and/or water stratification were the main factors controlling the accumulation of sedimentary OM in this area. Similar results were also reported in previous studies (Chen et al., 2016; Li et al., 2017; Yan et al., 2015).

Terrestrial input flux and climate change

The clay mineral content can reflect the terrestrial inputs. Alkali metals such as Ca, Na, and K can easily migrate from feldspar to form clay minerals, with the ratio of Al₂O₃ typically increasing with increased formation of weathering products (Nesbitt and Young, 1982).

Due to the different degrees of chemical weathering under different climatic conditions, CIA values fall into three ranges: 50–65, 65–85, and 85–100, which reflect weak weathering in cold and dry climates, moderate weathering in warm and humid climates and strong weathering in hot and humid climates, respectively (Fedo et al., 1995; Nesbitt and Young, 1982) (Figure 9). To avoid the influence of the chemical composition of the source rocks, the trend in CIA rather than the CIA value itself is used to discuss climate change via the absolute climatic conditions. The variation curve of CIA reveals substantial climatic fluctuation in the WF Formation but a relatively stable climate in the LMX Formation, except for local climatic fluctuations, which is consistent with climate change during the Ordovician–Silurian transition (Brenchley et al., 2003; Harper et al., 2014) (Figure 3). The variation trend between TOC and CIA indicates that stable climatic conditions are more conducive to the preservation of sedimentary OM.

Figure 9.

Paleo-climate discrimination diagrams of shale samples from the Wufeng (WF), and upper and lower Longmaxi (LMX) Formations. The base plot was taken from Li et al. (2017) and Yan et al. (2015).

Al, K, and Rb are typical chemical elements used to indicate the terrestrial inputs. These elements are transported to sedimentary basins by rivers or wind in the form of silicate minerals, and then widely used to reconstruct continental inputs (Murphy et al., 2000). In the upper and lower LMX Formation, the average values of K/Al and Rb/Al were 0.41 and 0.41, and 16.46 and 16.93, respectively, and did not change significantly (Figure 3), suggesting that the terrigenous input remained relatively stable and the climate did not fluctuate significantly during this period. However, in the WF Formation, the values of K/Al and Rb/Al were relatively low (mean = 0.30 and 9.64 × 10⁻⁴, respectively) with large fluctuations (Figure 3). Combined with the paleogeography and paleoclimate characteristics, it is speculated that the cause of this phenomenon may be related to intermittent upwelling at the bottom of the water column during this period (Xiao et al., 2020; Zhao et al., 2019). The existence of upwelling hinders the precipitation of terrigenous clastic sediments, such as K and Al, and reduces the deposition of terrigenous detritus.

In addition, the relationship between TOC with K/Al and Rb/Al for the black shale of the study area showed a significant positive correlation in the lower LMX Formation (r = 0.49 and 0.65, respectively, p < 0.05) (Figure 10(a) and (b)), suggesting that the input of terrigenous detritus did not dilute the sedimentary OM, which may be related to the fact that detrital minerals promote the burial and preservation of OM through adsorption (Kennedy et al., 2014). However, in the upper LMX Formation, this relationship exhibited a negative correlation (Figure 10(b)). The reason for this phenomenon may be related to the tectonic movement during this period, whereby continuous topographic uplift led to a gradual decrease in sea level and a gradual increase in the relative terrigenous input in the basin, which limited the enrichment of sedimentary OM.

Figure 10.

Relationship between total organic carbon (TOC) and K/Al and Rb/Al, and Ti/Al in the Wufeng (WF), lower and upper Longmaxi (LMX) Formations.

Furthermore, studies have shown that the content of sedimentary OM in black shale is related to the deposition rate, with the value of Ti/Al increasing with the sediment accumulation rate (Murphy et al., 2000; Yan et al., 2015). The positive correlations between TOC and Ti/Al in the lower LMX Formation (r = 0.65, p < 0.01, n = 20) (Figure 10(c)) suggest that higher sedimentation rates promote the burial of sedimentary OM and hinder the oxidation of OM.

Main factors controlling OM enrichment

As mentioned above, the complexity of the OM enrichment process requires the consideration of many factors, including productivity, oxygenation level, sedimentation rate, and geological events, as well as the interrelationships between these factors (Chen et al., 2016; Lash and Blood, 2014; Murphy et al., 2000).

During the sedimentary period of the WF Formation, several paleo-uplifts occurred in the Yangtze plate due to tectonic compression, which divided the plate into distinct undercompensated marine basins. The formation of the Gondwana Glacier led to a sharp drop in global temperature and the migration of nutrient-and oxygen-rich water at high latitudes to the equator, which led to rising ocean currents and an increase in surface biological productivity and water column oxidation (Chen et al., 2011; Mu et al., 2011; Pang et al., 2018; Yan et al., 2008) (Figure 11(a)). The negative correlation between TOC and P/Ti and Ba_XS in the WF Formation was due to the water redox conditions, not a decrease in primary productivity. This is because sedimentary OM is more easily oxidized and decomposed under an oxic environment with only a small part of insoluble OM deposited at the bottom of the water column, which leads to minimal P and Ba content in the OM. This is supported by the highest Si_ex values and the positive correlation between TOC and redox conditions (Figures 3 and 8).

Figure 11.

Model of the depositional processes for Wc-1 well samples from the (a) Wufeng, (b) lower and (c) upper Longmaxi formations.

By the time of the lower LMX Formation, the paleoclimate had warmed, and the ice sheet was rapidly melting. The rising sea level led to regional expansion of the ocean basin and deepening of the sea level. At this time, the ancient ocean was in a stratified state (Cheng et al., 2013), and the high level of nutrients resulting from terrestrial input promoted the growth of planktonic algae in the upper water, which was shown by the higher values of TOC, Ba_XS, and P/Ti (Kennedy and Wagner, 2011). The anoxic water environment in the lower water column promoted the preservation of OM, which led to substantial enrichment of redox-sensitive elements (Figure 11(b)). However, it should be noted that the correlations between TOC and Ba_XS and P/Ti were positive and negative, respectively (Figure 6). The negative relationship between TOC and P/Ti reflects the effect of the anoxic water column, which leads to P recycling from OM into the overlying water columns, rather than a decrease in primary productivity (Latimer and Filippelli, 2002). The positive correlation between TOC and Ba_XS, which was due to the faster accumulation rate of OM, meant that the sulfate-reducing bacteria did not reduce barite in time and kept it in the sediments (Loucks and Ruppel, 2007).

During the sedimentary period of the upper LMX Formation, tectonic uplift and the fall in sea level may have promoted oxygenation in the study area, which changed the anoxic environment of the water column to a dysoxic environment (Li et al., 2017) (Figure 11(c)). Otherwise, lower P/Ti and Ba_XS values indicated slightly lower primary productivity in the upper LMX Formation than those in the lower LMX Formation (Table 2). Combined with an increase in terrigenous detrital material input and the dysoxic environment, the TOC content of the upper LMX Formation continued to decrease, as shown in Figure 3.

Conclusions

Biomarkers in samples of the WF–LMX Formation obtained from the Wc-1 well in northeastern Chongqing suggest that the main origin of the sedimentary OM in the black shale was planktonic algae.

Combined with regional tectonic evolution and changes in sea level, the geochemical analysis of P/Ti and Ba_XS as a paleoproductivity index indicates that the depositional period of the WF–LMX Formation was characterized by high paleoproductivity.

The changing trends of Pr/Ph, U_EF, V_EF, and Mo_EF in the redox index indicate that the redox conditions of the water column in the study area were oxic/dysoxic, anoxic, and dysoxic in the WF, lower LMX, and upper LMX Formations, respectively.

The main factors controlling OM enrichment differed for the three units. For the WF Formation, the redox condition was the main influencing factor. For the lower LMX Formation, high paleoproductivity and anoxic environment were the main factors. For the upper LMX Formation, the paleoproductivity and terrigenous detrital material inputs were the main influencing factors.

Footnotes

Declaration of conflicting interests

The author(s) declared no potential conflicts of interest with respect to the research, authorship, and/or publication of this article.

Funding

The author(s) disclosed the following financial support for the research, authorship, and publication of this article: Authors would like to thank the financial support from the National Natural Science Foundation of China (No. 41772129).

ORCID iD

Zaitian Dong

References

Algeo

Kuwahara

Sano

, et al. (2011) Spatial variation in sediment fluxes, redox conditions, and productivity in the Permian-Triassic panthalassic ocean. Palaeogeography Palaeoclimatology Palaeoecology 308(1–2): 65–83.

Algeo

Lyons

(2006) Mo-total organic carbon covariation in modern anoxic marine environments: Implications for analysis of paleoredox and paleohydrographic conditions. Paleoceanography 21(1): n/a–n/a.

Algeo

Marenco

Saltzman

(2016) Co-evolution of oceans, climate, and the biosphere during the 'ordovician revolution': A review. Palaeogeography Palaeoclimatology Palaeoecology 458(1): 1–11.

Algeo

Maynard

(2004) Trace-element behavior and redox facies in core shales of upper pennsylvanian Kansas-type cyclothems. Chemical Geology 206(3–4): 289–318.

Algeo

Rowe

(2012) Paleoceanographic applications of trace-metal concentration data. Chemical Geology 324–325: 6–18.

Azevedo

Aquino

Simoneit

BRT

, et al. (1992) Novel series of tricyclic aromatic terpanes characterized in tasmanian tasmanite. Organic Geochemistry 18(1): 9–16.

Bao

Yin

, et al. (2020) Influence of reservoir properties on the methane adsorption capacity and fractal features of coal and shale in the upper permian coal measures of the South Sichuan coalfield, China. Energy Exploration & Exploitation 38(1): 57–78.

Bock

Mclennan

Hanson

(1998) Geochemistry and provenance of the middle ordovician austin glen member (normanskill formation) and the taconian orogeny in new England. Sedimentology 45(4): 635–655.

Boström

Kraemer

Gartner

(1973) Provenance and accumulation rates of opaline silica, Al, Ti, Fe, Mn, Cu, Ni, and Co in pacific pelagic sediments. Chemical Geology 11(2): 123–148.

10.

Brenchley

Carden

Hints

, et al. (2003) High-resolution stable isotope stratigraphy of upper ordovician sequences: Constraints on the timing of bioevents and environmental changes associated with mass extinction and glaciation. Geological Society of America Bulletin 115(1): 89–104.

11.

Chen

Zhou

, et al. (2016) The geochemical characteristics and factors controlling the organic matter accumulation of the late Ordovician-early silurian black shale in the upper Yangtze basin, South China. Marine and Petroleum Geology 76: 159–175.

12.

Chen

(2016) Influencing factors analysis of lower palaeozoic shale gas content in foreland fold thrust belts of South dabashan, North-Eastern chongqing. Petroleum Geology and Engineering 30(4): 18–22. (in Chinese with English abstract).

13.

Chen

Zhu

Qin

, et al. (2014) Reservoir evaluation of the lower silurian longmaxi formation shale gas in the Southern Sichuan basin of China. Marine and Petroleum Geology 57: 619–630.

14.

Chen

Zhu

Wang

, et al. (2011) Characteristics and significance of mineral compositions of lower silurian longmaxi formation shale gas reservoir in the Southern margin of Sichuan basin. Acta Petrologica Sinica 32(5): 775–782. in Chinese with English abstract).

15.

Cheng

Wang

Chen

, et al. (2013) Sedimentary and burial environment of black shales of Sinian to early palaeozoic in upper Yangtze region. Acta Petrologica Sinica 29(8): 2906–2912. in Chinese with English abstract).

16.

Clark

Blumer

(1967) Distribution of n-Paraffins in marine organisms and sediment. Limnology and Oceanography 12(1): 79–87.

17.

Demaison

Moore

(1980) Anoxic environments and oil source bed genesis. Organic Geochemistry 2(1): 9–1209.

18.

Dong

Shen

Xiao

, et al. (2011) Construction and structural analysis of regional geological sections of the Southern daba Shan thrust-fold belts. Acta Petrologica Sinica 27(3): 689–698.

19.

Dymond

Collier

(1996) Particulate barium fluxes and their relationships to biological productivity. Deep Sea Research Part II: Topical Studies in Oceanography 43(4–6): 1283–1308.

20.

Erickson

Helz

(2000) Molybdenum (VI) speciation in sulfidic waters: Stability and lability of thiomolybdates. Geochimica et Cosmochimica Acta 64(7): 1149–1158.

21.

Fan

Peng

Melchin

(2009) Carbon isotopes and event stratigraphy near the Ordovician-Silurian boundary, Yichang, South China. Palaeogeography Palaeoclimatology Palaeoecology 276(1-4): 160–169.

22.

Fedo

Wayne Nesbitt

Young

(1995) Unraveling the effects of potassium metasomatism in sedimentary rocks and paleosols, with implications for paleoweathering conditions and provenance. Geology 23(10): 921.

23.

Finnegan

Bergmann

Eiler

, et al. (2011) The magnitude and duration of late Ordovician-early silurian glaciation. Science (New York, N.Y.) 331(6019): 903–906.

24.

Gao

Sun

Zhang

, et al. (2012) Petroleum systems of the hebaochang gas field in the Southern of Sichuan basin, SW China. Energy Exploration & Exploitation 30(4): 623–643.

25.

Guo

Zeng

(2015) The structural and preservation conditions for shale gas enrichment and high productivity in the Wufeng-Longmaxi formation, southeastern Sichuan basin. Energy Exploration & Exploitation 33(3): 259–276.

26.

Han

Calvin

(1969) Hydrocarbon distribution of algae and bacteria, and microbiological activity in sediments. Proceedings of the National Academy of Sciences of the United States of America 64(2): 436–443.

27.

Harper

DAT

Hammarlund

Rasmussen

CMO

(2014) End ordovician extinctions: A coincidence of causes. Gondwana Research 25(4): 1294–1307.

28.

Hatch

Leventhal

(1992) Relationship between inferred redox potential of the depositional environment and geochemistry of the upper Pennsylvanian (missourian) stark shale member of the Dennis limestone, Wabaunsee county, Kansas, U.S.A. Chemical Geology 99(1–3): 65–82.

29.

Helz

Miller

Charnock

, et al. (1996) Mechanism of molybdenum removal from the sea and its concentration in black shales: EXAFS evidence. Geochimica et Cosmochimica Acta 60(19): 3631–3642.

30.

Jones

Manning

DAC

(1994) Comparison of geochemical indices used for the interpretation of palaeoredox conditions in ancient mudstones. Chemical Geology 111(1-4): 111–129.

31.

Kennedy

Löhr

Fraser

, et al. (2014) Direct evidence for organic carbon preservation as clay-organic nanocomposites in a Devonian black shale; from deposition to diagenesis. Earth and Planetary Science Letters 388: 59–70.

32.

Kennedy

Wagner

(2011) Clay mineral continental amplifier for marine carbon sequestration in a greenhouse ocean. Proceedings of the National Academy of Sciences 108(24): 9776–9781.

33.

Khan

Feng

Zhang

, et al. (2019) Biogenic silica and organic carbon fluxes provide evidence of enhanced marine productivity in the upper Ordovician-lower silurian of South China. Palaeogeography. Palaeoclimatology. Palaeoecology 534: 109278.

34.

Lash

Blood

(2014) Organic matter accumulation, redox, and diagenetic history of the marcellus formation, southwestern Pennsylvania, appalachian basin. Marine and Petroleum Geology 57: 244–263.

35.

Latimer

Filippelli

(2002) Eocene to miocene terrigenous inputs and export production: Geochemical evidence from ODP leg 177, site 1090. Palaeogeography Palaeoclimatology Palaeoecology 182(3–4): 151–164.

36.

Lézin

Andreu

Pellenard

, et al. (2013) Geochemical disturbance and paleoenvironmental changes during the early toarcian in NW Europe. Chemical Geology 341: 1–15.

37.

Liang

Bai

Zou

, et al. (2016) Shale gas enrichment pattern and exploration significance of well wuxi-2 in northeast Chongqing, NE Sichuan basin. Petroleum Exploration and Development 43(3): 386–358.

38.

Zhu

, et al. (2017) Origin of organic matter and paleo-sedimentary environment reconstruction of the triassic oil shale in Tongchuan city, Southern ordos basin (China). Fuel 208: 223–235.

39.

(2019) Reservoir characteristics and pore evolution Simulation research of Longmaxi Formation shale in Chengkou Area, Northeastern Chongqing. China: China University of Mining and Technology.

40.

Zhang

Ellis

, et al. (2017) Depositional environment and organic matter accumulation of upper Ordovician-lower silurian marine shale in the upper Yangtze platform, South China. Palaeogeography Palaeoclimatology Palaeoecology 466: 252–264.

41.

Liu

Luo

, et al. (2007) Structural style and deformation mechanism of the southern dabashan foreland fold-and-thrust belt, Central China. Frontiers of Earth Science in China 1(2): 181–193.

42.

Liu

Zhang

, et al. (2017) Evidences of biogenic silica of Wufeng-Longmaxi formation shale in Jiaoshiba area and its geological significance. Journal of China University of Petroleum (Edition of Natural Science) 41 (1): 34–41. (in Chinese with English abstract)

43.

Liu

Gong

, et al. (2019) Geochemical characteristics of the lower silurian longmaxi formation on the Yangtze platform, South China: Implications for depositional environment and accumulation of organic matters. Journal of Asian Earth Sciences 184: 104003.

44.

Lüning

Craig

Loydell

, et al. (2000) Lower silurian hot shales in North Africa and Arabia: regional distribution and depositional model. Earth-Science Reviews 49(1–4): 121–200.

45.

Loucks

Ruppel

(2007) Mississippian Barnett shale: Lithofacies and depositional setting of a deep-water shale-gas succession in the fort worth basin, Texas. AAPG Bulletin 91(4): 579–601.

46.

McLennan

(2001) Relationships between the trace element composition of sedimentary rocks and upper continental crust. Geochemistry, Geophysics, Geosystems 2.

47.

Michalopoulos

Aller

(2004) Early diagenesis of biogenic silica in the amazon Delta: Alteration, authigenic clay formation, and storage. Geochimica et Cosmochimica Acta 68(5): 1061–1085.

48.

Zhou

Liang

, et al. (2011) Early paleozoic sedimentary environment of hydrocarbon source rocks in the middle-Upper yangtze region and petroleum and gas exploration. Acta Petrol Sin 85(4): 526–532. (in Chinese with English abstract).

49.

Murphy

Sageman

Hollander

, et al. (2000) Black shale deposition and faunal overturn in the devonian appalachian basin: Clastic starvation, seasonal water-column mixing, and efficient biolimiting nutrient recycling. Paleoceanography 15(3): 280–291.

50.

Nesbitt

Young

(1982) Early proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature 299(5885): 715–717.

51.

Ourisson

Albrecht

Rohmer

(1982) Predictive microbial biochemistry – From molecular fossils to prokaryotic membranes. Trends in Biochemical Sciences 7(7): 236–239.

52.

Pang

Jiao

, et al. (2018) Characteristics of organic pores and composition of bio-precursors in the Wufeng and Longmaxi formation shales, Southern Sichuan basin, China. Energy Exploration & Exploitation 36(4): 645–664.

53.

Peters

Moldowan

(1991) Effects of source, thermal maturity, and biodegradation on the distribution and isomerization of homohopanes in petroleum. Organic Geochemistry 17(1): 47–61.

54.

Peters

Walters

Moldowan

(2005) The Biomarker Guide, Biomarkers and Isotopes in Petroleum Exploration and Earth History. Cambridge: Cambridge University Press

55.

Rimmer

(2004) Geochemical paleoredox indicators in Devonian-Mississippian black shales, Central Appalachian basin (USA). Chemical Geology 206(3–4): 373–391.

56.

Ross

DJK

Bustin

(2009) Investigating the use of sedimentary geochemical proxies for paleoenvironment interpretation of thermally mature organic-rich strata: Examples from the Devonian-Mississippian shales, Western Canadian sedimentary basin. Chemical Geology 260(1–2): 1–19.

57.

Schoepfer

Shen

Wei

, et al. (2015) Total organic carbon, organic phosphorus, and biogenic barium fluxes as proxies for paleomarine productivity. Earth-Science Reviews 149: 23–52.

58.

Tang

Zhu

, et al. (2019) Geochemical characteristics and origin of hydrocarbons in the mesoproterozoic reservoirs in the Liaoxi depression, NE China. Energy Exploration & Exploitation 38(2): 333–347.

59.

Taylor

McClcnnan

(1985) The Continental Crust: Its Composition and Evolution. An Examination of the Geochemical Record Preserved in Sedimentary Rocks. Carlton: Blackwell Scientific Publications.

60.

Tuo

Zhang

(2016) Organic matter properties and shale gas potential of paleozoic shales in Sichuan basin, China. Journal of Natural Gas Science and Engineering 28: 434–446.

61.

Tribovillard

Algeo

Lyons

, et al. (2006) Trace metals as paleoredox and paleoproductivity proxies: an update. Chemical Geology 232(1–2): 12–32.

62.

Tribovillard

Riboulleau

Lyons

, et al. (2004) Enhanced trapping of molybdenum by sulfurized marine organic matter of marine origin in mesozoic limestones and shales. Chemical Geology 213(4): 385–401.

63.

Wang

Liu

, et al. (2015) Shale gas reservoir characteristic analysis for Shuijingtuo formation in the Mountain region of the outside basin, Chongqing northeast. Mud Logging Engineering 26: 57–60. (in Chinese with English abstract).

64.

Wedepohl

(1971) Environmental influences on the chemical composition of shales and clays. Physics and Chemistry of the Earth 8: 307–333.

65.

Westermann

Stein

Matera

, et al. (2013) Rapid changes in the redox conditions of the Western Tethys ocean during the early Aptian oceanic anoxic event. Geochimica et Cosmochimica Acta 121: 467–486.

66.

Wilkin

Arthur

Dean

(1997) History of water-column anoxia in the black sea indicated by pyrite framboid size distributions. Earth and Planetary Science Letters 148(3-4): 517–525.

67.

Xiao

Liu

Ran

, et al. (2020) Geochemistry and sedimentology of the upper ordovician-lower silurian black shale in the Northern margin of the upper Yangtze platform, South China: Implications for depositional controls on organic-matter accumulation. Australian Journal of Earth Sciences 67(1): 129–150.

68.

Yan

Chen

Wang

, et al. (2009) Geochemical changes across the Ordovician-Silurian transition on the Yangtze platform, South China. Science in China Series D: Earth Sciences 52(1): 38–54.

69.

Yan

Wang

, et al. (2015) Geochemical characteristics in the Longmaxi formation (early silurian) of South China: Implications for organic matter accumulation. Marine and Petroleum Geology 65: 290–301.

70.

Yan

Wang

Chen

, et al. (2008) Sedimentary environment and development controls of the hydrocarbon sources beds: the upper ordovician Wufeng formation and the lower silurian Longmaxi formation in the Yangtze area. Acta Geologica Sinica 3: 321–327.

71.

Zhang

, et al. (2020) Characteristics of microorganisms and origin of organic matter in Wufeng formation and Longmaxi formation in Sichuan basin, South China. Marine and Petroleum Geology 111: 363–374.

72.

Zhao

Min

, et al. (2019) Development of upwelling during the sedimentary period of the Organic-rich shales in The Wufeng and Longmaxi formations of the upper Yangtze region and its impact on organic matter enrichment. Journal of Marine Science and Engineering 7(4): 99.

Origin and paleoenvironment of organic matter in the Wufeng–Longmaxi shales in the northeastern Sichuan basin

Abstract

Keywords

Introduction

Geological setting

Samples and methods

Samples

Analytical methods

Data presentation

Results

TOC abundance

Major and trace elements

Biomarkers

Discussion

Origin of sedimentary OM

Paleoproductivity

Basin restriction

Redox conditions

Terrestrial input flux and climate change

Main factors controlling OM enrichment

Conclusions

Footnotes

Declaration of conflicting interests

Funding

ORCID iD

References