Origin of chert nodules in the Ediacaran Doushantuo Formation black shales from Yangtze Block, South China

Marine and Petroleum Geology 114 (2020) 104227

Contents lists available at ScienceDirect

Marine and Petroleum Geology journal homepage: www.elsevier.com/locate/marpetgeo

Research paper

Origin of chert nodules in the Ediacaran Doushantuo Formation black shales from Yangtze Block, South China

T

Ping Gaoa,b,∗, Zhiliang Hec, Gary G. Lashd, Shuangjian Lib, Rongqiang Zhangb a

School of Energy Resources, China University of Geosciences (Beijing), Beijing, 100083, China Petroleum Exploration & Production Research Institute, SINOPEC, Beijing, 100083, China c Ministry of Science and Technology, SINOPEC, Beijing, 100728, China d Department of Geology and Environmental Sciences, State University of New York - Fredonia, Fredonia, NY, 14063, USA b

A R T I C LE I N FO

A B S T R A C T

Keywords: Silicification Carbonate replacement Anaerobic oxidation of methane Anoxia Organic matter

Chert nodules are widely developed in Ediacaran deposits of South China, yet the mechanism responsible for their formation remains disputed. Petrological and geochemical studies of chert nodules and host shale of the Ediacaran Doushantuo Formation of the Yangtze Block, South China, were carried out to more fully assess silicification of these deposits. Doushantuo chert nodules display a concentrically layered internal structure dominated by quartz and lesser concentrations of calcite, carbonate fluorapatite, organic matter, pyrite, and sphalerite. Individual layers become thinner from nodule centers to edges. Our results suggest that (1) Doushantuo host black shale accumulated under persistent anoxic bottom water conditions and (2) chert nodules formed as a consequence of the anaerobic oxidation of methane focused along sulfate-methane transition zones at shallow burial depth. Nodules likely originated from porous proto-nodules that consisted of 13C-depleted authigenic calcite produced by the anaerobic oxidation of methane. Decaying of organic matter within nodules helped to create porosity that was filled by authigenic calcite and also released phosphate and zinc ions to pore water, thus favoring precipitation of carbonate fluorapatite and sphalerite. Silica supplied principally from the diagenesis of clay minerals of the host shale cemented the nodules and replaced early formed calcite. The progressive infilling of residual porosity by 18O-depleted pore fluids associated with increasing burial depth is reflected in the strongly negative δ18Ocarb values documented from the nodules. Thus, the formation of Doushantuo chert nodules reflects a protracted diagenetic history initiated by bacterial sulfate reduction and anaerobic oxidation of methane close to and within the sulfate-methane transition zone and continued in association with the diagenetic release of silica from host shale as well hydrocarbon generation with increasing burial depth.

1. Introduction Chert nodules are common to modern and ancient marine sediments and sedimentary rocks (Maliva and Siever, 1989; Loi and Dabard, 2002; Xiao et al., 2010; Stefurak et al., 2014; Muscente et al., 2015; Qu et al., 2017). Three models accounting for the silicification of marine sediment and the formation of nodules have been postulated, including biological mediation, diagenetic replacement, and direct precipitation (Knauth, 1994; Loi and Dabard, 2002; Dong et al., 2009; Xiao et al., 2010; Marin-Carbonne et al., 2014; Wen et al., 2016; Gao et al., 2018). The absence of significant silica-secreting organisms during Precambrian time suggests that contemporary oceans contained higher concentrations of dissolved silica than modern oceans (Siever, 1992; Maliva et al., 2005; Geilert et al., 2014; Stefurak et al., 2014; Marin-

∗

Carbonne et al., 2014); thus, silicification would have occurred via inorganic pathways. However, its worth considering the degree to which abiotic silica precipitation was facilitated by the emergence of Siaccumulating bacteria near the close of the Precambrian (Dong et al., 2015). Direct precipitation of silica from seawater appears to have played a minimal role in silicification of Holocene and Phanerozoic deposits (Racki and Cordey, 2000; Maliva et al., 2005), with the main occurrences of siliceous hydrothermal deposits (Chen et al., 2009). Thus, the mechanism(s) responsible for silicification during the Precambrian remain poorly understood. In general, the formation of chert nodules is associated with a decreasing clastic input, which stabilizes the sediment-water interface (SWI) and diagenetic fronts within the sediment column (Loi and Dabard, 2002; Dabard and Loi, 2012). Indeed, the significance of

Corresponding author. School of Energy Resources, China University of Geosciences (Beijing), Beijing, 100083, China. E-mail addresses: [email protected], [email protected] (P. Gao).

https://doi.org/10.1016/j.marpetgeo.2020.104227 Received 4 April 2019; Received in revised form 27 November 2019; Accepted 5 January 2020 Available online 07 January 2020 0264-8172/ © 2020 Elsevier Ltd. All rights reserved.

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Fig. 1. (A) Simplified paleogeographic map showing the major geologic elements of the Yangtze Block during the Ediacaran and the location of the Songlin section; (B) lithologic column of the Lower Ediacaran Doushantuo Formation of the Songlin section (Jiang et al., 2011).

Songlin section is located ~20 km west of Zunyi City in Guizhou Province and is inferred to be deposited in a shelf lagoon at that time based on paleogeographic reconstruction (Fig. 1A). The DST comprises a succession of marine carbonate and shale deposits that are widely distributed across the Yangtze Block (Jiang et al., 2011). The DST of the Songlin section, approximately 80 m thick, rests unconformably on Nantuo Formation tillites (Fig. 1B). The DST is subdivided into four lithostratigraphic members. Member I, the stratigraphically lowermost unit, comprises ca. 5 m of 13C-depleted dolomite in a sharp contact with underlying Cryogenian diamictite of the Nantuo Formation (Fig. 1B). Member II consists of ca. 25 m of silty mudstone intercalated with dolomites (Fig. 1B). Member III is ca. 50 m thick and consists of black shale in the lower part of the unit and intercalated dolomite and shale in its upper part (Fig. 1B). Member IV is composed of ca. 7m-thick black shale intercalated with phosphorite in its upper part and contains abundant cm-sized chert nodules (Figs. 1B and 2). The DST is overlain by massive dolomite of the Dengying Formation (Fig. 1B). Stable carbon isotope chemostratigraphy of the DST exposed in the

reduced clastic sediment supply on chert nodule formation is revealed by their common association with condensed successions, including black shale deposits associated with transgressive system tract and basal highstand system tract successions (Loi and Dabard, 2002; Loi et al., 2010). Such a scenario of shallow diagenesis is consistent with observations of recent sediments that display strongly enriched concentrations of dissolved silica at 5–50 cm burial depth (Froelich et al., 1988; Khalil et al., 2007; Ehlert et al., 2016). The well-exposed Ediacaran Doushantuo Formation (DST) of South China contains abundant chert and phosphate nodules (Xiao et al., 2010; Jiang et al., 2011; Shen et al., 2011; Muscente et al., 2015; Xin et al., 2015; Zhao et al., 2016; Qu et al., 2017; Shang et al., 2018). Chert nodules hosted in DST dolomite intervals from the Yangtze Gorges area of South China, including the Julongwan section (Fig. 1), have been extensively studied (Xiao et al., 2010; Muscente et al., 2015; Zhao et al., 2016; Qu et al., 2017). Xiao et al. (2010) proposed that the silica precipitation was facilitated by localized pH change linked to sulfate reduction of organic matter and pyrite precipitation in the anoxic zone. Most recently, a complete and continuous sequence of silica phases from opal-CT to microcrystalline quartz has been discovered in the silica matrix of chert nodules, suggesting diagenetic transformation from opal-CT to quartz (Shang et al., 2018). However, it's questionable whether opal-CT, a metastable phase of silica, could be preserved in the silica matrix over a long period of geological time. Similar chert nodules were usually contained in DST black shale, but they have received relatively little attention. The paper reports results of a comprehensive integrated study utilizing petrological and geochemical data carried out on DST chert nodules hosted in black shale of South China. The principal goal is the further elucidation of silicification during the Precambrian.

2. Geological setting and studied stratigraphic section The South China is composed of Yangtze Block and Cathaysia Block (Fig. 1A). The Yangtze Block comprises four main sedimentary facies during the early Ediacaran, including inner shelf, intrashelf basin/shelf lagoon, shelf margin, and slope-basin facies from the northwest to southeast (Fig. 1A; Jiang et al., 2011; Xiao et al., 2012). The studied

Fig. 2. Outcrop photograph of chert nodules in DST black shales. The length of lens cap is 77 mm. 2

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

dispersive X-ray spectroscopy (EDS) system detectors (Oxford Instruments, UK) at the Analytical Laboratory of Beijing Research Institute of Uranium Geology (ALBRIUG). Imaging was performed on backscatter electron (BSE) mode at 15 kV and a working distance of ~5.0 mm under vacuum. Following removal of carbonate from two aliquots of sample powders with dilute HCl, total organic carbon (TOC) and total sulfur (TS) contents of one aliquot were measured using a Leco CS-230 analyzer with analytical precision better than 10%. A second aliquot was used for the measurement of stable carbon isotope compositions of organic matter (OM) using a Finnigan MAT 253 isotope mass spectrometer. Replicate analyses of studied samples and international standards (IAEA-600 Caffeine and USGS24 Graphite) were conducted with an analytical precision better than 0.1‰. This work was conducted at the State Key Laboratory of Petroleum Resources and Prospecting, China University of Petroleum, Beijing (SKLPRP-CUPB). Stable carbon and oxygen isotopic compositions of carbonate fractions of sample powders were measured at the ALBRIUG using the phosphoric acid treatment method that widely reported by previous authors (e.g., Cui et al., 2017; Lash, 2018). The Chinese national standard GBW04416 was tested routinely every 5 samples to monitor precision and accuracy; analytical precision was better than 0.2‰. Results are reported in δ notation in per mille (‰) relative to the Vienna-Pee Dee Belemnite (V-PDB) standard. Major, trace, and rare earth element (REE) concentrations were measured at the ALBRIUG. Major oxides of host rocks were determined on a PANalytical AxiosmAX X-ray fluorescence analyzer on fusion glasses of a 1:10 ratio of sample to Li2B4O7 in accordance with Gao et al. (2016). Major element (Fe and Al) concentrations of chert nodules were determined on a PerkinElmer 5300DV ICP-OES using sample powders digested in a mixture of super-pure HNO3+HCl + HF in high pressure Teflon cups as outlined by Gao et al. (2016). Trace and REE concentrations of all samples were determined on a PerkinElmer NexION300D ICP-MS using sample powders digested as above. Analytical precision was better than 5% for major and trace elements and 10% for REE. Enrichment factors (EFs) of selected elements were calculated as X-EF = [(X/Al)sample/(X/Al)PAAS], where X and Al represent weight percent concentrations of elements X and Al of the sample and the PostArchean Australian Shale (PAAS) standard (Taylor and McLennan, 1985). In order to avoid the artificial effect, Ce and Eu anomalies were calculated according to Lawrence et al. (2006); i.e., Ce/Ce* = CeN/ ((PrN)2/NdN), and Eu/Eu* = EuN/(SmN × (SmN/NdN)1/2), where N refers to elemental concentrations normalized to PAAS (McLennan,

Jiulongwan and nearby sections (Fig. 1A) has been described by multiple authors (Jiang et al., 2011; McFadden et al., 2009). The δ13Ccarb profile of the DST generally displays three Ediacaran Negative (EN) excursions: EN1 that corresponds to dolomite of Member 1; EN2 found near the boundary of members II and III and EN3 spanning the upper part of Member III into Member IV (Jiang et al., 2011; Cui et al., 2016, 2017) (Fig. 1B). EN2 is thought to be associated with the widespread Gaskiers glaciation (Condon et al., 2005; Tahata et al., 2013). The longlasting EN3 excursion is believed to be correlated with the global Shuram Excursion recognized worldwide (Jiang et al., 2007; McFadden et al., 2009; Cui et al., 2016). Since the DST of Songlin section was deposited in the strongly stratified shelf lagoon facies, carbonate deposits in Member II and Member III may be formed below the chemocline and may have significant negative δ13Ccarb values (Fig. 1B; Jiang et al., 2011). The age of the DST is constrained between 635.2 ± 0.6 Ma and 551.1 ± 0.7 Ma based on U–Pb dating of zircons collected from volcanic ash beds near the base and top of the formation, respectively (Condon et al., 2005). The DST appears to comprise two and one-half second-order sequences (SS) of ca. 35 Ma for SS1 (635-600 Ma), 35 Ma for SS2 (600-565 Ma), and ca. 14 Ma for TST deposits of SS3 (565551 Ma) (Fig. 1B; Yang et al., 2015). Abundant chert nodules of the bottom of Member IV are closely associated with a condensed interval of black shale (Figs. 1B and 2). The nodule-bearing strata represent early TST deposits that accumulated no later than 565 Ma. 3. Samples and analytical methods 3.1. Samples Ten chert nodule samples and seven host rock samples were collected from the Songlin section (Figs. 1 and 2). Detailed information of the studied samples is provided in Table 1. All samples were prepared for analysis by grinding to a 200 mesh size in an automatic agate mortar. 3.2. Analytical methods Optical observation of polished thin sections was performed on a Zeiss Axioskop 40 A pol microscope at the Laboratory of Structural and Sedimentological Reservoir Geology, SINOPEC. Thin sections were coated with platinum for analysis using a FEI Nova Nano SEM450 scanning electron microscope (SEM) equipped with X-MaxN electron Table 1 Bulk geochemical parameters of DST chert nodules and host shale. Sample

Lithology

Locationa (m)

Size (cm × cm)

TOC (%)

TS

δ13Corg (‰ V-PDB)

δ13Ccarb

δ18Ocarb

SL-54-P1 SL-55-P1 SL-55-P2 SL-56-P1 SL-56-P2 SL-56-P3 SL-60-P1 SL-60-P2 SL-60-P3 SL-60-P4 SL-53 SL-54 SL-55 SL-55S SL-56 SL-56S SL-57

Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Chert nodule Black shale Black shale Black shale Black shale Black shale Black shale Argillaceous dolomite

4.65 5.15 5.15 5.25 5.25 5.25 5.45 5.45 5.45 5.45 4.4 4.65 5.15 5.15 5.25 5.25 5.65

1.5 1.2 1.6 1.8 2.4 2.5 1.9 1.8 1.5 1.4 / / / / / / /

1.3 1.4 1.9 1.6 0.8 0.7 3.1 3.0 1.0 1.2 3.0 3.1 3.4 3.3 4.0 4.3 2.3

0.09 0.2 0.3 0.08 0.03 0.5 0.4 0.3 0.4 0.2 2.7 2.0 1.5 1.3 2.5 3.9 0.3

ND −31.2 −31.2 ND ND ND −30.2 −31.7 −31.9 −31.3 −32.0 −31.8 −31.9 −31.7 −31.9 ND −31.2

−4.8 −3.8 −3.4 −5.7 −3.0 −3.3 −3.7 −3.5 −3.4 −2.8 −1.8 −1.8 −1.1 −1.1 −4.5 −4.5 −9.4

−16.0 −12.5 −14.2 −15.1 −14.8 −13.9 −15.5 −13.8 −12.9 −14.0 −5.5 −4.9 −5.9 −5.1 −12.1 −10.2 −3.6

× × × × × × × × × ×

1.6 2.1 1.9 1.8 2.4 3.2 1.9 1.8 1.5 1.5

/, not available; ND, not detected. a , The location is based on the distance to the boundary between Doushantuo and Dengying Formations. 3

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

relict calcite, surrounded by OM and microquartz (Figs. 4F–I and 5B). Larger masses (as much as 1.5 mm in diameter) appear to have formed by the coalescence of smaller masses (Fig. 4H). Amorphous solid bitumen (SB) is observed in intercrystalline pores within crystals of quartz and calcite (Fig. 4H). Finally, pyrite framboids 2–10 μm in diameter are randomly distributed throughout nodule cores (Fig. 5A and B). Core-rim contacts generally display a zigzag shape similar to the crystal form of calcites (Figs. 4I and 5C). Rims are generally thinner and more mineralogically varied relative to cores, including calcite, quartz, SB, CFA, metal sulfides (pyrite and sphalerite), and clay (Figs. 4I and 5C–K). Pyrite, the dominant metal sulfide is present in four forms, including massive (Fig. 5D and G), globular (Fig. 5E and F), framboidal (Fig. 5E), and subhedral (Fig. 5H). Massive pyrite is closely associated with CFA and SB (Fig. 5G). Nano-scale globular pyrite, commonly filled with SB (Fig. 5E), is present as aggregates (Fig. 5E) and chain-like bodies (Fig. 5F). Pyrite framboids are generally < 5 μm in diameter (Fig. 5E, G and I) and are closely associated with SB (Fig. 5E) and massive sphalerite (Fig. 5I). Subhedral pyrite is generally associated with massive sphalerite and CFA (Fig. 5H and I). Massive sphalerite is associated with pyrite and CFA (Fig. 5H–J). CFA is concentrated in nodule rims but is more abundant in host rocks than in nodules (Fig. 5D). It is noteworthy that clay minerals in nodule rims are found as the stacked sheets (Fig. 5K). 4.1.2. Host rocks The host shale of the chert nodules comprises quartz, dolomite, feldspar (including K-feldspar, albite, and orthoclase), CFA, metal sulfides (pyrite and sphalerite), SB, and few muscovites (Figs. 3B and 5D, L).

Fig. 3. Representative XRD pattern of DST (A) chert nodule and (B) host shale; Q = quartz; Cal = calcite; Dol = dolomite; CFA = carbonate fluorapatite; Mu = muscovite; Na = albite; Or = orthoclase; Py = pyrite.

4.2. Nodule geochemistry

1989).

Total organic carbon (TOC), total sulfur (TS), δ13Corg, δ13Ccarb, and δ Ocarb values of chert nodules and host rocks are presented in Table 1. TOC of chert nodules ranges from 0.7 to 3.1 wt% (average = 1.6 wt%; n = 10), less than that of host rocks (2.3–4.3 wt%; average = 3.3 wt%; n = 7). TS of chert nodules is relatively low, ranging from 0.03 to 0.5 wt% (average = 0.25 wt%; n = 10), markedly less than that of host shale (0.3–3.9 wt%; average = 2.0 wt%; n = 7). The δ13Corg of chert nodules ranges from −31.9 to −30.2‰ (average = −31.3‰; n = 6), approximately equal to that of the host rocks (average = −31.8‰; n = 6) (Fig. 6A). δ13Ccarb of chert nodules varies from −5.7 to −2.8‰ (average = −3.7‰; n = 10), which is slightly lower than that of the host shale (−4.5 to −1.1‰; average = −2.5‰; n = 6), excluding an outlier of argillaceous dolomite sample that may preserve the record of the Shuram Carbon Negative Excursion (Cui et al., 2016, 2017) (Fig. 6B). δ18Ocarb value of chert nodules are generally less than −10‰, much depleted in 18O relative to host rock (average = −6.8‰; n = 7; Fig. 6B). 18

4. Results 4.1. Petrological and mineralogical features 4.1.1. Chert nodules Chert nodules are randomly distributed throughout black shale of DST Member IV (Figs. 1B and 2). Nodules display a variety of shapes, including globular, ovoid, and oblate, and range from 1.2 to 3.2 cm in diameter. Coalesced smaller nodules are rarely observed (Fig. 2). The contacts of nodules and host shale are sharp and readily clearly observed. The wrapping up by laminated host shale (Fig. 2) suggests that nodules formed in the early diagenetic history of these deposits prior to sediment compaction. XRD analysis reveals that DST chert nodules are composed principally of quartz, followed by small amounts of calcite, dolomite, carbonate fluorapatite (CFA), and pyrite (Fig. 3A). Spherical phosphate nodules (Fig. 4J–L) dominated by CFA and OM, which is commonly rimed by quartz (Fig. 4J and K), can be found within the black shale. The spherical shapes of these nodules suggest that they, too, were formed in association with early diagenesis, before the host sediment experienced mechanical compaction. Optical and SEM observations reveal that nodules are generally composed of two distinct layers (Fig. 4A and B). Nodule cores are dominated by variable amounts of microcrystalline quartz and calcite, with rare occurrence of OM (Fig. 4A–I). Nodule rims are composed mainly of calcite, OM, and quartz (Fig. 4A and B). Increasing nodule diameter is accompanied by an increasing thickness of nodule core. Silicification of nodule cores appears to have been strongest along the edges of calcite crystals (Fig. 4E). Relict calcite and ghosts of calcite precursors are common in nodule cores (Figs. 4C–E and 5A). Microfossils, perhaps including acanthomorphic acritarchs, are well preserved in the silica and CFA matrix of chert nodules (Figs. 4C and D and 5B). Nodules cores display masses of megaquartz, some containing

4.3. Element geochemistry 4.3.1. Major and trace elements Major and trace element concentrations of chert nodules and host rocks are presented in Table 2. Major elements, especially Fe and Al, of chert nodules and host rocks display depleted elemental concentrations relative to PAAS (Table 2). Host rocks are enriched in the redox-sensitive trace elements Mo, U and V relative to PAAS whereas chert nodules are characterized by depleted redox-sensitive element concentrations (Table 2). Chert nodules and host rocks are strongly enriched in Zn relative to PAAS (Table 2). 4.3.2. Rare earth elements Rare earth element (REE) and Y concentrations and associated parameters of analyzed chert nodule and host rock samples are 4

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Fig. 4. Optical photographs showing textures of DST chert nodules and host rocks (PPL = plane-polarized light; CPL = cross-polarized light). (A) chert nodule dominated by microquartz (MiQ) and calcite (Cal) in its core; the nodule rim is enriched in OM; (B) calcite nodule has been moderately silicified mainly in its core; (C) strongly silicified calcite of a nodule core; some calcite fragments have been preserved. Inset illustrates detail of a preserved soft body microfossil consisting of microquartz and OM in the core and megaquartz (MeQ) in the rim; (D) globular microfossil (outlined by the yellow dashed line) composed of MiQ and small amount of OM in the core and MeQ in the rim. Some relicts of finely-crystalline calcite are indicated by yellow arrows; (E) silicified calcite rimmed by quartz (Q); (F, G) oblateshaped silica mass contained within a chert nodule composed of coarsely-crystalline quartz and calcite relicts; the host nodule is enriched in OM and composed principally of finely crystalline quartz; (H) a large silica mass that may have formed by coalescence of smaller silica masses. A small amount of solid bitumen (SB) is present within intercrystalline pores of the silica mass; (I) the rim of a chert nodule composed mainly of calcite and SB. Note the large silica mass consisting MeQ and OM; (J) two phosphate nodules (yellow dashed circles) and a chert nodule (red dashed circle) within black shale; (K) detail of the phosphate nodule shown in Fig. 4J composed of cryptocrystalline CFA and OM; its rim has been silicified; (L) detail of the chert nodule shown in Fig. 4J. The presence of CFA suggests that the precursor nodule was a phosphate nodule. . (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

5

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Fig. 5. Backscattered electron micrographs showing textures of DST chert nodules and associated host shale. Areas delineated by red stars and solid squares numbered 1–8 were analyzed by EDS; (Cal = calcite; Q = quartz; Py = pyrite; CFA = carbonate fluorapatite; Sph = Sphalerite; SB = solid bitumen; K–F = Kfeldspar). (A) inferred relict calcite (yellow dashed areas) and pyrite framboids (red arrows and inset) in the silica core of a chert nodule; (B) a CFA mass (delineated by yellow dashed lines) and minor pyrite within a silica nodule core. The CFA mass contains two microfossils (probably acanthomorphic acritarchs) highlighted by the white dashed square; (C) the core of a chert nodule is dominated by quartz and contains a small globular CFA mass (white arrow). The nodule rim is characterized by a more varied mineral composition that includes calcite and quartz and subordinate solid bitumen (SB) and sphalerite (Sph). Quartz appears to be concentrated adjacent to SB. The boundary of the core and rim presents a zigzag shape similar to the crystal form of calcite; (D) massive Sph and Py as well as amorphous SB within a nodule rim. The host shale contains massive CFA and abundant pyrite. The boundary of the nodule and host rock is indicated by the yellow dashed line; (E) framboidal and globular pyrite associated with SB within a nodule. Globular pyrite appears to have preserved porosity; (F) SB associated with quartz and chain-like Py, which likely formed by aggregation of numerous nm-scale globular Py masses; (G) massive Py associated with Q, CFA, and SB within a chert nodule. Single pyrite framboid 5.347 μm in diameter and clay minerals can be observed in the silica matrix of a nodule; (H) massive Sph, subhedral Py, and amorphous SB within the silica matrix of a nodule; (I) massive Sph associated with a framboidal Py in the silica matrix of a nodule; (J) co-mingled Sph and CFA; (K) sheet-like clay minerals (probably illite-smectite mixed layer) within the rim of a nodule; (L) host shale composed principally of quartz, K-feldspar (K–F), CFA, Py, and SB. . (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

opposite pattern (Fig. 7A). Most chert nodules exhibit slightly negative Ce anomalies (Ce/Ce* = 0.80–1.00), slightly positive Eu anomalies (Eu/Eu* = 1.06–1.50), and low Y/Ho ratios (34.0–43.8). ΣREE values of the host shale are greater than those of chert nodules, varying from

presented in Table 3. Total rare earth elements (ΣREE) of chert nodules range from 5.43 to 33.7 ppm (average = 18.3 ppm; n = 10). REE concentrations normalized to PAAS largely exhibit a pattern of depletion of lighter elements (Fig. 7A), whereas few samples exhibit an 6

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Fig. 6. Cross plots of δ13Ccarb versus (A) δ18Ocarb and (B) δ13Corg.

Sulfidic or euxinic bottom-water conditions favor the formation of syngenetic framboids within the water column (Wilkin et al., 1996). Upon growing to a threshold diameter, these framboids will settle down to the SWI and subsequently be buried. Thus, syngenetic framboids are generally smaller and in a narrower size range compared to diagenetic framboids population (Wilkin et al., 1996; Wignall and Newton, 1998). Rickard (2019), based on a meta-analysis of sedimentary pyrite framboid size-frequency distributions, suggests that the mean diameter of modern syngenetic framboids formed within euxinic water columns is 4.7 μm and that of diagenetic framboids is 6.7 μm. Analyzed framboids of DST chert nodules and host shale display a mean diameter of < 4.7 μm and a standard deviation of < 2.0 μm (Fig. 8), suggesting that deposition of DST Member IV black shale was dominated by euxinic bottom water conditions (Wilkin et al., 1996; Wignall and Newton, 1998). The presence of a small number of large (> 7.0 μm) framboids (Fig. 8) may reflect brief episodes of non-sulfidic bottom water conditions (Wignall and Newton, 1998). Sedimentary Mo has been widely used as a proxy for bottom water redox potential owing to its generally strong enrichment in fine-grained organic-rich sediment deposited under oxygen-depleted conditions (Algeo and Lyons, 2006; Tribovillard et al., 2006; Scott and Lyons, 2012). In general, Mo concentrations in the range of 2–25 ppm imply anoxic conditions whereas its concentrations > 100 ppm can be interpreted to reflect deposition under persistent euxinic conditions in an

14.9 to 116 ppm (average = 78.6 ppm, n = 7). REE concentrations of host shale display modestly diminished lighter element concentrations, minimally negative Ce anomalies (Ce/Ce* = 0.63–0.90), positive Eu anomalies (Eu/Eu* = 0.91–1.53), and low Y/Ho ratios (31.8–35.7) (Fig. 7B). 5. Discussion 5.1. Oceanic conditions associated with accumulation of the Doushantuo shale Pyrite in organic-rich sediments is generally occurred in two general forms, cubic to pyritohedron crystals, and spheroidal aggregates of pyrite microcrystals (framboids) (Powell et al., 2003). The size distribution of pyrite framboids in modern and ancient sediments had provided a potential tool to distinguish anoxic conditions from oxicdysoxic conditions (Wilkin et al., 1996; Wignall and Newton, 1998; Rickard, 2019). Sedimentary pyrite framboids can be formed as either diagenetic or syngenetic phases, depending on bottom-water redox conditions. Under oxic-suboxic bottom-water conditions, the redox interface separating non-sulfidic from sulfidic water and the position of framboid formation are located within the sediment column, perhaps a few cm below the SWI (Wilkin et al., 1996). Such conditions favor the formation of diagenetic framboids in the pore water of sediments.

Table 2 Selected major and trace element concentrations for DST chert nodules and host shale. Sample

Al

Fe

Mo

U

V

Zn

Ba

Mo/TOC

EF-Mo

EF-U

SL-54-P1 SL-55-P1 SL-55-P2 SL-56-P1 SL-56-P2 SL-56-P3 SL-60-P1 SL-60-P2 SL-60-P3 SL-60-P4 SL-53 SL-54 SL-55 SL-55S SL-56 SL-56S SL-57 PAASa

0.21 0.26 0.29 0.38 0.15 0.28 0.20 0.26 0.40 0.41 3.43 3.05 2.17 2.02 3.36 3.36 0.45 10.01

0.29 0.15 0.17 0.60 0.21 0.46 0.27 0.48 0.25 0.32 2.27 1.57 1.24 0.83 1.87 1.86 0.41 5.03

0.80 0.60 1.10 0.92 0.38 0.48 1.23 1.00 0.91 1.00 4.02 4.12 3.40 3.07 6.76 5.70 1.99 1.00

0.47 0.84 0.73 0.88 0.50 0.37 0.75 0.69 1.21 1.12 3.87 4.54 4.37 4.50 6.28 4.35 1.81 3.10

2.41 6.06 6.95 4.12 0.46 2.71 5.42 3.94 5.89 11.0 44.2 47.3 36.8 34.4 42.7 43.9 15.6 150

145 543 1040 130 64.7 230 180 635 898 1480 160 142 148 113 514 311 56.5 85.0

129 236 115 187 114 109 82.0 149 178 265 1130 888 559 588 826 834 155 650

0.60 0.43 0.57 0.59 0.48 0.66 0.40 0.33 0.94 0.86 1.32 1.34 0.99 0.93 1.67 1.32 0.87 /

37.2 23.1 38.5 24.3 24.5 17.0 61.6 38.6 23.0 24.2 11.7 13.5 15.7 15.2 20.1 17.0 44.3 /

7.14 10.3 8.25 7.55 10.5 4.17 12.2 8.65 9.87 8.74 3.64 4.81 6.50 7.19 6.04 4.18 13.0 /

/, not available. a , PAAS from Taylor and McLennan (1985). 7

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Table 3 REE concentrations and associated parameters for DST chert nodules and host shale. Sample

La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Y

∑REE

Y/Ho

Ce/Ce*

Eu/Eu*

SL-54-P1 SL-55-P1 SL-55-P2 SL-56-P1 SL-56-P2 SL-56-P3 SL-60-P1 SL-60-P2 SL-60-P3 SL-60-P4 SL-53 SL-54 SL-55 SL-55S SL-56 SL-56S SL-57

1.49 8.96 6.69 2.19 1.15 1.72 4.46 4.69 6.53 4.49 16.5 16.4 9.82 13.7 21.4 19.6 3.05

2.57 14.8 11.3 3.76 1.62 2.99 7.32 7.46 10.2 7.39 31.4 29.0 18.0 24.0 36.4 33.0 4.76

0.36 1.57 1.23 0.59 0.23 0.46 0.85 0.91 1.47 0.88 4.61 4.21 2.81 4.03 5.60 5.15 0.70

1.62 5.48 4.32 2.56 1.12 2.02 3.13 3.46 6.25 3.47 20.8 19.0 12.5 14.8 25.0 22.8 3.12

0.36 0.75 0.65 0.56 0.21 0.45 0.51 0.63 1.32 0.65 4.51 4.12 2.78 3.50 5.43 4.95 0.61

0.092 0.20 0.16 0.16 0.062 0.11 0.11 0.17 0.34 0.18 1.02 1.11 0.59 0.74 1.27 1.14 0.20

0.34 0.68 0.58 0.63 0.25 0.45 0.50 0.60 1.37 0.66 4.50 4.35 2.84 3.47 5.46 4.79 0.63

0.066 0.10 0.10 0.11 0.049 0.082 0.085 0.10 0.25 0.12 0.85 0.82 0.53 0.60 1.03 0.93 0.12

0.39 0.50 0.49 0.65 0.30 0.48 0.44 0.54 1.34 0.69 4.63 4.55 2.93 3.20 5.88 5.23 0.68

0.079 0.093 0.097 0.13 0.06 0.097 0.091 0.11 0.26 0.14 0.93 0.94 0.59 0.63 1.17 1.04 0.14

0.22 0.26 0.29 0.35 0.18 0.26 0.26 0.30 0.69 0.40 2.44 2.60 1.58 1.68 3.19 2.86 0.40

0.037 0.041 0.046 0.059 0.026 0.043 0.041 0.048 0.11 0.069 0.40 0.41 0.26 0.27 0.53 0.45 0.065

0.24 0.26 0.29 0.33 0.16 0.23 0.24 0.27 0.58 0.42 2.14 2.25 1.50 1.58 2.90 2.70 0.41

0.027 0.040 0.040 0.051 0.021 0.031 0.033 0.039 0.082 0.063 0.29 0.32 0.21 0.22 0.43 0.37 0.053

2.84 3.36 3.37 4.83 2.63 3.30 3.22 3.88 9.49 4.85 30.3 33.3 18.8 20.5 38.7 36.1 5.04

7.87 33.7 26.3 12.1 5.43 9.41 18.1 19.3 30.8 19.6 95.0 90.1 57.0 72.4 116 105 14.9

36.0 36.1 34.7 38.0 43.8 34.0 35.4 36.6 36.2 34.4 32.6 35.4 31.8 32.6 33.1 34.7 35.7

0.92 0.95 0.93 0.80 1.00 0.82 0.93 0.91 0.85 0.96 0.89 0.90 0.82 0.63 0.84 0.82 0.89

1.15 1.50 1.29 1.26 1.42 1.06 1.09 1.31 1.18 1.34 1.01 1.20 0.94 0.91 1.04 1.03 1.53

Fig. 7. PAAS normalized REE and Y values; (A) DST chert nodules and normal seawater (Elderfield and Greaves, 1982); (B) host shale of the Songlin section (PAAS from McLennan, 1989).

chert nodules deposited under anoxic bottom water conditions in a strongly restricted marine environment.

unrestricted ocean (Scott and Lyons, 2012). Intermediate Mo enrichments (25–100 ppm) may reflect either intermittent euxinia or Mo depletion in a restricted basin (Algeo and Lyons, 2006; Scott and Lyons, 2012). DST black shales display slight enrichment of Mo (2.0–6.8 ppm; average = 4.2 ppm; n = 7), suggesting deposition under anoxic bottom conditions. Moreover, the patterns of Mo-TOC co-variation can provide valuable information about paleohydrographic conditions, especially the degree of watermass restriction, but only under anoxic conditions (Algeo and Lyons, 2006; Algeo and Rowe, 2012). In sulfidic silled basins, aqueous Mo concentrations can be decreased due to Mo uptake in sediments and insufficient Mo replenishment from deep water (Algeo and Lyons, 2006; Algeo and Rowe, 2012), thus leading to a systematic decrease in Mo/TOC ratios with increasing restriction. In our study, Mo/TOC of DST black shales shows the low values ranging from 0.9 to 1.7 (Table 2), which is much lower than that of the Black Sea (4.5 ± 1; Algeo and Lyons, 2006), suggesting strong restriction of watermass. Algeo and Tribovillard (2009) and Tribovillard et al. (2012) described the use of patterns of authigenic U and Mo covariation as a means of discriminating redox conditions of diverse marine depositional systems. Mo-EF values of the analyzed DST samples plot close to modified seawater values and define a trend of increasing U enrichment with increasing Mo enrichment (Fig. 9), suggestive of anoxic bottom water conditions (Algeo and Tribovillard, 2009). In summary, the DST Member IV black shales that host abundant

5.2. Microenvironment of nodule formation The association of nano-scale hollow globular pyrite and very small pyrite framboids in chert nodules (Fig. 5E and F) suggests that pyrite was formed in association with bacterial activity within an anoxic-euxinic water column. However, the presence of large framboids (> 5 μm; Figs. 5A, G and 8), subhedral pyrite crystals (Fig. 5H), and pyrite masses (Fig. 5G) may reflect the slow formation of pyrite associated with BSR within the reducing pore water environment related to the sinking of the chemocline into the sediment column (Wilkin et al., 1996; Wignall and Newton, 1998). The Ce anomaly serves as a redox proxy for marine sediment when REE can be demonstrated to be primarily of hydrogenous origin (Chen et al., 2015; Zhang et al., 2016). The Y/Ho ratio is useful for distinguishing REE of hydrogenous origin from that derived from lithogenic sources. ΣREE values of analyzed DST nodules are plotted in the left (i.e., lower ΣREE concentrations) of host rock values and the ideal hydrogenous-lithogenous trend of the Y/Ho versus ΣREE crossplot (Fig. 10). The strongly depleted ΣREE concentration of the studied nodules likely reflects the diagenetic loss of REE by the replacement of authigenic quartz. It is noteworthy that Y/Ho values of the nodules 8

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Fig. 10. Y/Ho versus ΣREE for DST chert nodules and host shale of the Songlin section. The red curve represents a mixing trend between the hydrogenous (modern seawater) component and terrigenous (lithogenous) component of an average upper continental crust composition (McLennan, 2001); percentages provided as the amount of hydrogenous component (Chen et al., 2015; Zhang et al., 2016). Chert nodules and host shale contain 20–30% hydrogenous REE. (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

Fig. 8. Size frequency distribution of pyrite framboids in the DST host shale (based on analysis of samples SL-55S, SL-56 and SL-57).

(Steiner et al., 2001), (2) strongly reducing pore water environment (Martinez-Ruiz et al., 1999; Xin et al., 2015), or (3) Ba interference during ICP-MS measurement (Dulski, 1994). The weak positive covariance of Eu and Ba concentrations of DST nodules (R2 = 0.27; n = 10) rules out the role of Ba interference. The lack of direct evidence of hydrothermal alteration of the studied deposits suggests that the positive Eu anomalies may reflect the anoxic microenvironment of formation. 5.3. Precipitation of authigenic calcite associated with nodule formation Authigenic calcite, characterized by lower δ13Ccarb values and extremely negative δ18Ocarb values (Table 1; Fig. 6B), is an important component in DST chert nodules, which records the diagenetic history of the nodules. 5.3.1. 13C-depletion of nodules The presence of 13C-depleted authigenic carbonate in a sedimentary deposit may be attributed to (1) anaerobic oxidation methane (AOM; Reeburgh, 1976; Niewöhner et al., 1998; Hinrichs et al., 1999; Borowski et al., 1999; Cui et al., 2016; Lash, 2015a, 2018), (2) thermal decarboxylation of OM (Raiswell and Fisher, 2000), and (3) thermochemical oxidation of methane (TOM; Hu et al., 2018). It is noteworthy that δ13Ccarb values of chert nodules are generally lower than those of host rocks (Fig. 6A). Thermal decarboxylation of OM in association with burial would be expected to produce 13C-depleted carbonate throughout the succession rather than confined to nodules. TOM is generally induced by exposure of high-valence metal oxides (i.e., Mn/ Fe oxides) to high temperatures (Hu et al., 2018). However, these metals in the DST shales are present as low-valence sulfides (Fig. 5D and L). Thus, thermal decarboxylation of OM and TOM can be ruled out as formation mechanisms of the DST nodules. The formation of 13C-depleted authigenic nodular and concretionary carbonate in association with BSR and AOM has been documented from modern and ancient sedimentary deposits (Raiswell, 1988; Cui et al., 2016; Lash, 2015a, 2018; Chuang et al., 2019). 13Cdepleted carbonates are common throughout the upper DST (Cui et al.,

Fig. 9. Mo-EF versus U-EF crossplot of Songlin section samples. Diagonal lines represent multiples (0.1, 0.3, 1, and 3) of the Mo:U ratio of present-day seawater. The light-green field represents the “particulate shuttle” trend, characteristic of depositional systems in which intense redox cycling of metal (especially Mn-) oxyhydroxides occurs within the water column (modified from Algeo and Tribovillard, 2009 and Tribovillard et al., 2012). (For interpretation of the references to colour in this figure legend, the reader is referred to the Web version of this article.)

suggest that only approximately 30% of the seawater signal survived during diagenetic alteration (Fig. 10). The oxidation of Ce(III) to Ce(IV) leads to its fractionation relative to other REEs under oxic conditions (Elderfield and Greaves, 1982; Chen et al., 2015). The slightly negative Ce anomalies or no Ce anomalies documented from DST nodule samples indicates no preferential removal of Ce to the host rocks during diagenesis loss (Chen et al., 2015), probably pointing to an anoxic diagenetic microenvironment. The analyzed nodules display positive Eu anomalies (Fig. 7A) which may be attributed to (1) the diagenetic effects of hydrothermal activity 9

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

2016, 2017). Cui et al. (2017) suggested that the magnitude of 13Cdepletion appears to be strongly dependent on the relative abundance of authigenic calcite. This observation compelled Cui et al. (2017) to postulate that the global Ediacaran Shuram Excursion, which appears to correlate with the negative excursion in the upper DST (EN3), was caused by a globally synchronized diagenetic event. δ13C–CH4 and δ13C-DIC (dissolved inorganic carbon) profiles of modern sediment columns display minimum values at the sulfate-methane transition zone (SMTZ), the thin horizon along which upward migrating methane encounters downward diffusing seawater sulfate to be consumed by AOM (Chuang et al., 2019). The depth of SMTZ is generally regulated by the rate of the upward methane flux and consequent rate of downward migrating seawater sulfate (Lash, 2015a; Cui et al., 2017). The postulated sulfate concentration of Ediacaran seawater was less than 10 mM (Loyd et al., 2012; Cui et al., 2017), which was markedly depleted relative to modern oceans (28 mM; Chuang et al., 2019), suggesting a very shallow SMTZ depth, perhaps several cm below the SWI (Cui et al., 2017). It is worth noting that although authigenic calcite can form below the SMTZ, within the methanogenic zone (MEZ), the preferential reduction of 13C-depleted CO2 by methanogens resulted in the 13C-enrichment of residual DIC and the consequent formation of 13C-enriched authigenic carbonates (Lash, 2018). Thus, we suggest that 13C-depleted calcite of DST nodules of the Songlin section was likely formed as a result of the AOM within the SMTZs.

and Lefticariu, 2007; Geilert et al., 2014). It has been hypothesized that the appearance of siliceous organisms in the Early Cambrian time (Xiao et al., 2005; Braun et al., 2007) diminished dissolved silica concentrations in the Phanerozoic global ocean (Maliva and Siever, 1989; Siever, 1992). Perry and Lefticariu (2007) postulated that dissolved silica concentrations of Precambrian seawater and pore water may have been as high as 60 ppm, perhaps ten times higher than that of modern seawater (Racki and Cordey, 2000). Silicification during the Precambrian was largely an abiological process restricted to shallow marine settings, including peritidal and superatidal environments (Maliva et al., 2005; Muscente et al., 2015). High rates of evaporation associated with such conditions elevate salinity and concentrations of dissolved silica in pore waters, thus promoting pervasive silicification at or just below the SWI (Maliva et al., 2005). Muscente et al. (2015) also discussed evidence of silicification of Ediacaran subtidal deposits. Earlier described evidence of replacement of carbonate and phosphate in DST chert nodules by quartz (Figs. 4A–H, J–L, and 5A) suggests that silicification occurred in a pore water environment characterized by elevated concentrations of dissolved silica provided by abiological processes (Knauth, 1994). Sources of the silica include detrital minerals, seawater, siliceous skeletons, and hydrothermal fluids (Maliva and Siever, 1989; Maliva et al., 2005; Perry and Lefticariu, 2007; Braun et al., 2007; Chen et al., 2009; Geilert et al., 2014). The Y/Ho ratio is used to elevate sources of REE in marine sediments (Xin et al., 2015) and of Si in chert (Zhao et al., 2016). In general, siliceous clastic rocks (represented by PAAS and NASC) display a Y/Ho ratio of ~27 (Gromet et al., 1984; Taylor and McLennan, 1985). Fresh water is characterized by Y/Ho ratios that approximate PAAS and NASC values whereas modern seawater generally has higher Y/Ho ratios, ranging from 48 to 59 (Nozaki et al., 1997) and occasionally as high as 80 (Tanaka et al., 2003). Y/Ho ratios of the studied DST chert nodules range from 34.0 to 43.8 (Table 3), indicating that the REE population was sourced mainly from terrigenous siliciclastic minerals (~70%) and subordinately by seawater (~30%; Fig. 10). The host shale of the DST chert nodules contains abundant siliciclastic minerals, including detrital quartz, K-feldspar, and clay (Fig. 5L). The presence of illite-smectite mixed layer clays within chert nodules (Fig. 5K and L) may reflect the effects of the transformation of smectite to illite. Shen et al. (2011) cited Ge/Si ratios as evidence that silica of the DST chert derived from clay diagenesis. Burial diagenesis results in the transformation of smectite to illite over a temperature range of 55–220 °C (Freed and Peacor, 1989). However, the well-preserved nature of soft-body organisms within Ediacaran chert nodules and chert (Fig. 3C and D; Xiao et al., 2010; Muscente et al., 2015; Qu et al., 2017; Shang et al., 2018) indicates that silicification occurred very early in the diagenetic history of these rocks. Amram and Ganor (2005) and Vorhies and Gaines (2009) make a compelling case that the dissolution or conversion of detrital clay minerals (e.g., detrital smectite) is accompanied by the release of silica to pore water via pH- or microbiallymediated processes, which ultimately precipitates as authigenic quartz. Thus, we suggest that silica incorporated into DST chert nodules originated mainly from clay diagenesis and secondarily from seawater.

5.3.2. 18O-depletion of nodules The authigenic carbonate of the DST chert nodules is characterized by strongly negative δ18Ocarb values (< −12‰), markedly more 18Odepleted than those of host carbonate (Fig. 6A). Similar anomalously 18 O-depleted compositions of authigenic carbonate are not uncommon (Raiswell and Fisher, 2000; Lash and Blood, 2004; Lash, 2015a; Cui et al., 2017) and have been attributed to a variety of processes, including an influx of 18O-depleted meteoric water or glacial meltwater, precipitation of 18O-enriched minerals (e.g., carbonate and silicates), precipitation of hydrothermal fluids, and recrystallization during deeper burial and formation of gas hydrates (see review in Raiswell and Fisher, 2000). Micro-textural observations and the geological framework of the DST chert nodules do not support any of these processes. It is worth considering that the 18O-depleted signature of authigenic calcite of the DST nodules does not reflect their early diagenetic history; i.e., AOM. Raiswell and Fisher (2000) and subsequently Lash (2005a) related 18O-depleted isotopic values of authigenic carbonate to the progressive infilling of porosity by 18O-depleted pore fluids during burial. Lash and Blood (2004) reported center-to-edge trends of decreasing 18O from carbonate concretions of the Upper Devonian Rhinestreet Shale of the Appalachian Basin, consistent with the progressive infilling of porosity toward the edges of concretion bodies during burial. According to this hypothesis, early and shallow calcite precipitation results in the formation of a porous but rigid compactionresistant nodule framework (Lash, 2015a). Indeed, the presence of a card-house clay-grain micro-texture in DST nodules of Member IV suggests that the nascent nodules were very porous yet compaction resistant (Dong et al., 2009). Our observations of SB within intercrystalline pores of studied DST nodules provides evidence that the open micro-fabric (Fig. 4H and I) was preserved in nodules during burial to the oil window. Thus, the negative δ18Ocarb values of DST chert nodules may reflect a mixed isotopic signature that initiated in association with precipitation of early and shallow relatively 18O-enriched calcite and continued with the addition of more 18O-depleted calcite cement at progressively greater burial depth and higher temperature as remnant porosity was filled.

5.5. Proposed model of nodule formation Our geochemical data suggest that DST black shale of the Songlin section accumulated under a persistent anoxic, perhaps euxinic, bottom water condition (Figs. 8 and 9). Assuming a low sulfate concentration of the global Ediacaran ocean, SMTZs where methane and sulfate were consumed by AOM to produce 13C-depleted calcite were likely to have existed a very short distance below the SWI (Dong et al., 2009; Cui et al., 2016, 2017). The preservation of soft-body fossils within DST nodules (Fig. 4C and D and 5B) offers the strongest evidence for the initiation of nodule growth at very shallow burial depth. However, the precipitation of 18O-depleted authigenic calcites suggests a protracted history of cementation through much of the burial history of the DST

5.4. Silicification of the Doushantuo nodules Previous studies have described widespread silicification during the Proterozoic Eon (Maliva and Siever, 1989; Maliva et al., 2005; Perry 10

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

6–9 at 25 °C (Langmuir et al., 1997). A high rate of BSR yields H+ thus creating a low-pH pore water microenvironment (pH = 6; Barth, 2003) that accelerates in the dissolution of calcite and precipitation of quartz. The replacement of calcite by quartz appears to have occurred progressively from nodule core to edge (Fig. 4A and B). However, replacement of CFA by quartz in most phosphate nodules appears to have started at nodule edges, which may reflect the more impermeable nature of these nodules (Fig. 4J and K). A small number of phosphate nodules have been completely silicified (Fig. 4L).

including to the point of hydrocarbon generation as suggested by the presence of SB in porosity preserved within the crystals of quartz of calcite (Fig. 4H). The formation processes of DST chert nodules can be briefly described as follows: (1) Proto-nodule formation - AOM. The consumption of methane and sulfate along the SMTZ by AOM quickly yields large amounts of HCO3−, thus increasing the alkalinity of pore waters and favoring the precipitation of 13C-depleted authigenic calcite (Borowski et al., 1999: Raiswell and Fisher, 2004; Lash, 2015a). Episodes of reduced sedimentation (burial) stabilize the SMTZ thus focusing the diagenetic effects of AOM, especially elevated alkalinity, along specific stratigraphic horizons, enabling the growth of nodular or even bedded authigenic calcite (Raiswell, 1988; Borowski et al., 1999; Rodriguez et al., 2000; Snyder et al., 2007; Dickens and Snyder, 2009). The associated BSR of OM yields gas bubbles (e.g., CO2 or ammonia), some of which aggregate into larger voids, within the SMTZ (Dong et al., 2009) may fill with carbonate cement. These similar structures have also been observed in the Pleistocene calcareous concretions, with the diameter of 30–350 μm (generally 100–130 μm), which are uniformly distributed in the whole nodules (Duck, 1995). In our study, optical observations clearly revealed that the core of nodules mainly consisted of one or numerous oblate structures with diameter of greater than 500 μm (Fig. 4A, B and F–I). Inferred gas bubbles of the DST nodules, some > 500 μm in diameter (Fig. 4A) are dominated by quartz, yet the preservation of relict calcite suggests that they were once filled with calcite (Fig. 4A–C and F–H). The spherical shape of small phosphate nodules (400–500 μm in diameter) hosted in DST shale (Fig. 4J and K) suggests formation prior to compaction, perhaps earlier than calcite precipitation. The close association of CFA and OM (Fig. 4J and K) in phosphate nodules indicates that CFA largely derived from the release of P from OM decaying (Föllmi, 1996). (2) Precipitation of metal sulfides within nodule rims. BSR of OM and AOM yields hydrogen sulfide (Callow and Brasier, 2009; Lash, 2015b) that may combine with available metal ions (e.g., Zn2+ and Fe2+) to form sulfide minerals, including sphalerite and pyrite (Fig. 5D). The strong positive correlation of Zn concentration and TOC in host shale (R2 = 0.64, n = 7) suggests that aqueous Zn may have derived from decayed OM. Moreover, the textural association of sphalerite and CFA (Fig. 5D and J) may reflect the addition of Zn released from Zn-bonded phosphate in OM and/or Fe-(oxyhydr) oxide particles that populated in the Ediacaran phosphorus-rich ocean (Fan et al., 2018). The likelihood of anoxic bottom- and pore-water conditions associated with deposition of the studied DST black shale exposed in the Songlin section suggests that pyrite precipitation and pyritization occurred in the presence of abundant hydrogen sulfide probably produced by the BSR and AOM. Thus, the precipitation of metal sulfides would have been contemporaneous with the precipitation of authigenic calcite in association with AOM-induced production of HCO3−. (3) Silicification. The preservation of calcite relicts and ghosts (Fig. 4A–H) within DST chert nodules suggests that silicification occurred later in the formation history of these deposits. In general, two basic conditions are required for silicification: (1) an abnormally high concentration of dissolved silica and (2) weakly acidic to weakly alkaline diagenetic microenvironment (pH = 6–9; Langmuir et al., 1997; Cui et al., 2017). Ediacaran sea water appears to have been enriched in aqueous Si relative to modern sea water (Maliva and Siever, 1989; Siever, 1992; Perry and Lefticariu, 2007). Moreover, aqueous silica could have been supplemented by clay diagenesis as evidenced by the lower Y/Ho ratios (Fig. 9) and the presence of smectite/illite mixed layer clay (Fig. 5K). Precipitation of quartz appears to require concomitant calcite dissolution (Muscente et al., 2015; Cui et al., 2017). Indeed, calcite and silica have opposing solubility-pH relationships over a pH range of approximately 9–10 (Knauth, 1994; Langmuir et al., 1997). Moreover, the acidification of pore water may induce calcite dissolution thus favoring SiO2 precipitation under pH conditions of

6. Conclusions Black shale of the Ediacaran DST Formation (Member IV) containing abundant chert nodules accumulated under persistent anoxic bottom water conditions. Methane ascending from the sediment column and seawater sulfate migrating downward from the SWI were consumed by AOM and BSR along the SMTZ. Periodic reductions in sedimentation (burial) rate stabilized the SMTZ and associated diagenesis resulting in the formation of the nodules. Chert nodules display a concentric internal structure dominated in their cores by quartz. Individual laminae thin outward from cores and display a more varied mineralogy, including calcite, CFA, OM, pyrite, and sphalerite. 13C-depleted authigenic calcite produced by AOM precipitated in gas bubbles generated by BSR. The degradation of OM also released phosphate and zinc ions to pore water, favoring the precipitation of CFA and sphalerite, especially in nodule rims. Silicification of the nodules may have triggered by an enhanced rate of BSR later in the diagenetic history of these deposits. The bulk of the silica may have derived from clay diagensis of shale host rocks and secondarily from seawater trapped within pores. The residual porosity of nodules was gradually filled by the 18O-depleted pore fluids during burial. The DST chert nodules, then, record a protracted diagenetic history that includes the release of silica by clay-diagenesis and consequent of replacement of early formed calcite by quartz. This mechanism could account for silicification during the Ediacaran or Neoproterozoic periods. Author contributions section Ping Gao: Investigation, Methodology, Formal analysis, Writing – Original Draft Preparation. Zhiliang He: Funding acquisition, Supervision. Gary G. Lash: Writing, Reviewing and Editing. Shuangjian Li: Sampling assistance, Funding acquisition. Rongqiang Zhang: Sampling assistance. Declaration of competing interest The authors declared that there is no conflict of interest. Acknowledgments This work was supported by the National Natural Science Foundation of China (41802170), the National Science and Technology Major Project of the Ministry of Science and Technology of China (2017ZX05008-002-004, 2017ZX05005-001), and the China Postdoctoral Science Foundation (2018M631688). Appendix A. Supplementary data Supplementary data to this article can be found online at https:// doi.org/10.1016/j.marpetgeo.2020.104227. References Algeo, T.J., Lyons, T.W., 2006. Mo–total organic carbon covariation in modern anoxic marine environments: implications for analysis of paleoredox and paleohydrographic conditions. Paleoceanography 21 PA1016.

11

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Cosmochim. Acta 48, 2469–2482. Hinrichs, K., Hayes, J.M., Sylva, S.P., Brewer, P.G., DeLong, F., EF, 1999. Methane-consuming archaebacteria in marine sediments. Nature 398, 802–805. Hu, W.-X., Kang, X., Cao, J., Wang, X.-L., Fu, B., Wu, H.-G., 2018. Thermochemical oxidation of methane induced by high-valence metal oxides in a sedimentary basin. Nat. Commun. 9, 5131. Jiang, G., Shi, X., Zhang, S., Wang, Y., Xiao, S., 2011. Stratigraphy and paleogeography of the ediacaran Doushantuo Formation (ca. 635–551Ma) in South China. Gondwana Res. 19, 831–849. Jiang, G., Kaufman, A.J., Christie-Blick, N., Zhang, S., Wu, H., 2007. Carbon isotope variability across the Ediacaran Yangtze platform in South China: implications for a large surface-to-deep ocean δ13C gradient. Earth Planet. Sci. Lett. 261, 303–320. Khalil, K., Rabouille, C., Gallinari, M., Soetaert, K., DeMaster, D.J., Ragueneau, O., 2007. Constraining biogenic silica dissolution in marine sediments: a comparison between diagenetic models and experimental dissolution rates. Mar. Geol. 106, 223–238. Knauth, L.P., 1994. Petrogenesis of chert. Rev. Mineral. Geochem. 29, 233–258. Langmuir, D., Hall, P., Drever, J., 1997. Environmental Geochemistry. Prentice Hall, New Jersey. Lash, G.G., 2015a. Authigenic barite nodules and carbonate concretions in the Upper Devonian shale succession of western New York – a record of variable methane flux during burial. Mar. Pet. Geol. 59, 305–319. Lash, G.G., 2015b. Pyritization induced by anaerobic oxidation of methane (AOM) – an example from the upper devonian shale succession, western New York, USA. Mar. Pet. Geol. 68, 520–535. Lash, G.G., 2018. Significance of stable carbon isotope trends in carbonate concretions formed in association with anaerobic oxidation of methane (AOM), Middle and Upper Devonian shale succession, western New York State, USA. Mar. Pet. Geol. 91, 470–479. Lash, G.G., Blood, D., 2004. Geochemical and textural evidence for early (shallow) diagenetic growth of stratigraphically confined carbonate concretions, Upper Devonian Rhinestreet black shale, western New York. Chem. Geol. 206, 407–424. Lawrence, M.G., Greig, A., Collerson, K.D., Kamber, B.S., 2006. Rare earth element and yttrium variability in South east queensland waterways. Aquat. Geochem. 12, 39–72. Loi, A., Dabard, M.-P., 2002. Controls of sea level fluctuations on the formation of Ordovician siliceous nodules in terrigenous offshore environments. Sediment. Geol. 153, 65–84. Loi, A., Ghienne, J.F., Dabard, M.P., Paris, F., Botquelen, A., Christ, N., Elaouad-Debbaj, Z., Gorini, A., Vidal, M., Videt, B., Destombes, J., 2010. The Late Ordovician glacioeustatic record froma high-latitude storm-dominated shelf succession: the Bou Ingarf section (Anti-Atlas, SouthernMorocco). Palaeogeogr. Palaeoclimatol. Palaeoecol. 296, 332–358. Loyd, S.J., Marenco, P.J., Hagadorn, J.W., Lyons, T.W., Kaufman, A.J., Sour-Tovar, F., Corsetti, F.A., 2012. Sustained low marine sulfate concentrations from the Neoproterozoic to the Cambrian: insights from carbonates of northwestern Mexico and eastern California. Earth Planet. Sci. Lett. 339–340, 79–94. Maliva, R.G., Knoll, A.H., Simonson, B.M., 2005. Secular change in the Precambrian silica cycle: insights from chert petrology. Geol. Soc. Am. Bull. 117, 835–845. Maliva, R.G., Siever, R., 1989. Nodular chert formation in carbonate rocks. J. Geol. 97, 421–433. Marin-Carbonne, J., Robert, F., Chaussidon, M., 2014. The silicon and oxygen isotope compositions of Precambrian cherts: a record of oceanic paleo-temperatures? Precambrian Res. 247, 223–234. Martinez-Ruiz, F., Ortega-Huertas, M., Palomo, I., 1999. Positive Eu anomaly development during diagenesis of the K/T boundary ejecta layer in the Agost section (SE Spain): implications for trace-element remobilization. Terra. Nova 11, 290–296. McFadden, K.A., Xiao, S., Zhou, C., Kowalewski, M., 2009. Quantitative evaluation of the biostratigraphic distribution of acanthomorphic acritarchs in the ediacaran Doushantuo Formation in the Yangtze Gorges area, South China. Precambrian Res. 173, 170–190. McLennan, S.M., 1989. Rare earth elements in sedimentary rocks; influence of provenance and sedimentary processes. Rev. Mineral. Geochem. 21, 169–200. McLennan, S.M., 2001. Relationships between the trace element composition of sedimentary rocks and upper continental crust. Geochem. Geophys. Geosyst. 2 (2000GC000109), 24. Muscente, A.D., Hawkins, A.D., Xiao, S., 2015. Fossil preservation through phosphatization and silicification in the Ediacaran Doushantuo Formation (South China): a comparative synthesis. Palaeogeogr. Palaeoclimatol. Palaeoecol. 434, 46–62. Niewöhner, C., Henson, C., Kasten, S., Zabel, M., Schultz, H.D., 1998. Deep sulfate reduction completely mediated by anaerobic methane oxidation in sediments of the upwelling area off Nambia. Geochem. Cosmochim. Acta 62, 455–464. Nozaki, Y., Zhang, J., Amakawa, H., 1997. The fractionation between Y and Ho in the marine environment. Earth Planet. Sci. Lett. 148, 329–340. Perry, E.C., Lefticariu, L., 2007. Formation and geochemistry of Precambrian cherts. In: Holland, H.D., Turekian, K.K. (Eds.), Treatise on Geochemistry. Pergamon, Oxford, pp. 1–21. Powell, W.G., Johnston, P.A., Collom, C.J., 2003. Geochemical evidence for oxygenated bottom waters during deposition of fossiliferous strata of the Burgess Shale Formation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 201, 249–268. Qu, Y., Wang, J., Xiao, S., Whitehouse, M., Engdahl, A., Wang, G., McLoughlin, N., 2017. Carbonaceous biosignatures of diverse chemotrophic microbial communities from chert nodules of the Ediacaran Doushantuo Formation. Precambrian Res. 290, 184–196. Racki, G., Cordey, F., 2000. Radiolarian palaeoecology and radiolarites: is the present the key to the past? Earth Sci. Rev. 52, 83–120. Raiswell, R., 1988. A chemical model for the origin of minor limestone-shale cycles by anaerobic methane oxidation. Geology 16, 641–644.

Algeo, T.J., Rowe, H., 2012. Paleoceanographic applications of trace-metal concentration data. Chem. Geol. 324–325, 6–18. Algeo, T.J., Tribovillard, N., 2009. Environmental analysis of paleoceanographic systems based on molybdenum–uranium covariation. Chem. Geol. 268, 211–225. Amram, K., Ganor, J., 2005. The combined effect of pH and temperature on smectite dissolution rate under acidic conditions. Geochem. Cosmochim. Acta 69, 2535–2546. Barth, H.-J., 2003. The influence of cyanobacteria on oil polluted intertidal soils at the Saudi Arabian Gulf shores. Mar. Pollut. Bull. 46, 1245–1252. Borowski, W.S., Paull, C.K., Ussler III, W., 1999. Global and local variations of interstitial sulfate gradients in deep-water, continental margin sediments: sensitivity to underlying methane and gas hydrates. Mar. Geol. 159, 131–154. Braun, A., Chen, J.-Y., Waloszek, D., Maas, A., 2007. Siliceous microfossils and biosiliceous sedimentation in the lowermost Cambrian of China. Geol. Soc., Lond., Spec. Publ. 286, 423–424. Callow, R.H.T., Brasier, M.D., 2009. Remarkable preservation of microbial mats in Neoproterozoic siliciclastic settings: implications for Ediacaran taphonomic models. Earth Sci. Rev. 96, 207–219. Chen, D., Wang, J., Qing, H., Yan, D., Li, R., 2009. Hydrothermal venting activities in the Early Cambrian, South China: petrological, geochronological and stable isotopic constraints. Chem. Geol. 258, 168–181. Chen, J., Algeo, T.J., Zhao, L., Chen, Z.-Q., Cao, L., Zhang, L., Li, Y., 2015. Diagenetic uptake of rare earth elements by bioapatite, with an example from Lower Triassic conodonts of South China. Earth Sci. Rev. 149, 181–202. Chuang, P.-C., Yang, T.F., Wallmann, K., Matsumoto, R., Hu, C.-Y., Chen, H.-W., Lin, S., Sun, C.-H., Li, H.-C., Wang, Y., Dale, A.W., 2019. Carbon isotope exchange during anaerobic oxidation of methane (AOM) in sediments of the northeastern South China Sea. Geochem. Cosmochim. Acta 246, 138–155. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., Jin, Y., 2005. U-Pb ages from the neoproterozoic Doushantuo Formation, China. Science 308, 95–98. Cui, H., Kaufman, A.J., Xiao, S., Zhou, C., Liu, X.M., 2017. Was the Ediacaran Shuram Excursion a globally synchronized early diagenetic event? Insights from methanederived authigenic carbonates in the uppermost Doushantuo Formation, South China. Chem. Geol. 450, 59–80. Cui, H., Xiao, S., Zhou, C., Peng, Y., Kaufman, A.J., Plummer, R.E., 2016. Phosphogenesis associated with the Shuram excursion: petrographic and geochemical observations from the ediacaran Doushantuo Formation of South China. Sediment. Geol. 341, 134–146. Dabard, M.-P., Loi, A., 2012. Environmental control on concretion-forming processes: examples from paleozoic terrigenous sediments of the north gondwana margin, armorican massif (middle ordovician and middle devonian) and SW sardinia (late ordovician). Sediment. Geol. 267, 93–103. Dickens, J., Snyder, G.T., 2009. Interpreting Upward Methane Flux from Pore Water Profiles. Fire in the Ice. National Energy Technology Laboratory Methane Hydrate Newsletter, pp. 7–10 Winter. Dong, L., Shen, B., Lee, C.-T.A., Shu, X., Peng, Y., Sun, Y., Tang, Z., Rong, H., Lang, X., Ma, H., Yang, F., Guo, W., 2015. Germanium/silicon of the Ediacaran-Cambrian Laobao cherts: Implications for the bedded chert formation and paleoenvironment interpretations. Geochem. Geophys. Geosyst. 16, 751–763. Dong, J., Zhang, S., Jiang, G., Zhao, Q., Li, H., Shi, X., Liu, J., 2009. Study on formation environment of carbonate concretions from fourth member of Doushantuo Formation in Yichang area, South China and its significance to source rock evaluation. Sci. China Ser. D Earth Sci. 39, 317–326 (in Chinese). Duck, R.W., 1995. Subaqueous shrinkage cracks and early sediment fabrics preserved in Pleistocene calcareous concretions. J. Geol. Soc. Lond. 152, 151–156. Dulski, P., 1994. Interferences of oxide, hydroxide and chloride analyte species in the determination of rare earth elements in geological samples by inductively coupled plasma-mass spectrometry. Fresenius’ J. Anal. Chem. 350, 194–203. Ehlert, C., Doering, K., Wallmann, K., Scholz, F., Sommer, S., Grasse, P., Geilert, S., Frank, M., 2016. Stable silicon isotope signatures of marine pore waters – biogenic opal dissolution versus authigenic clay mineral formation. Geochem. Cosmochim. Acta 191, 102–117. Elderfield, H., Greaves, M.J., 1982. The rare earth elements in seawater. Nature 296, 214–219. Fan, H., Wen, H., Xiao, C., Zhou, T., Cloquet, C., Zhu, X., 2018. Zinc geochemical cycling in a phosphorus-rich ocean during the Early Ediacaran. J. Geophys. Res.: Oceans 123 (8), 5248–5260. Freed, R.L., Peacor, D.R., 1989. Geopressured shale and sealing effect of smectite to illite transition. AAPG (Am. Assoc. Pet. Geol.) Bull. 73, 1223–1232. Froelich, P.N., Arthur, M.A., Burnett, W.C., Deakin, M., Hensley, V., Jahnke, R., Kaul, L., Kim, K.-H., Roe, K., Soutar, A., Vathakanon, C., 1988. Early diagenesis of organic matter in Peru continental margin sediments: phosphorite precipitation. Mar. Geol. 80, 309–343. Föllmi, K.B., 1996. The phosphorus cycle, phosphogenesis and marine phosphate-rich deposits. Earth Sci. Rev. 40, 55–124. Gao, P., Liu, G., Jia, C., Young, A., Wang, Z., Wang, T., Zhang, P., Wang, D., 2016. Redox variations and organic matter accumulation on the Yangtze carbonate platform during Late Ediacaran–Early Cambrian: constraints from petrology and geochemistry. Palaeogeogr. Palaeoclimatol. Palaeoecol. 450, 91–110. Gao, P., He, Z., Li, S., Lash, G.G., Li, B., Huang, B., Yan, D., 2018. Volcanic and hydrothermal activities recorded in phosphate nodules from the Lower Cambrian Niutitang Formation black shales in South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 505, 381–397. Geilert, S., Vroon, P.Z., van Bergen, M.J., 2014. Silicon isotopes and trace elements in chert record early Archean basin evolution. Chem. Geol. 386, 133–142. Gromet, L.P., Dymek, R.F., Haskin, L.A., Korotev, R.L., 1984. The “North American shale composite”: its compilation, major and trace element characteristics. Geochem.

12

Marine and Petroleum Geology 114 (2020) 104227

P. Gao, et al.

Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: its Composition and Evolution. Blackwell Scientific Publications, Oxford, pp. 311. Tribovillard, N., Algeo, T.J., Baudin, F., Riboulleau, A., 2012. Analysis of marine environmental conditions based on molybdenum–uranium covariation—Applications to Mesozoic paleoceanography. Chem. Geol. 325, 46–58. Tribovillard, N., Algeo, T.J., Lyons, T., Riboulleau, A., 2006. Trace metals as paleoredox and paleoproductivity proxies: an update. Chem. Geol. 232, 12–32. Vorhies, J.S., Gaines, R.R., 2009. Microbial dissolution of clay minerals as a source of iron and silica in marine sediments. Nat. Geosci. 2, 221–225. Wen, H., Fan, H., Tian, S., Wang, Q., Hu, R., 2016. The formation conditions of the early Ediacaran cherts, South China. Chem. Geol. 430, 45–69. Wignall, P.B., Newton, R., 1998. Pyrite framboid diameter as a measure of oxygen deficiency in ancient mudrocks. Am. J. Sci. 298, 537–552. Wilkin, R.T., Barnes, H.L., Brantley, S.L., 1996. The size distribution of framboidal pyrite in modern sediments: an indicator of redox conditions. Geochem. Cosmochim. Acta 60, 3897–3912. Xiao, S., Hu, J., Yuan, X., Parsley, R.L., Cao, R., 2005. Articulated sponges from the Lower Cambrian Hetang Formation in southern Anhui, South China: their age and implications for the early evolution of sponges. Palaeogeogr. Palaeoclimatol. Palaeoecol. 220, 89–117. Xiao, S., Schiffbauer, J.D., McFadden, K.A., Hunter, J., 2010. Petrographic and SIMS pyrite sulfur isotope analyses of Ediacaran chert nodules: implications for microbial processes in pyrite rim formation, silicification, and exceptional fossil preservation. Earth Planet. Sci. Lett. 297, 481–495. Xiao, S., McFadden, K.A., Peek, S., Kaufman, A.J., Zhou, C., Jiang, G., Hu, J., 2012. Integrated chemostratigraphy of the Doushantuo Formation at the northern Xiaofenghe section (Yangtze Gorges, South China) and its implication for Ediacaran stratigraphic correlation and ocean redox models. Precambrian Res. 192–195, 125–141. Xin, H., Jiang, S.-Y., Yang, J.-H., Wu, H.-P., Pi, D.-H., 2015. Rare earth element and Sr–Nd isotope geochemistry of phosphatic rocks in neoproterozoic ediacaran Doushantuo Formation in zhangcunping section from western hubei Province, South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 440, 712–724. Yang, A., Zhu, M., Zhang, J., Zhao, F., Lü, M., 2015. Sequence stratigraphic subdivision and correlation of the ediacaran (sinian) Doushantuo Formation of Yangtze plate, South China. J. Palaeogeogr. 17, 1–20 (in Chinese with English abstract). Zhang, L., Algeo, T.J., Cao, L., Zhao, L., Chen, Z.-Q., Li, Z., 2016. Diagenetic uptake of rare earth elements by conodont apatite. Palaeogeogr. Palaeoclimatol. Palaeoecol. 458, 176–197. Zhao, Y.-Y., Jiang, S.-Y., Li, D., Yang, J.-H., 2016. A petrographic and geochemical study of carbonate and silica phases from the Ediacaran Doushantuo Formation in the Three Gorges area of South China: implications for diagenetic conditions. Palaeogeogr. Palaeoclimatol. Palaeoecol. 463, 150–167.

Raiswell, R., Fisher, Q.J., 2000. Mudrock‐hosted carbonate concretions: a review of growth mechanisms and their influence on chemical and isotopic composition. J. Geol. Soc. 157, 239–251 London. Raiswell, R., Fisher, Q., 2004. Rates of carbonate cementation associated with sulphate reduction in DSDP/ODP sediments: implications for the formation of concretions. Chem. Geol. 211, 71–85. Reeburgh, W.S., 1976. Methane consumption in Cariaco Trench waters and sediments. Earth Planet. Sci. Lett. 28, 337–344. Rickard, D., 2019. Sedimentary pyrite framboid size-frequency distributions: a metaanalysis. Palaeogeogr. Palaeoclimatol. Palaeoecol. 522, 62–75. Rodriguez, N.M., Paull, C.K., Borowski, W.S., 2000. Zonation of authigenic carbonates within gas hydrate-bearing sedimentary sections on the Blake Ridge: offshore southeastern North America. In: Paull, C.K., Matsumoto, R., Wallace, P.J., Dillon, W.P. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results. 164. pp. 301–312. Scott, C., Lyons, T.W., 2012. Contrasting molybdenum cycling and isotopic properties in euxinic versus non-euxinic sediments and sedimentary rocks: refining the paleoproxies. Chem. Geol. 324–325, 19–27. Shang, X., Moczydłowska, M., Liu, P., Liu, L., 2018. Organic composition and diagenetic mineralization of microfossils in the Ediacaran Doushantuo chert nodule by Raman and petrographic analyses. Precambrian Res. 314, 145–159. Shen, B., Lee, C.-T., Xiao, S., 2011. Germanium/silicon ratios in diagenetic chert nodules from the Ediacaran Doushantuo Formation, South China. Chem. Geol. 280, 323–335. Siever, R., 1992. The silica cycle in the Precambrian. Geochem. Cosmochim. Acta 56, 3265–3272. Snyder, G.T., Dickens, G.R., Tomaru, H., Takeuchi, R., Komatsubara, J., Ishida, Y., Yu, H., 2007. Pore water profiles and authigenic mineralization in shallow marine sediments above the methane-charged system on Umitaka Spur, Japan Sea. Deep-Sea Research, Part II. Top. Stud. Oceanogr. 54, 1216–1239. Steiner, M., Wallis, E., Erdtmann, B.-D., Zhao, Y., Yang, R., 2001. Submarine-hydrothermal exhalative ore layers in black shales from South China and associated fossils insights into a Lower Cambrian facies and bio-evolution. Palaeogeogr. Palaeoclimatol. Palaeoecol. 169, 165–191. Stefurak, E.J.T., Lowe, D.R., Zentner, D., Fischer, W.W., 2014. Primary silica granules-A new mode of Paleoarchean sedimentation. Geology 42, 283–286. Tahata, M., Ueno, Y., Ishikawa, T., Sawaki, Y., Murakami, K., Han, J., Shu, D., Li, Y., Guo, J., Yoshida, N., Komiya, T., 2013. Carbon and oxygen isotope chemostratigraphies of the Yangtze platform, South China: decoding temperature and environmental changes through the Ediacaran. Gondwana Res. 23, 333–353. Tanaka, K., Miura, N., Asahara, Y., Kawabe, I., 2003. Rare earth element and strontium isotopic study of seamount-type limestones in Mesozoic accretionary complex of Southern Chichibu Terrane, central Japan: implication for incorporation process of seawater REE into limestones. Geochem. J. 37, 163–180.

13