Carbon Isotope Gradient of the Ediacaran Cap Carbonate in the Shennongjia Area and Its Implications for Ocean Stratification and Palaeogeography

0. INTRODUCTION

After the Neoproterozoic 'Snowball Earth' event, the cap carbonates overlying the Marinoan glacial diamictite are globally distributed (Hoffman and Schrag, 2002) due to rapid warming and extreme oceanic alkalinity during deglaciation (Hoffman et al., 1998) and possess characteristic negative carbon isotope values and enigmatic sedimentary textures (Shield et al., 2007; Hoffman et al., 2007; Jiang et al., 2006; Halverson et al., 2005; Hoffman and Schrag, 2002; James et al., 2001). These geochemical and sedimentary features, particularly from the cap carbonates in the Yangtze Block, South China can reveal the C-isotopic spatial differences among different sedimentary facies. Plenty of studies on C-isotopic geochemistry have shown that the cap carbonates, presented from the inner shelf, through the continental slope, to the basin in the Yangtze Block (Jiang et al., 2011, 2007; Zhu et al., 2003; Peng et al., 2016), possess a significant C-isotopic gradient and higher δ¹³C_carb values in the proximal shallow water than those in the distal deep water (Jiang et al., 2011; Shen et al., 2005; Zhou et al., 2004). Zhou et al. (2004)found a C-isotopic difference of 1‰-2‰ and attributed it to ocean stratification or temporal diachroneity. Shen et al. (2005)obtained a C-isotopic gradient of~3‰ in average and considered that it represents the primary C-isotopic difference associated with water depth which results from biological pumping and rapid rise of the post-glacial sea level (Hoffman et al., 1998). Afterwards, Jiang et al. (2007)proposed that the large surface-to-deep ocean C-isotopic gradient through the entire Doushantuo Formation is mainly dominated by anaerobic oxidation of dissolved organic carbon (DOC) and fluctuations of the chemocline in the ocean (Jiang et al., 2008).

However, paleogeography of the Ediacaran Doushantuo Formation in the Shennongjia area has still been ambiguous and little known, although Jiang et al. (2011)simply considered it as tidal flat settings. It is recorded that the Shennongjia area was in a tidal flat-to-shelf depositional environment from SE to NW during the Mesoproterozoic Period (Hu, 1997a, b). Therefore, in this study, we conducted a detailed study on the C-isotopic analysis of the Ediacaran Doushantuo cap carbonates in the Shennongjia area, combined with diagnostic sedimentary characteristics, to explore origins of the C-isotopic gradient as well as spatial variations, providing some enlightening information for the paleogeographic framework of the Shennongjia area during the Early Ediacaran time.

1. GEOLOGICAL SETTINGS

The Yangtze Block is bounded by the Qinling-Dabie orogenic belt on the north and Longmenshan fault zone on the west (Zhao and Cawood, 2012) with the Mesoproterozoic and Neoproterozoic strata exposed. The Shennongjia area, located in the northern Yangtze Block, is bounded by the Qingfeng fault and the Qinling orogenic belt on the north and next to the Huangling uplift on the southeast (Fig. 1a). It is a~1 800 km² structural dome, comprising the Shennongjia Group in the core and late Neoproterozoic to Phanerozoic strata along the rim (Fig. 1b). The Kongling high-grade metamorphic rocks consisting of Archean TTG and Paleoproterozoic meta-sedimentary rocks is~18 km southeast of the study area (Lu et al., 2016; Guo et al., 2014; Zhang et al., 2006a, b; Gao et al., 1999). The contact between the Kongling basement and the Shennongjia Group has not been observed (Qiu et al, 2011). The Shennongjia Group has been regarded as the Mesoproterozoic strata evidenced by a lot of geochronological data (Li et al., 2013; Li and Leng, 1991), but there still remains controversial. Different scenarios have always existed for the stratigraphic division of Shennongjia Group (Li et al., 2013; Li and Leng, 1991), due to the relative lack of reliable geochronological data. The Macaoyuan Group has been considered as the Molasse formation during the Early Neoproterozoic (Qiu et al., 2011; Li and Leng, 1991), but Wang et al. (2013)found that the Maocaoyuan Group should belong to Mesoproterozoic and was deposited in shallow water environment in terms of sedimentary facies and a constraint of precise zircon U-Pb ages. The Neoproterozoic strata mainly distribute along the rim and in the north of the Shennongjia area and are divided into Liantuo Formation, Pingqian Formation, Datangpo Formation, Nantuo Formation, Doushantuo Formation and Dengying Formation (Li and Leng, 1991). To date, few studies on the Neoproterozoic strata have been reported except for the black shale sandwiched in Nantuo diamictite of the Songluo Section (Ye et al., 2015) and the barite layers in Doushantuo carbonate of the Wushanhu Section (Killingsworth et al., 2013).

Figure  1.  (a) Simplified map showing the study area and the paleogeographic framework for the Yangtze Platform during the deposition of the Ediacaran Doushantuo Formation (modified from Wang et al., 2015 and Zhu et al., 2003). (b) Simplified geological map exhibiting the locations of the study sections and the Precambrian basement of the Shennongjia area, in northern Yangtze Block (modified from Li et al., 2013).

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2. SAMPLING SECTIONS AND METHODS

We sampled five sections of cap carbonates overlying the Nantuo diamictite including the Longxi, Muyu, Yazikou, Guogongping (the untested data) and Songluo sections in the Shennongjia area, Hubei Province, South China (Fig. 1b). At Longxi Section (31°38′20″ N, 110°48′00″ E), 2.5 m thick cap carbonate at the bottom of the Doushantuo Formation overlies 1.5 m thick Nantuo diamictite and underlies a suite of dark gray thin-bedded phosphoric shales. The Nantuo diamictite, overlying the dolostone-dominated Shennongjia Group, contains complex gravel components including coarse-grained conglomerate, brown ferruginous sandstone, and slight gray calcite pieces (Fig. 2a). Distinctive tepee-like structure with barite-and silica-filling sheet cracks is in the middle part of the cap carbonate (Fig. 2d). At Muyu Section (31°27′20″ N, 110°30′06″ E), ~50 km southwest of the Longxi Section, 1.1 m thick cap carbonate overlies~100 m thick Nantuo diamictite and underlies silty mudstone. At Yazikou Section (31°30′25″ N, 110°28′06″ E), ~5 km northwest of the Muyu Section, 0.9 m-thick cap carbonate overlies~100 m thick Nantuo diamictite and underlies silty mudstone. 4 cm-thick tuff bed occurs in the middle part of the cap carbonate (Fig. 2b). At Guogongping Section (31°28′01″ N, 110°09′02″ E), ~25 km west of the Yazikou Section, due to strata inversion, ~1 m thick cap carbonate overlies black shale containing carbonates and underlies ca.100 m-thick Nantuo diamictite (Fig. 2e). The slump structure (Fig. 2e) and wave ripples (Fig. 2f) are present in the cap carbonate. At Songluo Section (31°41′32″ N, 110°36′32″ E), ~20 km northwest of the Longxi Section, 0.4 m thick cap carbonate overlies~230 m thick Nantuo diamictite and underlies silty mudstone (Fig. 2c). In addition, the Chengkou Section in the Chengkou area of northeastern Chongqing is a referenced section in this study (Fig. 1a).

Figure  2.  Outcrop pictures of partial sections and sedimentary structures. (a) Longxi Section, (b) Yazikou Section, (c) Songluo Section, (d) Tepee-like structure with sheet cracks filled by barite and silica at Longxi Section, (e) Guogongping Section and its slump structure, and (f) Wave ripples at Guogongping Section.

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For δ¹³C_carb and δ¹⁸O_carb analysis, powders were micro-drilled from the dolomicritic matrix determined by petrographic observation. One hundred and fifty micrograms of carbonate powder was reacted online with anhydrous H₃PO₄ for 25 min at 70 ℃ for CO₂ collection. The carbon and oxygen isotopes of generated CO₂ was measured on a Finnigan MAT-253 mass spectrometer. Isotopic results are expressed in the standard δ notation as per mil (‰) deviation from V-PDB. The analytical precision is better than±0.1‰ for δ¹³C_carb and±0.3‰ for δ¹⁸O_carb based on duplicate analyses. All the analysis were conducted in the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences (Wuhan).

3. RESULTS

There are significant differences in the carbon and oxygen isotopic signatures from the Longxi, Muyu, Yazikou and Songluo sections (Table. 1, Fig. 3). At Longxi Section (Fig. 3a), δ¹³C_carb values of the cap carbonate show a range from-2.3‰ to 1‰ with an exception of 1.86‰ at the bottom of the cap carbonate. At 90 cm above the base of cap carbonate, δ¹³C_carb values reveal a positive excursion, up to 0.6‰ at 120 cm, and subsequently, rapidly decrease to the nadir (-2.28‰) at 140 cm where the tepee-like structure occurs. After that, δ¹³C_carb values slowly begin to rise, up to 0.66‰ at 200 cm, and then gently decrease until the top of the cap carbonate. δ¹⁸O_carb values range from-6‰ to-3‰ except a single data of-1.45‰ corresponding to δ¹³C_carb value of 1.86‰. At Muyu Section (Fig. 3b), δ¹³C_carb values in the lower part of the cap carbonate firstly increase and then decrease, ranging from-5‰ to-4‰, and a minor positive δ¹³C_carb shift occurs in the upper part of the cap carbonate. Nonetheless, δ¹⁸O_carb values have larger fluctuations, ranging from-10‰ to-6‰. At Yazikou Section (Fig. 3c), δ¹³C_carb values of the cap carbonate are considerably stable around-4‰, whilst δ¹⁸O_carb values are relatively stable, ranging from-9‰ to-7‰. At Songluo Section (Fig. 3d), δ¹³C_carb values of the cap carbonate show a range from-7‰ to-4‰. In the lower part, there is a gradually positive δ¹³C_carb excursion, with a peak of-4.07‰ at 9 cm above the base of cap carbonate. δ¹⁸O_carb values of the cap carbonate exhibit an oscillation rang ing from-7‰ to-5‰. Besides, the Chengkou Section, as a referenced section, has similarity with the Songluo Section in the C-isotopic variation tendency.

Table  1.  Carbon and oxygen isotopic compositions of Cap carbonates from four sections

Section	Sample ID	Height (cm)	Lithology	δ13C (%, VPDB)	δ18O (%, VPDB)
Longxi Section	LXC-3-1	3.0	Microcrystalline dolostone	-0.96	-4.96
	LXC-3-2	11.0	Microcrystalline dolostone	1.86	-1.45
	LXC-4-1	17.0	Microcrystalline dolostone	-0.97	-4.60
	LXC-4-2	22.0	Microcrystalline dolostone	-1.07	-4.80
	LXC-5	50.0	Microcrystalline dolostone	-0.37	-3.83
	LXC-6	62.0	Microcrystalline dolostone	-1.40	-4.31
	LXC-7	75.0	Microcrystalline dolostone	-1.36	-4.87
	LXC-8	84.0	Microcrystalline dolostone	-1.37	-4.71
	LXC-9	90.0	Microcrystalline dolostone	-1.36	-5.92
	LXC-10	116.0	Microcrystalline dolostone	0.57	-4.11
	LXC-11	124.0	Microcrystalline dolostone	0.38	-3.27
	LXC-12	131.0	Microcrystalline dolostone	-0.36	-4.59
	LXC-13	140.0	Microcrystalline dolostone	-2.28	-4.99
	LXC-14	154.0	Microcrystalline dolostone	-1.63	-4.51
	LXC-15	167.0	Microcrystalline dolostone	-0.60	-3.81
	LXC-16	188.0	Microcrystalline dolostone	-0.25	-5.33
	LXC-17-1	202.0	Gray black dolostone	0.61	-5.06
	LXC-17-2	202.0	Gray white dolostone	0.66	-5.06
	LXC-18-1	221.0	Gray black dolostone	0.53	-5.25
	LXC-18-2	222.0	Gray white dolostone	0.26	-5.17
	LXC-19	236.0	Microcrystalline dolostone	-0.11	-4.59
	LXC-20-1	250.0	Gray black dolostone	-0.80	-4.82
	LXC-20-2	250.0	Gray white dolostone	-0.87	-4.79
Muyu Section	MY-04-01	5.0	limestone	-4.30	-9.29
	MY-04-02	10.0	limestone	-3.95	-9.00
	MY-04-03	15.0	limestone	-4.00	-9.53
	MY-05-01	27.0	limestone	-4.67	-8.88
	MY-05-02	37.0	limestone	-4.10	-9.03
	MY-06	48.0	limestone	-3.97	-9.13
	MY-07	58.0	limestone	-3.91	-7.56
	MY-08	66.0	limestone	-4.02	-9.42
	MY-09	72.0	limestone	-4.15	-9.19
	MY-10	76.0	limestone	-4.16	-9.29
	MY-11	81.0	limestone	-3.36	-6.43
	MY-12	84.0	limestone	-3.93	-6.97
	MY-13	110.0	limestone	-5.36	-8.01
Yazikou Section	YZK-01-01	6.0	Microcrystalline dolostone	-3.81	-7.97
	YZK-02-01	19.0	Microcrystalline dolostone	-3.68	-8.28
	YZK-04-01	28.5	Microcrystalline dolostone	-3.72	-8.28
	YZK-07-01	59.5	Microcrystalline dolostone	-3.88	-7.48
	YZK-08-01 YZK-09-01	62.5 67.0	Microcrystalline dolostone Microcrystalline dolostone	-4.16-4.10	-8.46-7.49
	YZK-10-01	74.0	Microcrystalline dolostone	-4.24	-6.85
	YZK-11-01	82.0	Microcrystalline dolostone	-4.26	-7.84
Songluo Section	SL-ZC-1	0.5	Microcrystalline dolostone	-6.09	-6.59
	SL-ZC-2	2.0	Microcrystalline dolostone	-6.35	-6.88
	SL-ZC-3	9.0	Microcrystalline dolostone	-4.07	-5.66
	SL-ZC-4	11.0	Microcrystalline dolostone	-4.62	-5.46
	SL-ZC-5	16.5	Microcrystalline dolostone	-5.24	-6.31
	SL-ZC-6	21.0	Microcrystalline dolostone	-6.66	-6.84
	SL-ZC-7	24.0	Microcrystalline dolostone	-5.95	-6.34
	SL-ZC-8	29.0	Microcrystalline dolostone	-6.89	-6.90
	SL-ZC-9	38.0	Microcrystalline dolostone	-6.47	-6.60
	SL-ZC-10	43.0	Microcrystalline dolostone	-6.30	-5.21

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Figure  3.  Lithologic columns and carbon and oxygen isotope chemostratigraphy of the cap carbonate from most of the sections in this study. (a) The distinct C-isotope excursions and the corresponding O-isotope changes at Longxi Section, (b) The slight C-isotope excursions and the corresponding O-isotope changes at Muyu Section, (c) The extremely stable C-isotope excursions and the corresponding O-isotope changes at Yazikou Section, (d) The obviously positive C-isotope excursion and the corresponding O-isotope changes at Songluo Section, and (e) The isotope change trend at Chengkou Section (unpublished data as a reference in this study) which is similar to that in (d).

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There are no apparently linear correlation for the carbon and oxygen isotopes from four sections (Fig. 4), indicating that isotopic signals are not strongly modified by meteoric diagenesis (Knauth and Kennedy, 2009). In addition, all samples have greater δ¹⁸O_carb values than-10‰, which can be considered that C-isotopic compositions are not significantly influenced by diagenetic processes (Jacobsen and Kaufman, 1999; Kaufman and Knoll, 1995). Thus δ¹³C_carb values in this study can generally represent the primary C-isotopic compositions of coeval seawater.

Figure  4.  Cross-plot between δ¹³C_carb and δ¹⁸O_carb for the cap carbonates from the Longxi, Muyu, Yazikou, and Songluo sections in the Shennongjia area and the Platform-shelf, Shelf-slope, and Basinal sections in south China (derived from Shen et al., 2005).

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4. DISCUSSION

4.1 Possible Interpretations for Carbon Isotopic Stratification

If the C-isotope data in this study can basically represent the primary C-isotopic compositions of coeval seawater or syndepositional diagenetic signatures, our results reveal significantly distinct C-isotopic stratification with a gradient of~5‰ in average between the Longxi and Songluo sections (Figs. 3 and 4). In previous studies, the different magnitudes of Ediacaran ocean C-isotopic gradient exist between the surface and deep waters in different regions, such as 1‰ up to 10‰ in South China (Jiang et al., 2007; Shen et al., 2005; Zhou et al., 2004), up to 10‰-14‰ in the Quruqtagh region of northwestern China (Xiao et al., 2004) and Australia (Calver, 2000). Shen et al. (2008)and Huang et al. (2013)indirectly confirmed the post-glacial Neoproterozoic ocean stratification using the carbon and sulfur isotopes of the cap carbonates from northwest China and south China, respectively. Moreover, there also exists a surface-to-deep water C-isotopic gradient in modern marine, such as the Black Sea (e.g., Volkov, 2000; Deuser, 1970) and the Framvaren Fjord of southern Norway (Volkov, 2000). Therefore, it seems that the ocean C-isotopic gradient is globally prevalent regardless of the geological time.

As to this C-isotopic stratification of the Ediacaran cap carbonates, several possible interpretations have already been proposed. A first possible interpretation is meteoric alteration during the post-depositional processes (Knauth and Kennedy, 2009). But the possibility of such an interpretation is faint according to the correlation between carbon and oxygen isotopes (Fig. 4). A second possible interpretation is diachroneity of the cap carbonate precipitation which to some extent can interpret the C-isotopic variance in Namibia (Hoffman et al., 2007) and South Australia (Rose et al., 2010). The directly unambiguous influence of diachroneity on δ¹³C_carb values of the cap carbonate is that δ¹³C_carb values in the upper part of shallow water sections are approximately equal to those in the lower part of deep water sections. However, comparative analysis of δ¹³C_carb values among these sections in this study indicates that diachroneity of the cap carbonate deposition is not the foremost reason of C-isotopic stratification. While, the most plausible interpretation considered by Shen et al. (2005), which can also be applied to the C-isotopic stratification presented in this study is that the early Ediacaran ocean has the inherent characteristic of C-isotopic stratification associated with water depth collectively maintained by different environmental conditions among different facies in the aftermath of a Neoproterozoic glaciation. Shen et al. (2005)held that the effect of biological pumping might be the dominant factor to the origin of a large surface-to-deep seawater C-isotopic gradient. During the early Ediacaran Period, the immediately high p_CO2 in the atmosphere (Hoffman et al., 1998) could motivate photosynthesis to sequestrate more CO₂ via biological pumping. Photosynthesis can make the surface water enriched in ¹³C_carb through preferential uptake of ¹²C by primary producers whilst fallout carbonate particles (FCPs) with ¹³C-depleted organic matter are transported downwards and undergo remineralization in deep waters (Broecker and Peng, 1982) (Fig. 5a). Jiao et al. (2010)proposed a conceptual model of microbial carbon pump (MCP) temporarily storing the recalcitrant dissolved organic matter (RDOM) in contrast to the biological pump in modern marine, but considered that the biological pump is more effective than the MCP in oxic and eutrophic shallow seawaters. However, Jiang et al. (2007)believed that a strong Ediacaran surface-to-deep seawater C-isotopic gradient was most likely to derive from anaerobic oxidation of DOC via sulfate reduction in long-term anoxic deep waters. Shen et al. (2008)found sulfur isotopic signatures to verify the existence of long-term syn-glacial sulfate reduction in Ediacaran anoxic bottom waters. Although the ¹³C-enriched DIC generated by primary producers in the surface seawater can directly deposit as platform facies, FCPs with ¹³C-enriched DIC may capture ¹³C-depleted cements during the process of downward transport into the deep seawater (Jiang et al., 2007). Therefore, in this study, we think that a C-isotopic gradient of~5‰ between the Longxi and Songluo sections may be the results from both biological pump favored by Shen et al. (2005)in surface waters and anaerobic oxidation of DOC (Jiang et al., 2007) in deep waters. In addition, methane contributions may be involved into carbon cycling in deep waters (Fig. 5b), which needs further investigation.

4.2 Origins of Carbon Isotope Excursions

At Longxi Section (Fig. 3a), the average δ¹³C_carb value of~-1‰ generally indicates that the cap carbonate may be from primary productivity in the surface ocean and deposit in shallow water environment above the chemocline (Jiang et al., 2011). A single positive δ¹³C_carb value (1.86‰) at the bottom of the cap carbonate is insufficient to demonstrate the incorporation of methanogenesis. Although methanogenesis can generate the positive δ¹³C_carb values in carbonates in sulfide-rich or euxinic seawaters (Shen et al., 2016), the cap carbonate at Longxi Section may form in sulfate-rich or oxic sedimentary environment as a whole, which is difficult to satisfy the occurrence of methanogenesis. Except the single positive C-isotope value, δ¹³C_carb values in the lower part of the cap carbonate generally remains stable, indicating that the depositional setting is relatively steady during late transgression (Fig. 5a). Because the primary sulfate of oceans might have been greatly consumed and the riverine input of sulfate was limited during the glaciation, the deep-water ocean was very likely in a long-term sulfide-rich or euxinic environment. Under this circumstance, methane was produced via methanogenesis and further stored in the form of methane hydrate subsequently, which could lead directly to positive C-isotopic anomalies of carbonates in deep waters (Ader et al., 2009) (Fig. 5a). Therefore, we believe that the positive C-isotopic anomalies in the lower part of cap carbonates at both the Songluo and Chengkou sections (Fig. 3d, 3e) are most likely due to methanogenesis in deep waters.

Given that tepee-like structure with barite-and silica-filling sheet-cracks can be regarded as structures of subaerial exposure in a tidal flat environment (James et al., 2001; Kennedy, 1996; Plummer, 1978), the positive δ¹³C_carb excursion below the tepee-horizon may result from downward shift of chemocline during early regression. Due to post-glacial isostatic rebound (Zhou et al., 2010), intermittent exposure of carbonate sediments occurred at Longxi Section during late regression, resulting in the formation of tepee-like structure and negative C-isotopic anomalies within the tepee-horizon (Fig. 3a). In the process, the increasing sulfate supply from enhanced weathering to the ocean could induce anaerobic oxidation of the oceanic DOC via sulfate reduction, transferring light carbon to the ¹³C-enriched DIC reservoir in surface waters and producing negative C-isotopic anomalies in surface water sections (Jiang et al., 2007). In addition, intermittent exposure would trigger the mixing of seawater and freshwater with ¹³C-depleted DIC, becoming a possibility of causing negative C-isotopic anomalies (Fig. 5b). Meanwhile, the sea-level fall might trigger destabilization of methane hydrate fixed in seafloor sediments, resulting in methane release into the deep waters. In the anoxic environment, anaerobic oxidation of methane (AOM) mediated by archaea and sulfate-reducing bacteria (Wang et al., 2008; Jiang et al., 2006, 2003) could prompt the deposition of the cap carbonate with light carbon isotopes (Fig. 5b). Hence, the carbon isotopes in the upper part of cap carbonates from both the Songluo and Chengkou sections can return to nomal in the presence of AOM.

In the uppermost part of the cap carbonate at Longxi Section, the positive δ¹³C_carb excursion accompanied by the negative δ¹⁸O_carb excursion occurs, which is exceedingly common in other Doushantuo cap carbonates close to the proximal inner shelf in the Yangtze Gorges (Jiang et al., 2003) with higher δ¹³C_carb values than those in the outer shelf, slope and basin (Shen et al., 2005; Zhou et al., 2004). And yet, differed from that in the middle part of the cap carbonate, the positive δ¹³C_carb excursion on the top of the cap carbonate is inclined to the results from enhanced primary productivity (Shimura et al., 2014; Shen et al., 2005) and organic carbon burial (Jiang et al., 2010), possibly due to strong terrestrial input from extreme chemical weathering indicated by Ohno et al. (2008)and Sawaki et al. (2010)during early transgression (Fig. 5c).

Figure  5.  Schematic diagram interpreting origins of carbon isotope excursions in Fig. 3. (a) In the coexistence of photosynthesis in surface waters and DOC in deep waters, the ocean had a large C-isotopic gradient. During late transgression, Fallout carbonate particles (FCPs) with ¹³C-enriched DIC were directly preserved at Longxi Section while FCPs with ¹³C-depleted organic matter were transported downwards to the deep waters. Due to the long-term sulfide-rich or euxinic environment, methane generated by methanogenesis was stored as methane hydrate in the seafloor sediments, resulting in the positive C-isotopic excursion at Songluo Section. During early regression, downward shift of chemocline led to the positive C-isotopic excursion at Longxi Section; (b) During late regression, anaerobic oxidation of DOC accelerated by enhanced sulfate from oxidative weathering and mixing of freshwater with ¹³C-depleted DIC into the surface ocean could lead to the negative C-isotopic excursion at Longxi Section. Whilst destabilization of methane hydrate and AOM terminated the positive C-isotopic excursion at Songluo Section; (c) During early transgression, enhanced primary productivity and organic matter burial caused the positive C-isotopic excursion at Longxi Section.

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Compared with that at Longxi Section, the cap carbonate at Muyu Section possesses faintly similar carbon isotopic variation tendency (Fig. 3b). We believe that the Muyu Section might be nearby below the chemocline and the carbon isotopic variations were likely to be mildly influenced by the seawater environment above the chemocline. At Yazikou Section (Fig. 3c), δ¹³C_carb values of the cap carbonate remain extremely stable, indicating that the cap carbonate may deposit in the cryptic oceanic environment perturbed by neither the surface seawater nor the deep seawater.

4.3 Possible Implications for Palaeogeographic Framework

From the previous studies in South China (e.g., Jiang et al., 2011, 2007; Shen et al., 2005) and other regions (e.g., Xiao, 2004; Calver, 2000; Volkov, 2000), it is not difficult to find that the conspicuous C-isotopic gradient often exists in the surface-to-deep water ocean. By comparison with the C-isotopic gradient from platform-to-basin facies studied by Shen et al. (2005) (Fig. 4), we boldly speculate that a surface-to-deep water palaeogeographic framework from SE to NW may exist in the Shennongjia area during the Early Ediacaran Period.

Currently we have a limited number of diagnostic sedimentary evidences, probably providing some support for the palaeogeography in this study. Firstly, the tepee-like structure from Longxi Section (Fig. 2d), as an exposed structure, may indicate a surface water environment. The slump structures and wave ripples at Guogongping Section (Fig. 2e, 2f) probably represent a shallow-water slope environment, indicating the strong hydrodynamic condition. The ca.3 m-thick shale in Nantuo diamictite of the Songluo Section (Ye et al., 2015) is likely to indicate a deep-water euxinic environment. Equally importantly, compared with the Chenkou Section near the Southern Qinling Sea (Fig. 1a), the Songluo Section underwent an extremely parallel carbon cycling process during the early Ediacaran Period (Fig. 3d, 3e). Secondly, the thickness difference of Marinoan diamictites mentioned in Section 2, to some extent, can also be one of evidences of estimating paleo-water depth (Zhou et al., 2010). It is essential that the Shennongjia area had slightly apparent paleo-water depth changes from SE to NW during the Mesoproterozoic Period proposed by Hu. (1997a, b). Hence, the Shennongjia area during the Ediacaran Period may have inherited the palaeogeography of the Mesoproterozoic Period and been influenced by subsequent post-glacial tectonic uplift or subsidence, which needs further investigation and verification.

5. CONCLUSIONS

In this study, we investigate five sections of the Doushantuo cap carbonate in the Shengnongjia area and conclude as follows:

(1) There exists a large δ¹³C_carb gradient (~5‰) between the Longxi and Songluo sections commonly maintained by photosynthesis in shallow waters and anaerobic oxidation of DOC in deep waters.

(2) The fluctuations of carbonate carbon isotope compositions observed at Longxi, Muyu and Songluo sections, may be the results from post-glacial sea-level elevation changes. The stabilization of carbonate carbon isotope composition at Yazikou Section indirectly provides origin of the ocean C-isotopic stratification with support.

(3) Potential oceanic stratification revealed by δ¹³C_carb gradient implies a significant difference in sedimentary paleo-water depths from SE to NW in the Shennongjia area during the Early Ediacaran Period.