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Volume 31 Issue 1
Jan.  2020
Article Contents
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New Insight into Factors Controlling Organic Matter Distribution in Lower Cambrian Source Rocks: A Study from the Qiongzhusi Formation in South China

  • Sedimentary organic matter (OM) is a major reservoir of organic carbon in the global carbon cycle. Despite many studies, there still exist many debates on the mechanism of OM accumulation and preservation in marine sediments. We present a new field study of a Lower Cambrian shallow marine shelf sequence in the northern edge of the Yangtze Plate, China. Our results show that palynological OM and biogenic silica (Bio-Si) could be used alongside more conventional redox and paleo-productivity proxies to study the distribution of OM in marine sediments. The qualitative and quantitative study of palynological OM provides more detailed information on the nature of sedimentary organic carbon, which can be helpful in the assessment of primary productivity and OM preservation. In addition, the presence of Bio-Si stimulates the physical preservation of OM. Further analysis indicates that an increase in Bio-Si can promote OM preservation. This case-study provides insight into the intertwined factors controlling OM accumulation in the Early Cambrian.
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    Zhou, C. M., Jiang, S. Y., 2009. Palaeoceanographic Redox Environments for the Lower Cambrian Hetang Formation in South China: Evidence from Pyrite Framboids, Redox Sensitive Trace Elements, and Sponge Biota Occurrence. Palaeogeography, Palaeoclimatology, Palaeoecology, 271(3/4): 279-286.
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New Insight into Factors Controlling Organic Matter Distribution in Lower Cambrian Source Rocks: A Study from the Qiongzhusi Formation in South China

    Corresponding author: Shucan Zheng,
  • 1. State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences, Wuhan 430074, China
  • 2. School of Earth Sciences, China University of Geosciences, Wuhan 430074, China
  • 3. University of the Littoral Opal Coast, UMR 8187 LOG, F-59000 Lille, France
  • 4. University of Lille, CNRS, UMR 8198 Evo-Eco-Paleo, F-59000 Lille, France
  • 5. Petroleum Exploration and Production Research Institute, SINOPEC, Beijing 100083, China

Abstract: Sedimentary organic matter (OM) is a major reservoir of organic carbon in the global carbon cycle. Despite many studies, there still exist many debates on the mechanism of OM accumulation and preservation in marine sediments. We present a new field study of a Lower Cambrian shallow marine shelf sequence in the northern edge of the Yangtze Plate, China. Our results show that palynological OM and biogenic silica (Bio-Si) could be used alongside more conventional redox and paleo-productivity proxies to study the distribution of OM in marine sediments. The qualitative and quantitative study of palynological OM provides more detailed information on the nature of sedimentary organic carbon, which can be helpful in the assessment of primary productivity and OM preservation. In addition, the presence of Bio-Si stimulates the physical preservation of OM. Further analysis indicates that an increase in Bio-Si can promote OM preservation. This case-study provides insight into the intertwined factors controlling OM accumulation in the Early Cambrian.

  • The Shatan section is on the northern edge of the Yangtze Plate, South China. It is a fresh outcrop that spans from the Upper Ediacaran to the Middle Cambrian.

    The Lower Cambrian sedimentary accumulations in South China are important marine source rocks, characterized by high total organic carbon (TOC) and sulfide contents. The rocks are mostly dark carbonate and fine-grained clastics, rich in silica minerals (Hu et al., 2018; Li and He, 2014; Wang et al., 2014; Guo et al., 2007a; Steiner et al., 2001). These source rocks are widely distributed in the Middle and Upper Yangtze Region (Jiang X F et al., 2018; Zhang, 2014; Hu et al., 2012; Jiang G Q et al., 2012). The northern and northeastern Yangtze Plate was a carbonate platform near the continental margin (Fig. 1a). The filling was mainly affected by the emerged Hannan old land, Motianling old land and Kangdian old land. The paleo-island of central Hubei only existed in the Terreneuvian and has been submerged since Cambrian Series 2 (Fig. 1a) (Tian et al., 2001). The sea level generally rose in the Cambrian, and series of short-period transgressive-regressive cycles occurred (Haq and Schutter, 2008; Miller, 2005). As a result, black shale, carbonates or black siliceous shale accumulated in the lower Cambrian over the entire Yangtze Plate (Chen et al., 2009; Guo et al., 2007a). The contemporaneous strata are characterized by black mudstone and chert in the southeast margin region (Jin et al., 2016; Guo et al., 2013, 2007b; Chen et al., 2009; Yang et al., 2003; Zhao et al., 1999).

    Figure 1.  (a) Paleogeographic map of the North Yangtze Plate during the Early Cambrian (modified from Li and He, 2014; Wang et al., 2012; Tian et al., 2001); (b) stratigraphic column, system tract and the sea-level changes of the Qiongzhusi Fm., Shatan Section. TST. Transgressive system tracts; HST. high stand system tracts.

    Samples are taken from the Qiongzhusi Formation (Fm.) which spans from the Cambrian Fortunian to Stage 3. This Formation accumulates shallow shelf facies, and has a total thickness of 459 m. It uncomformably overlies the dolomite strata of the Ediacaran Dengying Fm. and is comformably overlain by the oolitic limestone of the Xiannüdong Fm.. The Qiongzhusi Fm. is divided into 5 beds according to the lithologic characteristics (Fig. 1b). Lithologically, the lower part of the Qiongzhusi Fm. is composed of laminated silty mudstone (Bed 1) and black carbonaceous mudstone with phosphatic nodules (Bed 2). A large set of laminated calcareous mudstone and fine-grain sandstone is deposited in its middle part (Bed 3). Bed 4 is constituted by black shale with phosphorite nodules. The uppermost part of the formation (Bed 5) is composed of laminated calcareous mudstone and carbonaceous silty mudstone inter-layered with calcareous nodules. The sea-level changes could be echoed by the lithology and stratigraphic characteristics in the section. Two transgressive system tracts (TST) and two high stand systems tracts (HST) are revealed (Fig. 1b). The thin layers (3-7 cm) of carbonaceous mudstone and black shale in beds 1, 2 and 4 correspond to the water-deepening events of the TSTs (Fig. 1b). The relatively thick layers (8-25 cm) of laminated calcareous mudstone and carbonaceous silty mudstone in beds 3 and 5 correspond to the HSTs (Fig. 1b).

    Index small shelly fossils belonging to the Lower Cambrian Fortunian assemblages, Anabarites trisulcatus-Protohertzina anabarica and Paragloborilus subglobosus-Purella squamulosa, are abundant in the base of the Qiongzhusi Fm., especially in the phosphorite nodule level, representing a Nemakit-Daldynian age (equivalent to the early Meishucunian) (Guo et al., 2014; Steiner et al., 2004). The middle and upper parts of the Qiongzhusi Fm. are correlated to Cambrian Stage 3 according to the trilobite fossils from the carbonaceous claystone-siltstone, including Eoredlichia, Yunnanocephalus sp., Y. nanjiangensis and Wutingaspis sp. etc. (Zheng et al., 2014; Mo, 2012; Steiner et al., 2001; Luo et al., 1994).

  • A total of 62 samples were collected evenly throughout the section for palynological study. Fifty grams of each sample were processed using standard palynological extraction method (Lei et al., 2012; Traverse, 2007). The samples were observed using transmitted-light microscope (Leica DM5500) at the State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (Wuhan). To evaluate the relative abundance and variation of OM, a tablet with known number of synthesized lycopodium spores (27 637 pieces in each tablet) was added in each sample for the calculation of each type of the OM (see formulas below).

    Since AOM (amorphous or structureless OM) appeared flocculated in the thin section and its quantity was way more than other types of OM, which made it difficult to count, Adobe Photoshop image processing software was used to measure the area of the AOM. Palynomorphs and SOM (structured OM) were counted by numbers. The statistical formula used is as follows

    X1. AOM area concentration (μm2/50 g); S. AOM area (μm2); Cly. lycopodium spore concentration (27 637 pieces/50 g); Nly. counted lycopodium spores (pieces); X2. palynomorphs (or SOM) concentration (pieces/50 g); N. palynomorphs (or SOM) numbers.

  • Bulk geochemical and TOC concentration analyses were conducted on 54 samples distributed uniformly throughout the section to investigate redox conditions and bulk accumulation of OM in the sediment.

    TOC abundance was determined by a Carbon/Sulfur Analyzer CS-2000 after dissolution of powdered samples in 10% HCl and subsequent air drying. This was performed at the Wuxi Research Institute of Petroleum Geology, China, with an analytical precision better than ±0.2%.

    Major element abundances were measured by wavelength-dispersive X-ray fluorescence (XRF) analysis on fused glass beads using the XRF-1800 apparatus at Hubei Geological Research Laboratory, China. The analytical uncertainty is < 5%.

    Trace elements were measured using Agilent 7500a inductively coupled plasma mass spectrometry (ICP-MS) at the State Key Laboratory of Geological Processes and Mineral Resources, China University of Geosciences (Wuhan). Analysis of the international rock standards BHVO-2 and BCR-2 indicated that the analytical precision was better than 5% (Xu et al., 2012).

    Elements such as Al and Ti are proxies for the detrital siliciclastic fraction of sediments. Aluminum is a main component of aluminosilicates such as clay minerals, and Ti is a trace element in clay minerals and rutile (Tribovillard et al., 2006). In this study, elements are normalized by the ratios of each element to aluminium (X/Al), which precludes dilution effect by carbonate/authigenic minerals (Tribovillard et al., 2006). Major and trace elements and TOC are shown in Table S1.

  • In paleo-marine environments, Bio-Si equals the total silica in every sample deducting its terrigenous input, and therefore can be calculated as follows (Arsairai et al., 2016; Xiang et al., 2013; Ross and Bustin, 2009)

    Excess SiO2 represents Bio-Si. SiO2 (sample) and Al (sample) are the tested concentrations of SiO2 and Al in each sample. (SiO2/Al) PAAS represents ratio of SiO2/Al of the Post Archean Australian Shale (Taylor and Mclennan, 1985).

  • Fifteen samples from the Qiongzhusi Fm. which represented all changes in lithotype throughout the section were examined using Scanning Electron Microscopy (SEM) to characterize OM in-situ. Fresh surfaces of untreated rock samples were sprayed with carbon and microphotographs were taken under secondary electron mode (SE) at low voltage (3-5 kV) and backscatter electron (BSE) mode at 20 kV. In total 75 photographs of in-situ OM were taken. Energy spectrum measurements of the OM samples were performed using electron dispersive X-ray spectroscopy system (EDS) detectors to analyze their chemical components.

  • To document the variations in organic facies, three types of OM are classified according to the partition scheme suggested by Batten (1996) and Li and Batten (2005), which identifies three types of OM: AOM, SOM and palynomorphs. Under transmitted light, AOM is translucent or opaque, flocculated or gelatinous, and shows no identifiable biological structure or stable contour (occasionally cell-structure remnants are observed; Fig. 2). AOM is further subdivided into two types, namely, opaque AOM (Figs. 2a, 2b) which corresponds to black structureless aggregates, and transparent AOM which consists of flocculent or spongy aggregates (Fig. 2b). Gelatinous OM visible in Fig. 2c most likely corresponds to cyanobacterial mats (Li and Batten, 2005; Batten, 1996; Tyson, 1987). Mixed opaque and transparent AOM (plus tiny pyrite grains) is abundant in the measured section (Fig. 2d).

    Figure 2.  Representative images of AOM from the Qiongzhusi Fm., Shatan Section. (a) Opaque AOM in Sample NS-009; (b) opaque and transparent AOM and lycopodium spores in Sample NS-120; (c) transparent gelatinous AOM in Bed 1; (d) mixed AOM in Bed 1.

    All organic-walled micro fossils in the palynology slides are considered to be palynomorphs in the study. The results show that the overwhelming majority of palynomorphs are strongly carbonized, unidentifiable spherical acritarchs (Figs. 3a-3d), of which some are surrounded by extracellular polymeric substances (EPS) (Fig. 3a). The SOM are black cellulosic fragments with clear edges and structures (Figs. 3e-3h) derived from phytoplankton or macroalgae. Some are dark black opaque fragments (Figs. 3e, 3f) under transmitted light due to advanced degradation and/or oxidation (Graz et al., 2010). There are dark brown opaque fragments showing visible internal structures (Fig. 3g), and cuticular fragments (Fig. 3h) that display traces of unidentifiable internal structures under transmitted light. The advanced degradation stage imparts a dark brown to black color with fragment dimensions variable between 20 and 150 μm.

    Figure 3.  Representative images of palynomorphs and SOM in the Qiongzhusi Fm., Shatan Section. (a)-(d) Palynomorphs; (e)-(h) SOM; (e), (f) dark black cellulosic opaque fragments with clear edges; (g) dark brown corroded-outlined opaque fragments; (h) cuticular fragments.

  • Results of TOC and palynological OM are shown in Fig. 4, displaying obvious coincidences. Curves of palynological OM show the detailed information on the nature of sedimentary organic carbon, which helps in the assessment of the proportion of different types of OM.

    Figure 4.  Quantitative distribution of TOC and palynological OM in the Qiongzhusi Fm., Shatan Section.

    Estimated by area (the mean particle size of SOM and palynomorphs are 36.7 μm), total AOM in the whole section accounts for more than 80% of all the OM content, and in each bed, AOM content is much greater than that of other types of OM as well. In beds 1 and 2, all types of OM are relatively enriched in the whole section (Figs. 4b-4d).

    AOM accounts for approximately 72% of all OM content. In Bed 3, all types of OM decrease (Figs. 4b-4d). Palynomorphs and SOM in the lower part of Bed 3 (Stage 2) show higher values than in the upper part (Stage 3) (Figs. 4c, 4d). AOM accounts for 76.3% of all the OM content; which is still the dominant (Fig. 4b). In beds 4 and 5, all the OM content increases in values. AOM accounts for 92.3%. Palynomorphs and SOM content are also higher than that in other beds (Figs. 4b-4d). The total palynological OM is the summation of AOM, SOM and palynomorphs estimated by area (Fig. 4e). Its variation trend resembles the AOM, as AOM dominates the palynological OM content.

  • The analytical results are illustrated in Fig. 5. To gain more insight into the preservation of OM in relation to paleoenvironmental conditions, U/Th, V/Cr, Mo/Al and Corg/P are used as the redox proxies (Jones and Manning, 1994; Hatch and Leventhal, 1992) and Cu/Al, Ni/Al, Zn/Al are used as the paleo-productivity proxies (Shen et al., 2015; Verlaan, 2008; Meyers et al., 2005; Piper and Perkins, 2004). The interpretations of the trace metals of interest is based on previous work by Little et al. (2015), Algeo and Tribovillard (2009), Algeo and Lyons (2006), Tribovillard et al. (2006), Turgeon and Brumsack (2006), and Algeo and Maynard (2004). As shown in Fig. 5, in beds 1 and 2, the mean value of TOC is 2.5% (Fig. 5a), and the mean value of V/Cr is 2.42 (Fig. 5b) which are relatively high in the whole section. The values of other redox proxies, namely, U/Th, Corg/P and Mo/Al are uniformly high (Figs. 5c-5e). The mean values of the paleo-productivity proxies Ni/Al, Zn/Al and Cu/Al are 2.1 ppm/%, 7.4 ppm/% and 5.6 ppm/%, respectively (Figs. 5f-5h). In Bed 3, TOC contents drop to the lowest level of the sequence (0.45% average and 1.02% maximum) (Fig. 5a), and the average ratio of V/Cr decreases to 0.82 (Fig. 5b), U/Th, Corg/P and Mo/Al ratios drop at the same time (Figs. 5c-5e). While the paleo-productivity proxies Ni/Al Zn/Al and Cu/Al mean ratios respectively increase to 3.24 ppm/%, 13.06 ppm/% and 6.23 ppm/% (Figs. 5f-5h). In Bed 4, V/Cr, U/Th, Corg/P and Mo/Al ratios shows marked increase (Figs. 5b-5e), the mean value of V/Cr rise to 1.7 (Fig. 5b). The Ni/Al, Zn/Al and Cu/Al mean ratios are 2.74 ppm/%, 6.23 ppm/% and 5.09 ppm/%, respectively (Figs. 5f-5h), which show an inconspicuous decline. TOC contents show a sharp increase with an average of 1.9% (Fig. 5a). In Bed 5, the V/Cr ratio continues to increase to 2.35 on average with a maximum of 5.48 (Fig. 5b). U/Th and Mo/Al ratios increase as well (Figs. 5c-5e). Average values of Ni/Al, Zn/Al and Cu/Al are 5.7 ppm/%, 22.5 ppm/% and 14.1 ppm/%, respectively, showing a slightly increase (Figs. 5f-5h). TOC values keep stable, showing the mean value of 1.6% (Fig. 5a).

    Figure 5.  Stratigraphic distribution of TOC, redox and paleo-productivity proxies and total palynological OM content in the Qiongzhusi Fm., Shatan Section. V/Cr values on the left with the red line (< 2) indicate an oxic condition, in the part between the red and blue lines (2-4.25) indicate a dysoxic condition, and on the right of the blue line (> 4.25) indicate anoxic condition (Hatch and Leventhal, 1992).

  • To examine the relationship between Bio-Si and OM content, we first determine the origin of silicon in the investigated section. Th and Zr are two elements dominant in land-derived clastic fraction, whilst Si may have mixed origins, e.g., the quartz fraction of the clastic supply, bio-opal of siliceous organisms or hydrothermal silica (Adachi et al., 1986).

    As shown in Fig. 6b, in beds 1 and 2, Si/Al which represents the total silicon is experiencing a decline (Fig. 6b; blue arrow). This trend is revealed by the maximum value of 11.20 at the base of Bed 1, and the minimum value of 7.60 in the upper part of Bed 2. While the values of Th/Al and Zr/Al in beds 1 and 2 increase (Figs. 6c and 6d; red arrows), the maximum value and minimum value of Th/Al are 3.25 and 4.09, respectively, and the maximum value and minimum value of Zr/Al are 56.62 and 84.94, respectively. From the bottom of Bed 3 till the top of Bed 4, there is no significant variation of the Si/Al value (Fig. 6b) except for the peak value of 12.12 in the lower part of Bed 3. The mean values of Si/Al in beds 3, 4 are both 8.7. Regarding to Th/Al and Zr/Al, no strong fluctuation of the curves is observed in beds 3, 4 (Figs. 6c and 6d). The mean values of Th/Al in beds 3, 4 are 3.43 and 3.42 respectively, and the mean values of Zr/Al in beds 3, 4 are 69.48 and 68.64 respectively. In Bed 5, there is a coincident decline of Si/Al, Th/Al and Zr/Al in the lower part (Figs. 6b-6d; blue arrows). The mean values of Si/Al, Th/Al and Zr/Al are 8.87, 3.10 and 54.66 respectively in Bed 5. In our study, the Si/Al ratio does not follow the same vertical distribution as Th/Al or Zr/Al from Bed 1 to Bed 4 (Figs. 6b-6d), which suggests that Si does not share a common carrier phase with Th and Zr from land. In contrast, the curves of Si/Al, Th/Al and Zr/Al in Bed 5 are obviously coherent, representing terrestrial input of clastic quartz in the upper part of the section.

    Figure 6.  Stratigraphic distribution of TOC and the mineral proxies in the Qiongzhusi Fm., Shatan Section.

    In addition, the Fe-Al-Mn ternary plot (Fig. 7) shows that samples fall in part I, which corresponds to sediments that are not influenced by hydrothermal processes. The non-hydrothermal zone and the hydrothermal zone in the Al-Fe-Mn ternary plot (Fig. 7) are divided by the chemical characteristics of samples from various deposits, such as biogenic cherts, metalliferous deposits and pelagic oozes (see e.g., Adachi et al., 1986). The hydrothermal origin of silica in the study is thus ruled out. Hence, the contrasting behavior of Si points to a probable biogenic origin which would indicate the presence of siliceous organisms in the paleo-depositional environment, i.e., sponges or radiolarian (Maliva et al., 1989; Yamamoto, 1987) from Bed 1 to Bed 4. In Bed 5, Si may have mixed origins from both the clastic supply and the bio-opal of siliceous organisms.

    Figure 7.  A1-Fe-Mn diagram showing the effect of hydrothermal emanations in the Qiongzhusi Fm., Shatan Section.

    As shown in section 2.3, Bio-Si is represented by excess SiO2. In beds 1 and 2, the mean value of Bio-Si is 48.3% which is relatively high in the whole section (Fig. 6e). In Bed 3, the mean value of Bio-Si drops slightly to 38.4%. In Bed 4, values show an inconspicuous increase, and the mean value rises to 49.1%. In Bed 5, the Bio-Si content experiences a decline, and the mean value drops to 39.8% (Fig. 6e).

  • Variations in the extent of OM loading are commonly linked to stratigraphic or spatial shifts in redox conditions (Poulton et al., 2010; Tyson, 2005). Broadly, the redox variation expressed by our data resembles the pervasively oxygenated pattern of Early Cambrian inner-shelf oxygen minimum zones (OMZs) (Guilbaud et al., 2018).

    As shown in Fig. 5, the variation trend of the redox proxies, namely, V/Cr, U/Th, Corg/P and Mo/Al are accordant (Figs. 5b-5e). Four stages of redox evolution can be recognized (Fig. 5). Beds 1, 2 are the first stage, and the mean value of V/Cr is 2.42 (Fig. 5b), indicating hypoxic-anoxic conditions (Hatch and Leventhal, 1992). Bed 3 is the second stage where the average ratio of V/Cr is 0.82 (Fig. 5b) which coincides with the oxygenation event of Cambrian Stage 2 (oxygenated surface seawater in entire South China, Jin et al., 2016; Chen et al., 2015; Yang et al., 1999). In some deep-water sections in the Yangtze Plate, bottom waters are anoxic (Yang et al., 1999), the measured section is composed of shallow-marine sediments, and the bottom waters remain oxygenated. Bed 4 is the third stage, and the dominating redox state is still oxic with some values in the dysoxic zone (Hatch and Leventhal, 1992). Bed 5 is the fourth stage, with the mean ratios of V/Cr at 2.35 (Fig. 5b) indicating a change from oxic to anoxic condition (Hatch and Leventhal, 1992).

    The coupling between TOC and all the redox proxies is significant in our section (Figs. 5a-5e), indicating an obvious influence of redox conditions on TOC distribution. Moreover, Fig. 8 shows that there are significant correlations between TOC and the redox proxies in the measured section. The correlation coefficients between TOC and U/Th, Mo, Corg/P are RU/Th=0.70, RMo=0.64 and RCorg/P=0.77, respectively (Figs. 8a-8d). Therefore, we can conclude that organic carbon distribution is well related to paleo-oxygenation levels of the depositional environment. On the contrary, as illustrated in Fig. 5, it is observed that the stratigraphic distribution of Ni/Al, Zn/Al and Cu/Al does not match with that of TOC (Figs. 5f-5h). Beds 1 and 2 show the lowest paleo-productivity but high TOC contents (Fig. 5a), and the high paleo-productivity proxy values in Bed 3 are not accompanied by organic carbon enrichment (Figs. 5f-5h). In Bed 4, TOC and V/Cr, U/Th, Corg/P, Mo/Al show a corresponding increase (Figs. 5a-5e), whereas the productivity proxies Ni/Al, Zn/Al and Cu/Al only start to "take off" in Bed 5 (Figs. 5f-5h). In addition, cross-plots opposing TOC to paleo-productivity proxies, namely, Ni, Zn and Cu, show no statistically significant correlations, as the coefficients are RNi/Al=0.019, RCu/Al=0.016 and RZn/Al=0.084, respectively (Fig. 8e). These observations suggest that the marine productivity doesn't exert the primary control on organic carbon distribution.

    Figure 8.  Geochemical indices vs. TOC abundance from the Qiongzhusi Fm. in the Shatan Section.

    However, the correlation between TOC and V/Cr is high only for beds 1 to 4 where the redox conditions are relatively steady, with a correlation coefficient RV/Cr=0.64 (Fig. 8c). In Bed 5 where the redox conditions are fluctuating, no significant correlation between TOC and V/Cr is observed. In addition, the TOC, Zn/Al, Ni/Al and Cu/Al curves all show the same "depressed" distribution at around the middle part of Bed 5 (Figs. 5a-5h, In this case, paleo-productivity and TOC seem to be closely related in Bed 5. This observation implies that the low oxygen level is the prerequisite of organic carbon enrichment as discussed above, and restriction of TOC by the paleo-productivity could only be revealed during the deposition episode in which the redox condition is generally anoxic as in Bed 5.

    Our reexamining of the "productivity model" and the "preservation model" suggests that neither of them is complete enough to assess sedimentary OM distribution. It also reveals that the factors controlling OM distribution are multivariant. Any factors that shorten the OET of OM should be taken into consideration to improve the assessment model. The cognition of the paleoenvironment thus provides references that can be compared to other influencing factors, for example, the Bio-Si, and the origin and type of OM.

  • As discussed above, variations in TOC are linked to stratigraphic shifts in redox conditions. However, a closer look at the V/Cr curve indicates that many values in beds 4 and 5 fall into the oxic area (V/Cr < 2) (Fig. 5b). But as TOC and palynological OM values are relatively high in these beds (Figs. 5a, 5i), the oxic condition does not seem to have a negative effect on OM preservation. One possible interpretation is that the chemically-recalcitrant components related to the origin of OM might be an intrinsic factor of OM preservation. Although there is no direct evidence in this paper to show the selective preservation with OM abiotic condensation or aggregation, the discussion of the origin of OM based on the palynological OM study provides us insight to the early stages of OM preservation.

    As shown in Section 3.1, the measured section is characterized by abundant AOM contents that are dominated by amorphous aggregates, which was derived from bacteria, cyanobacteria, algae and fragments or detritus of macroalgae (Bhattacharya and Dutta, 2015; Moczydłowska and Willman, 2009; Li and Batten, 2005). A significant correlation is observed between TOC and AOM with a correlation coefficient of R=0.499 (Fig. 9a). The correlation coefficient between the palynomorphs and SOM is R=0.482 (Fig. 9b). And it is worth noting that all types of OM-constituents are related to some extent with TOC (Figs. 4a-4d). This could provide indirect evidences that both the micro phytoplankton and macroalgae are important OM precursors.

    Figure 9.  TOC and palynology OM correlation from the Qiongzhusi Fm. in the Shatan Section. (a) TOC abundance vs. AOM values; (b) palynomorphs values vs. SOM values.

    Previous studies show that algaenans produced by algae are independent long-chain aliphatic compounds with hydroxyl or ester functional groups, and the cutans in the cell wall are non-hydrolyzable components like cuticles of higher plants (Gelin et al., 1999). They are selectively protected and degradation-resistant (Obeid et al., 2015; Zhang et al., 2007). In addition, DOC (dissolved organic carbon) or synthesized TEP extracellular polymers) and EPS (extracellular polymeric substances) secreted by the phytoplankton and bacteria (Ding et al., 2015; Verdugo and Santschi, 2010; Flemming et al., 2007) could form three-dimensional biopolymer networks resistant to remineralization in the water which collide and condensate into polymeric AOM (Verdugo et al., 2004; Mecozzi et al., 2001; Ciglenečki et al., 2000; Lee and Wakeham, 1992). The abundance of macroalgae fragments in the Neoproterozoic to the Cambrian indicates the prosperity of both green algae and brown algae (Du and Tian, 1986). FTIR spectroscopic data and pyrolysis experiments prove that macroalgae in the Neoproterozoic-Cambrian interval produce biopolymers similar to algaenans, which promotes their preservation (Sharma et al., 2009).

    Anoxic conditions only have a strong effect on the preservation of the more labile (reactive) fraction of sedimentary OM (Burdige, 2006). Therefore, the chemically-recalcitrant components related to the origin of OM should be considered as an intrinsic factor of OM preservation especially when the conditions are not severely anoxic.

  • Researches show that the preservation of OM is affected by adsorption onto inorganic matrices, e.g., mineral particles of silica, calcium carbonate and clay minerals, which leads to physical protection of OM (Batten and Freeland, 2007). The carbonaceous mudstone and black shale in beds 1, 2 and 4 correspond to the water-deepening event (Fig. 1c), and TOC values and palynological OM content are relatively high in these Beds (Figs. 4a-4d). This indicates that the increased input of clay mineral caused by the sea-level rising has favored the physical protection of OM.

    What is more notably related to the OM preservation in the Shatan section in the study is the Bio-Si. It has been used as a proxy to indicate the OM enrichment in time and space in many studies (Lyle et al., 1988; Banahan and Goering, 1986).

    Figure 6 shows a significant coupling between TOC and Bio-Si distribution in the measured section. The red arrows show increases in both values and the blue arrows shows decreases (Figs. 6a, 6e). Moreover, Fig. 8f reveals a significant correlation between TOC and Bio-Si with the correlation coefficient of R=0.69. In contrast, the variation curves show no coincidence between TOC and Th/Al and Zr/Al (Figs. 6a-6d). Besides, cross-plots opposing TOC to Th/Al and Zr/Al show no statistically significant correlations, as the coefficients are RTh/Al= -0.105 and RZr/Al=0.099, respectively (Figs. 8g, 8h). It indicates that there is a strong link between the distribution of marine Bio-Si and the accumulation of organic carbon in the paleo-marine environment, and that the sedimentary organic carbon is not controlled by weathering intensity.

    In order to gain some direct insight to the bonding between OM and Bio-Si, SEM observation is carried out in the measured section. Large amounts of OM have been recovered in all of the investigated samples (Fig. 10). The in-situ images show that almost all the visible OM in the original rock samples are without fixed structure (Figs. 10a-10c). This type of OM observed under the SEM has been defined as dispersed OM previously (Zhang et al., 2017), which corresponds to the amorphous organic aggregates, or AOM in terms of palynological OM (Zhang et al., 2017; Li and Batten, 2005).

    Figure 10.  In-situ dispersed OM and the EDS spots from the Qiongzhusi Fm. in the Shatan Section. (a), (b) BSE images of dispersed OM in Bed 1 and Bed 4; (c) SE image of partial enlarged image of (b). Cnt. Counting rate of the X-ray; Kev. X-ray energy value of the element.

    The symbiotic relationship between phytoplankton and micro-siliceous organisms are believed to have favored OM preservation (Lu et al., 2018). The formation of biogenic amorphous silica by micro-organisms has been widely studied (Robinson and Sullivan, 1987; Volcani et al., 1981). Micro-siliceous organisms in Early Cambrian oceans, including radiolarians and sponges (Ma et al., 2019; Zhang and Feng, 2019; Chang et al., 2018, 2017; Cao, 2014; Wu et al., 2014; Braun et al., 2007; Xiao et al., 2005; Yuan et al., 2002; Zhang and Pratt, 1994), extracted dissolved silica from the surface layer of ocean water to form their skeletons (Dennett, 2002; Caron et al., 1995). As discussed above, AOM is derived from algae, acritarchs and bacteria (Bhattacharya and Dutta, 2015; Moczydłowska, 2011; Moczydłowska and Zang, 2006; Moczydłowska and Willman, 2009). The phytoplankton account for a large part of the total productivity of the ocean surface, and are likely to be used as food by siliceous skeleton planktons (Banahan and Goering, 1986).

    Table 1 shows the energy dispersing spectrum (EDS) values of the test points in the measured section. Seen from the high atomic percentage of C which ranges from 57.72% to 88.69% (blue dash box), the chemical characterization of OM is dominated by organic compounds. The values in the red dash box show high silicon content, after correction for the silica that could be involved in clay minerals (red dash box). As illustrated in Section 3.3., the origin of Si is determined to come from biogenetic silica. This in-situ observation suggests that the OM in our measured section is closely bonded to Bio-Si.

    C O Na Mg Al Si P S Cl K Ca Ti Fe
    a1 wt.% 80.10 5.56 2.81 10.38 1.13
    at.% 88.69 4.62 1.39 4.92 0.39
    a2 wt.% 70.29 17.97 2.36 0.59 1.45 3.08 1.96 2.30
    at.% 79.37 15.23 1.39 0.33 0.73 1.49 0.68 0.78
    a3 wt.% 57.98 16.98 0.50 3.74 0.71 10.24 0.53 2.99 3.07 2.63
    at.% 72.82 16.01 0.33 2.09 0.33 4.67 0.22 1.16 1.16 0.71
    b1 wt.% 61.08 14.29 4.10 0.70 1.54 11.54 0.71 0.25 0.73 5.07
    at.% 75.75 13.31 0.94 0.43 0.85 6.12 0.34 0.11 0.28 1.89
    b2 wt.% 52.67 16.72 1.57 0.90 6.27 18.16 3.00 0.71
    at.% 67.37 16.05 1.05 0.57 3.57 0.94 1.18 0.27
    b3 wt.% 66.42 22.39 0.49 0.49 1.70 5.77 1.57 1.00
    at.% 75.55 19.12 0.29 28.00 0.86 2.80 0.75 0.35
    c1 wt.% 52.64 18.85 1.42 5.25 12.36 2.85 2.23 0.49 3.89
    at.% 68.53 18.42 0.91 3.04 5.83 1.14 0.87 0.16 1.09
    c2 wt.% 62.85 21.28 0.81 3.66 9.57 0.06 2.13
    at.% 73.15 18.90 0.48 1.93 4.75 0.03 0.77
    c3 wt.% 44.21 28.40 1.49 1.09 4.50 11.29 0.92 1.18 0.58 2.13 4.22
    at.% 57.72 27.84 1.02 0.70 2.61 6.30 0.46 0.58 0.26 0.77 1.65

    Table 1.  EDS data of Fig. 10 (wt.%. mass percentage of the elements, at.%. atomic percentage of the elements)

    Additionally, AOM, the dominant organic content in our test, belongs to the category of "polymeric OM" (Li and Batten, 2005). The gelatinous silicon from the dead micro-siliceous organisms could have been involved in the process of OM polymerization, and from the gelatinous particulate organic aggregation of increased size and higher sinking rate (Lebrato et al., 2013, 2011), leading to lower OET of OM in the water column.

    Besides, mineral-organic interactions can protect unstable organic components, for example, shielding active protein from chemical attacks, like enzymatic attacks or abiotic acid hydrolysis (Zhu et al., 2014). FTIR and MAS-NMR spectroscopic study of Bio-Si shows that there is an important amount of OM that remains linked to silica, probably via Si-C or Si-N electrostatic bonds (Gendron-Badou et al., 2003). As a result, correlations between OM and Bio-Si are the ubiquity of many observations from the organic-rich siliceous depositions.

    This observation seems to be independent from other environmental factors. As mentioned above, the bonding between OM and Bio-Si is known to be a protecting agent for OM against decay (Vandenbroucke and Largeau, 2007). Consequently, Bio-Si could be an additional factor involved in OM preservation and accumulation, acting mainly as a physical protection factor.

  • In the study, we show how the distribution of OM is incompletely accounted for by the paleoenvironmental conditions, and that previous geochemical indicators of redox conditions and paleo-productivity are not the only driving force for sedimentary OM preservation.

    We emphasize herein that the qualitative and quantitative study of OM palynofacies could be used for OM source analysis and OM preservation evaluation, which substantially expands the function of the TOC on primary productivity assessment and the content of resources. It can be used even in settings where preservation of OM is poorer and are also applicable over a wide range of thermal maturity. The AOM derived from algae and bacteria in our test represents the prevailing organic facies in Early Cambrian continental margin. One reason for the good preservation of these organic forms is the selective protection inherent to their own constituent chemistry, which forms the main source of the TOC here.

    According to our results, Bio-Si is a proxy for OM enrichment. The activity of micro-siliceous organisms (in Early Cambrian, radiolarians and sponges) favors the preservation of sedimentary OM. It is relatively independent from paleoenvironmental conditions and has a positive effect on OM preservation. This also offers a realistic account of the OM preservation mode during the Proterozoic-Phanerozoic transition. These factors seem to have acted in (partial) synergy with the impact of some other mechanisms.

  • We are grateful to the editors for editorial handling and the anonymous reviewers for their critical and constructive comments that have greatly improved the quality of this paper. We express our sincere thanks to the palaeontology research team of the UMR 8198-Evo-Eco-Paleo, CNRS, for allowing the research stay of the senior author at Lille University. We thank Dr. Sebastiaan van de Velde for the language correction. We also thank Prof. Jianxin Yu, senior engineer Zhongbao Liu, and Dr. Thomas Harvey for inspiring discussion. This work was supported by the National Natural Science Foundation of China (No. 41430101) and the State Special Fund from Ministry of Science and Technology (No. 2017ZX05036002). The final publication is available at Springer via

    Electronic Supplementary Material: Supplementary material (Table S1) is available in the online version of this article at

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