A petrographic and geochemical study of carbonate

Palaeogeography, Palaeoclimatology, Palaeoecology 463 (2016) 150–167

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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 Yan-Yan Zhao a,b,c,⁎, Shao-Yong Jiang d,e, Da Li e, Jing-Hong Yang e a

College of Marine Geosciences, Ocean University of China, Qingdao 266100, China Key Lab of Submarine Geosciences and Prospecting Techniques, Ministry of Education, Qingdao 266100, China c CAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Science, University of Science and Technology of China, Hefei 230026, China d State Key Laboratory of Geological Processes and Mineral Resources, Faculty of Earth Resources, China University of Geosciences, Wuhan 430074, China e State Key Laboratory for Mineral Deposits Research, Department of Earth Sciences, Nanjing University, Nanjing 210093, China b

a r t i c l e

i n f o

Article history: Received 6 May 2016 Received in revised form 30 September 2016 Accepted 6 October 2016 Available online 8 October 2016 Keywords: Oxygen isotope Trace element REEs Diagenesis Chert Nodules

a b s t r a c t The carbonate of the Ediacaran Doushantuo Formation contains various silica phases in the Three Gorges area of South China. To better understand the diagenetic conditions under which these silica phases formed, we carried out a petrographic and geochemical study of the silica phases and the coadjacent carbonate phases at the Xiaofenghe section (XFH), in the Three Gorges area of South China. Five types of silica phases can be distinguished based on petrographic characteristics: S1 and S2 are composed of microcrystalline quartz, in chert nodules and disseminated in the host dolostone, respectively. S3 is composed of the megaquartz, while S4 and S5 are microcrystalline quartz and fibrous silica, respectively, both replacing calcite cement. Six distinct types of carbonate can be recognized, including the host matrix carbonate (C1), sparry calcite nodules (C2), calcite around chert nodules (C3), dispersed calcite cement within chert nodules (C4), fine micropeloidal calcite cements (C5) and calcite veins (C6). Sedimentary structures suggest that formation of the S1 through S5 silica phases and C2 through C5 carbonate phases occurred during early burial diagenesis, and therefore their formation recorded early diagenetic conditions. Assuming that Ediacaran surface seawater temperature ranged from 28 to 39 °C, surface seawater δ18O values can be calculated based on the oxygen isotopic composition of the host dolostone matrix (C1), which range from −6 to 0‰, −4 to 0‰, and −2 to 0‰, for the D2, D3 and D4 subunits of the Doushantuo Formation, respectively. Based on the δ18O values of chert nodules and beds (mainly phase S1), diagenetic temperatures can be constrained to between 27 and 60 °C. Under such low diagenetic temperatures, the primary geochemical signatures of the host dolostone, such as REE + Y patterns, δ13CCarb and δ34SPy values, are likely to be preserved. Those silica phases without evidence of contamination of continental silicate detritus show characteristic REE + Y patterns, with minor light REE depletions, positive La, Eu and Gd anomalies, negative Ce anomalies and slightly super-chondritic Y/Ho ratios, implying that the source of silica may be derived from seawater. Negative Ce anomalies in the silica phases in D2 through D4 subunits indicate that the bottom seawater or early diagenetic porewater was oxic at the time of silica precipitation. © 2016 Elsevier B.V. All rights reserved.

1. Introduction The geochemical and fossil records of the Ediacaran Period have been extensively studied to explore the impact of changing climate and oceanic redox conditions on the development of early multicellular life. The Doushantuo Formation, in the Three Gorges area of South ⁎ Corresponding author at: CAS Key Laboratory of Crust-Mantle Materials and Environments, School of Earth and Space Science, University of Science and Technology of China, Hefei 230026, China. E-mail address: [email protected] (Y.-Y. Zhao).

http://dx.doi.org/10.1016/j.palaeo.2016.10.004 0031-0182/© 2016 Elsevier B.V. All rights reserved.

China, shows many of the geochemical trends and features thought to be typical of the global Ediacaran marine record (McFadden et al., 2008; Li et al., 2010). However, there are some indications that post-depositional processes have modified the primary geochemical signals (Knauth and Kennedy, 2009; Derry, 2010; Grotzinger et al., 2011). For example, the validity of carbonate isotope ratios has come under increased scrutiny, as a result of the observation that carbon and oxygen isotopes covary in several Ediacaran sections (Knauth and Kennedy, 2009; Derry, 2010; Grotzinger et al., 2011). Diagenetic and hydrothermal processes can influence the geochemistry of sedimentary carbonates, and are the most likely mechanism

Y.-Y. Zhao et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 463 (2016) 150–167

responsible for the observed coupling of carbon and oxygen isotope signals (Jiang et al., 2006; Ling et al., 2007; Ader et al., 2009; Sawaki et al., 2010). Bristow et al. (2011) suggested that the highly 13C-depleted carbonates (with δ13C values as low as − 48‰ VPDB; Jiang et al., 2003; Wang et al., 2008; Zhou et al., 2016) in the cap dolostones at the base of the Doushantuo Formation were precipitated from hydrothermal fluids (N 300 °C) that preferentially flowed through the cap dolostones after deposition. Chen et al. (2015) demonstrated that diagenesis can alter the REE distribution in marine sediments. Therefore, additional methods are needed to assess the extent of diagenetic overprints before using the geochemical data to reconstruct ancient seawater chemistry and paleoceanographic environments. The carbonates of the Ediacaran Doushantuo Formation, in the Three Gorges area of South China, are diagenetically silicified (McFadden et al., 2009; Xiao et al., 2010, 2012; Zhu et al., 2013; Liu et al., 2014a, 2014b). Silica occurs in numerous forms, including disseminated macro- and micro-crystalline quartz, chert nodules, and sheet-like chert bands (McFadden et al., 2009; Xiao et al., 2012; Wen et al., 2016), which precipitated during post-depositional processes (Xiao et al., 2010; Shen et al., 2011). The chert nodules in the Doushantuo Formation overlying the cap carbonate at Xiaofenghe (XFH) section, in particular, preserved assemblages of fossil prokaryotic and eukaryotic organisms, including cyanobacteria, multicellular algae, acanthomorphic (spiny) acritarchs, and microfossils interpreted as animal resting eggs and embryos (Zhang et al., 1998; Yin et al., 2007; Xiao et al., 2012; Liu et al., 2013). Studying the silica phases preserved in the Doushantuo Formation can provide information not only about the post-depositional processes, but also about sources of silica and the conditions of fossil preservation. A number of previous studies have demonstrated that the silicon, hydrogen, and oxygen isotopic compositions of cherts have the potential to serve as records of environmental conditions and can record formation and burial temperatures, oxygen isotopic compositions of the ancient oceans, water depth, and paleoclimate (Knauth and Epstein, 1976; Jiang et al., 1993, 1994; Knauth and Lowe, 2003; Kasting et al., 2006; Marin et al., 2010; Van den Boorn et al., 2010; Marin-Carbonne et al., 2012, 2014; Ramseyer et al., 2013; Wen et al., 2016). For example, the δ18O and δ2H values of the 3.42-Ga Buck Reef Chert in South Africa record a minimum diagenetic temperature below ~ 40 °C (Hren et al., 2009). Furthermore, the concentrations of rare earth elements (REEs) in chert vary systematically, depending on the relative influence of hydrothermal, terrestrial, and marine inputs in the basin (Owen et al., 1999b; Friend et al., 2008; Shen et al., 2011). Murray (1994) demonstrated that REE patterns in chert can distinguish between continental margin, ridge proximal, and pelagic deposits. Specifically, cherts precipitated directly from seawater are typically characterized by positive La anomalies, LREE depletions, negative Ce anomalies (if seawater was locally oxygenated), as well as slightly higher Y/Ho ratios relative to the chondritic value of ~ 28 (Lawrence et al., 2006; Bolhar and Van Kranendonk, 2007; Van den Boorn et al., 2010). Sedimentological, chemostratigraphic, and paleontological data suggest that the Doushantuo Formation in the Three Gorges area of South China was deposited in a subtidal or intertidal shelf environment (Zhu et al., 2013), and that the basin was at least intermittently connected with the global ocean during the Ediacaran (Vernhet, 2007; Zhu et al., 2007; Jiang et al., 2008; Shen et al., 2008; Ader et al., 2009; Li et al., 2010; Jiang et al., 2011; Lu et al., 2013; Xiao et al., 2012, 2014; Zhu et al., 2013). Despite numerous studies of the silica phases in the Doushantuo Formation overlying the cap carbonate, the source of silica is still controversial, and the diagenetic temperature and redox conditions are not fully understood. Xiao et al. (2010) proposed that the chert nodules contained in the Doushantuo Formation precipitated rapidly from silica-rich Precambrian seawater during early diagenesis, based on the sulfur isotopic composition of in situ pyrites. Shen et al. (2011) suggested that the chert nodules formed in a restricted environment with a strong terrestrial influence. Wen et al. (2016) suggested

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that Ge and Si in the cherts were derived mainly from seawater, with minor detrital and negligible hydrothermal contamination. To better constrain the source and post-depositional history of these silica phases, we performed a detailed geochemical study of the silica phases and adjacent carbonate phases in the Doushantuo Formation overlying the cap carbonate at XFH. 2. Geological setting The Ediacaran succession in the Three Gorges area of South China consists of the Doushantuo Formation and the Dengying Formation (Fig. 1). The Doushantuo Formation has been radiometrically constrained to between 635.2 ± 0.6 and 551.0 ± 0.7 Ma (Condon et al., 2005), and has been the focus of extensive paleontological and geochemical study over the past decade (Yin et al., 2007; Zhou et al., 2007; Shen et al., 2008; McFadden et al., 2008, 2009; Zhu et al., 2008; Jiang et al., 2011; Xiao et al., 2012; Derkowski et al., 2013; Kikumoto et al., 2014; Shields-Zhou and Zhu, 2013; Tahata et al., 2013; Liu et al., 2014a, 2014b; Sawaki et al., 2014; She et al., 2014; Wang et al., 2014a, 2014b; Xin et al., 2015, 2016). The Xiaofenghe section is situated on the eastern flank of the Huangling Anticline, approximately 6 km west of Xiaofeng Village, and is approximately 35 km updip from the Jiulongwan section (Fig. 1). This section contains the thickest exposure of the Doushantuo Formation (approximately 220 m) in the Three Gorges area (Yin et al., 2007;

Fig. 1. Simplified geological map showing the occurrence of upper Neoproterozoic sedimentary rocks of the Huangling Anticline, on the Three Gorges area of South China (insert). Map modified from Xiao et al. (2012). Asterisks denote the sample localities.

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Xiao et al., 2012; Zhu et al., 2013). Diapause embryos of large acanthomorphic acritarchs and spiny organic-walled microfossils, are preserved in the chert nodules at this locality (Yin et al., 2007). In this study, we examined the Doushantuo Formation at XFH in two sections, exposed along either bank of a small stream in Xiaofenghe Village. The lower Doushantuo Formation at XFH was measured along the northern bank (NXFH), however, the upper part of the formation is not exposed above 140 m (McFadden et al., 2009; Xiao et al., 2012; Zhu et al., 2013). A second stratigraphic section, which extends up to the Doushantuo-Dengying Formation contact, was measured along the southern bank (SXFH; Yin et al., 2007; Zhu et al., 2007, 2013; Xiao et al., 2012). Although the contact between the Nantuo Formation and the Doushantuo Formation is easily recognized at NXFH, the three sedimentary disconformities cannot be identified due to the poor outcrop exposure (Xiao et al., 2012; Zhu et al., 2013). The cap carbonate subunit (D1) at NXFH is approximately 4.8 m thick, directly overlying the diamictites of the Nantuo Formation (Fig. 2). The carbonate is similar to other Ediacaran cap carbonates

described elsewhere in South China (Zhu et al., 2007, 2013; Xiao et al., 2012), and is composed of sparry dolostone at the base, overlain by microlaminated and organic-rich dolomicrite with irregular quartz veins, brecciated dolomite, and abundant tepee-like structures. The upper part of the cap carbonate subunit is laminated, and partially silicified (Zhu et al., 2007; Xiao et al., 2012; Hohl et al., 2015). Member II of the Doushantuo Formation (D2) is overlying the cap carbonate. The lower D2 subunit is composed of black shale, interbedded with laminated silty or muddy dolostone with chert nodules (Xiao et al., 2012). The upper D2 subunit contains argillaceous, thinto medium-bedded dolostones containing abundant pea-sized chert nodules and phosphatic grains (Figs. 2, 3A and B), with the abundance of phosphatic grains increasing upsection. The base of the Member III (D3) occurs approximately 100 m above the base of the section, at a flooding surface, where can be identified as the boundary of the D2 and D3 subunits (Fig. 2). The phosphatic-dolomitic siliceous mudstone transitions into dolomitic mudstone and is overlain by lime mudstone (Fig. 2). The uppermost D3 subunit at the

Fig. 2. Profiles of δ13C (‰-VPDB) and δ18O (‰-VSMOW) of carbonate and δ18O (‰-VSMOW) of silica phases in the D2, D3 and D4 subunits of the Doushantuo Formation, at NXFH and SXFH sections in the Three Gorges area, South China. Stratigraphic columns on the left show the main rock types of the Doushantuo and Dengying Formations. The 0 m datum is the top of the Nantuo diamictites in section NXFH, and the erosional surface in section SXFH.


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Fig. 3. Field photographs of the Doushantuo Formation overlying the cap carbonate: (A) Chert nodule within the dolostone at NXFH (yellow line). (B) Chert nodules in Sample S08 at SXFH. Yellow points indicate the locations of subsamples for δ18O analysis of chert, and δ18O and δ13C analysis of the coadjacent calcite. Calcite is located outside of the chert nodule at point 2, and inside of the chert nodule at point 3. (C) Erosional surface (yellow arrows) overlain by a black shale unit in the SXFH section, correlatable with the NXFH section. (D) Chert band at SXFH (yellow arrows). (E) Banded dolostone without chert nodules at SXFH. (F) Chert nodules in the dolostone at SXFH.

NXFH section is characterized by massive cherty limestone with thin (possibly microbial) laminations interbedded with argillaceous limestone. Generally, the upper part of the NXFH section contains greater abundances of carbonate, phosphatic clasts, and silica than the Jiulongwan section, but less shale (Xiao et al., 2012), which suggests that the Doushantuo Formation at NXFH was deposited in shallower water and closer to the locus of carbonate production (Jiang et al., 2011; Zhu et al., 2013; Cui et al., 2015; Hohl et al., 2015). Xiao et al. (2012) proposed that the 140-m-thick succession of the Doushantuo Formation at NXFH can be correlated with the ~70 m of the lower Doushantuo Formation exposed at Jiulongwan, based on bio- and chemostratigraphic data, and suggested that the carbonates of the Doushantuo Formation at XFH may have been deposited in non-euxinic (oxic or ferruginous) waters, in a shallow environment surrounding a deeper euxinic lagoon. Hohl et al. (2015) proposed that the shallow marine depositional

environment at XFH was affected by several episodes of freshwater influx, on the basis of isotopic and REE + Y data. The Doushantuo Formation is well exposed along a road at the SXFH section. The black shale with erosional surface at approximately 10 m above the base of the SXFH section can be correlated with that at about 100 m of the NXFH section (Figs. 2 and 3C; Zhu et al., 2013). The D3 subunit overlying the erosional surface is composed of thinly bedded muddy dolomite, with phosphatic-chert nodules (Fig. 2), the abundance of which decrease upsection. This lithology is overlain by a black shale that transitions into dolomite interbedded with cherty limestone. Member VI (D4) is composed of black shale interbedded with muddy limestone and dolostone (Fig. 2). The overlying Dengying Formation is composed of the bedded dolostones. It should be noted that the boundary between the Doushantuo and Dengying formations is not clear in outcrops, due to the similarity in lithology and color (Zhu et al., 2013).

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3. Analytical methods Bulk rock samples were initially crushed into chips, with a subset of clean chips free of surface alteration selected for further analysis. The selected chips were then further crushed in a stainless steel mortar and pulverized in an agate mill. The samples containing nodules and bands were cut into thin slices and polished. Chert or calcite nodules and bands can be identified on the basis of colors, and separated from the dolostone host rock by hand or toothdrill. 3.1. XRD analysis Approximately 5 g of powdered bulk rocks was treated with an excess of 10% HCl at room temperature for 24 h to completely remove carbonate phases. It should be noted that the volumes of acid are depended on the weight of powder and carbonate contents and must be enough to completely remove the carbonate phases. The residual insoluble silica phases were rinsed three times using distilled water, and dried at 110 °C. Due to quartz particles settling faster than clay minerals during the rinsing process, we collected only the lowermost part of the suspension liquid, to remove the majority of clay minerals. The mineralogical characteristics of the insoluble materials were determined using X-ray diffractometry (XRD) techniques as a semi-quantitative check, which was carried out at University of Science and Technology of China in Hefei. The samples were scanned with a Rigaku D/max-IIIa diffractometer equipped with a Cu-target tube and a curved graphite monochromator, operating at 37.5 kV and 20 mA. Samples were step-scanned from 3° to 70° with a step size of 0.02° (2θ). The side-packing method, proposed by the National Bureau of Standards (NBS), was used to prepare the XRD samples.

compositions of both bulk rocks and subsamples were analyzed using a GasBench II in continuous flow mode, with carbonates digested in concentrated phosphoric acid at 72 ± 1 °C (Zha et al., 2010). The extracted CO2 was measured on a Finnigan MAT253 mass spectrometer at the CAS Key Laboratory of Crust-Mantle Materials and Environments, University of Science and Technology of China in Hefei. Details of this procedure and analytical techniques were described by Zha et al. (2010). Both carbon and oxygen isotopic compositions are reported in delta notation relative to the Vienna Pee Dee Belemnite (VPDB) standard. Precision and reproducibility of δ13C and δ18O analyses are better than ±0.2‰ based on replicate analyses of internal standards and samples. Two calcite reference materials were used to monitor daily isotope analyses: International Standard NBS-19, with δ13C = + 1.95‰ and δ18O = − 2.20‰, and National Standard of China GBW04417, with δ13C = −6.06‰ and δ18O = −24.12‰. It should be noted that the δ18O values of carbonates are reported using the Vienna Standard Mean Ocean Water (VSMOW) standard during our discussion. Values were converted between the VPDB-VSMOW standards using the equations δ18O (VSMOW) = 1.03091 × δ18O (VPDB) + 30.91 and δ18O (VPDB) = 0.97002 × δ18O (VSMOW) − 29.98 (Coplen et al., 1983; Hoefs, 2009).

Selected fragments of chert nodules and bands were treated with 10% HCl to completely remove carbonate. The treated fragments were carbon coated, and imaged using a FEI Sirion 200 scanning electron microscope (SEM) at the University of Science and Technology of China in Hefei. An accelerating voltage of 15 kV and a working distance of 10 mm were used during sample imaging.

3.4.2. Oxygen isotopic compositions of silica phases and chert nodules/ bands Oxygen isotope analyses were conducted on residual silica phases and separated chert nodules/bands after they were treated with HCl to completely remove carbonate phases. An aliquot of 1.1 to 1.8 mg of sample powder was analyzed via the laser fluorination technique, using a 25 W MIR-10 CO2 laser at the Laboratory for Chemical Geodynamics, University of Science and Technology of China in Hefei. Oxygen was directly transferred to a MAT253 mass spectrometer for measurement of the 18O/16O and 17O/16O ratios (Zheng et al., 2002). Oxygen isotope data are reported in delta notation relative to Vienna Standard Mean Ocean Water (VSMOW). The quartz reference mineral GBW04409, with a δ18O value of +11.11‰, was used to test reproducibility. The reproducibility of repeated δ18O measurements for each standard or sample on a given day was better than ±0.1‰ (1σ). All oxygen isotope results are listed in Appendix Table 1.

3.3. Trace and rare earth element analysis

4. Petrography

To completely remove the carbonate component prior to analysis, powdered bulk rock samples were treated with 10% HCl at 60 °C for 24 h, after which the silica phases were collected and rinsed three times using distilled water, then dried at 110 °C. Residual silica phases were then weighed, to determine the percentage of carbonate in the original sample. Due to the high organic matter content, the trace elements in the silica phases were analyzed using ALS Minerals ME-MS81 procedure. The silica phases were mixed with lithium metatorate/tetroborate flux, and fused in a furnace at 1025 °C. The resulting melt was then cooled and dissolved in an acid mixture containing nitric, hydrochloric and hydrofluoric acid. The solution was then analyzed via the inductively coupled plasma-mass spectrometry (ICP-MS) using a Perkin Elmer Elan 9000 at ALS Chemex Co., Ltd. in Guangzhou, China. Standard reference materials (OREAS 120, STSD-1, SY-4 and TRHB) were measured during the analytical process, and measurement accuracy was determined to be within 10%.

Doushantuo Formation strata overlying the cap carbonate at both NXFH and SXFH are composed of 7 to 99 wt.% carbonate minerals and 1 to 93 wt.% HCl-insoluble materials, including quartz, phosphate, pyrite, clays and organic matter (Xiao et al., 2012; Hohl et al., 2015). In the field, fresh surfaces of carbonates are gray to dark-gray, and sometimes gray-yellow (Fig. 3). Laminations and bedding planes are apparent in the dolostone matrix (Fig. 3C–E). Chert nodules and bands are dark gray to black in color (Fig. 3B) due to fine-grained opaque materials, such as pyrite, organic matter, clay, and phosphatic grains (Xiao et al., 2012). Chert nodules are spherical to oblate in shape and occur sporadically throughout the section, but are generally oriented parallel to bedding (Fig. 3B and F). Chert bands are also parallel to bedding and laminations (Fig. 3D). The laminations and bedding planes are locally disrupted or compacted around the chert nodules (Fig. 4A and F), and are sometimes apparent within the chert nodules (Xiao et al., 2012), suggesting that the chert nodules and bands probably formed in situ prior to compaction.

3.4. Stable carbon and oxygen isotope analysis

4.1. Host dolostone and calcite cements

3.4.1. Carbon and oxygen isotopic compositions of carbonate The carbon and oxygen isotopic compositions of host dolostone matrix (C1), carbonate nodules (C2), and carbonates associated with chert (C3 and C4) were analyzed separately. The carbon and oxygen isotopic

The host dolostone matrix of the Doushantuo Formation overlying the cap carbonate at XFH (C1) is mainly composed of a microcrystalline to microsparry dolomite (Fig. 2, Xiao et al., 2012; Hohl et al., 2015), without observable recrystallization in thin section (Fig. 4A, C and G).

3.2. SEM


One generation of calcite cements, which stains red with alizarin, is present in the dolostone matrix (Fig. 4B, D, E and H). The calcite cements is present in five distinct forms: sparry calcite nodules (C2, Fig. 4C and H), carbonates around chert nodules (C3, Fig. 4D), cements dispersed within chert nodules (C4, Fig. 4E), fine micropeloidal calcite cements (C5, Fig. 4E), and calcite veins (C6, Fig. 4F). In thin section, the cements are mainly composed of equant, medium- to coarsely- crystalline calcite, white in the outcrops (Fig. 4B–H). It should be noted that calcite nodules are typically oriented parallel to bedding, and range from one to several mm in diameter. C1, C2, C3 carbonates can be sampled by drilling the white part of hand specimens, while disseminated C4 cement can be sampled by drilling the chert nodules. However, calcite cements (C5) or veins (C6) cannot be sampled for geochemical analyses due to their diffuse distribution (Fig. 4D) although they can be observed in host rocks under a microscope (Fig. 4A, C and D). 4.2. Silica phases and chert nodules/bands Five distinct types of silica phases (S1 to S5) can be identified in thin sections. S1 is microcrystalline quartz found in chert nodules, with darkcolored and irregular boundaries. The nodules can be up to several centimeters in diameter, and are composed of equant or anhedral mosaics fabrics (Fig. 4A, B, E and G). Calcite relics (Fig. 4E) and apatite (Fig. 4B) can be observed in the chert nodules, suggesting that the chert probably replaced original calcite. S2 is the most abundant silica type, mainly composed of microcrystalline quartz disseminated within the host rock (Fig. 4C and D) without replacement of adjacent carbonate relics of previous structures. S2 usually appears as pale gray birefringent material under polarized light, or is characterized by an irregular, sweeping extinction with defined, crenulated crystal boundaries (Marin et al., 2010). S3 is drusy megaquartz disseminated within the dolostone, often cross-cut by calcite veins (Fig. 4A and C). S4 is the microcrystalline chert or megaquartz replacing rhombic calcite around its edges or interior (Fig. 4C and D). S5 is radial spherulitic quartz or chalcedony with irregular sweeping extinction in plane-polarized light (Fig. 4H). The petrographic evidence shows that S1, S4 and S5 formed via the replacement of previous calcite cements. The S1, incorporating relict calcite C4 (Fig. 4B), may represent a continuous aggregation of S4 and S5, as crystal morphology changed in response to progressive restriction of the fluid system during early burial. The distribution features suggest that S1, S4 and S5 postdate the calcite cements (including C2, C3, C4 and C5). On the other hand, S2 is disseminated within the host dolostone matrix without any replacement of calcite, which suggests that S2 was precipitated directly from the porewater of the carbonate matrix. S3 is cross cut by calcite veins (C6, Fig. 4A and C), indicating that S3 also precipitated from porewater, and predated the calcite veins. The finegrained and laminated carbonate matrix is typically deformed around the early diagenetic chert nodules of S1 and S3 (Fig. 4A and G), implying that the formation of these phases predated extensive burial compaction. Thus, the S1–S5 phases probably precipitated during early burial. With this constraint, the silicification of the D2 to D4 subunits probably occurred within several meters of the sediment-water interface, and S1–S5 may preserve information about the diagenetic history of the host dolostone. Large microbial mat fragments can be observed in the centers of some chert nodules (Xiao et al., 2010). Three known taxa of acanthomorphic acritarchs, Eotylotopalla delicata, Ericiasphaera magna and Polygonium cratum, have been reported from the chert horizons (Yin et al., 2007; McFadden et al., 2009). After HCl treatment, the chert nodules and bands contain amorphous or microcrystalline quartz with variable porosity (Fig. 5). XRD analysis only detects quartz in the silica phases (Fig. 6), indicating that other components, such as phosphate and clay materials, were relatively insignificant. Therefore, we subsequently used quartz interchangeable with the silica phases in our discussion. The δ18OQz values

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referred not only to the oxygen isotopic composition of chert nodules and beds, but also to the compositions of silica phases more generally. 5. Geochemical results 5.1. Stable carbon and oxygen isotopic compositions No samples between 6.3 and 49.3 m at the NXFH section were suitable for carbonate isotopic analysis, because the exposed outcrop is mainly composed of black shale interbedded with fine-grained phosphatic siltstone. The carbonate matrix in overlying strata, from 49.3 m to the top of the section, have δ13C values ranging from − 1.39 to + 6.27‰, and δ18O values ranging from +18.77 to +30.33‰ (Appendix Table 1). The δ13C curve of the D2 subunit shows a negative shift from + 6.44 to + 1.98‰, then back to + 6.01‰ (Fig. 3A). The D3 subunit through lower D4 interval exhibits a negative shift from +6.17 down to −1.39‰, followed by a positive shift in the upper part of D4 subunit (Fig. 3A and B). The δ18O values appear to be independent of stratigraphic position, and no stratigraphic trends are apparent from D2 to D4 subunits. These results are consistent with those previously reported in Zhu et al. (2013), Xiao et al. (2012) and Liu et al. (2013). Selected subsamples collected by drilling, including carbonate nodules (C2), carbonates around chert nodules (C3), and carbonate within chert nodules (C4), were measured to compare with those of host dolostone matrix (C1). These subsamples have δ13C values ranging from + 2.23 to + 5.14‰ in the D2 subunit, + 1.03 to + 6.85‰ in the D3 subunit and −1.09 to +7.97‰ in the D4 subunit (Appendix Table 1). The δ18O values range from +18.66 to +22.93‰ in the D2 subunit, from + 19.47 to + 27.88‰ in the D3 subunit and from + 18.77 to + 23.52‰ in the D4 subunit. In general, the δ18O values of these subsamples are relatively low compared with the corresponding dolostone matrix (Fig. 7). The differences in δ18O values between the subsamples and their coadjacent dolostone matrix vary from −9.84 to −0.76‰ (Appendix Table 1). No consistent pattern in δ13C values relating the subsamples to the coadjacent dolostone matrix can be observed. The δ18O values of the silica phases (δ18OQz) range from +22.27 to +24.68‰ in the D2 subunit, from +21.04 to +27.89‰ in the D3 subunit and from +24.79 to +26.31‰ in the D4 subunit (Appendix Table 1). Only a few chert nodules and bands could be recognized based on the distinct color in outcrops, and have been analyzed for their oxygen isotopic compositions. The other silica phases can be recognized in thin sections, but are too finely disseminated to be separated from one another for isotopic analysis. The sampled chert nodules and beds have δ18O values ranging from + 20.12 to +28.27‰ (Appendix Table 1), which are similar to, or slightly higher than, the δ18O values of the various other silica phases. This indicates that the different silica phases shared similar silica sources and depositional conditions, because the other silica phases are mixtures of different silica phases with varying proportions. 5.2. Trace and rare earth element concentrations Trace element concentrations in cherts may record the influence of hydrothermal, terrestrial or marine inputs on fluids in the basin (Owen et al., 1999b; Friend et al., 2008; Shen et al., 2011). REE concentrations were normalized against Post Archean Australian Shale (PAAS; Taylor and McLennan, 1985) to remove the odd-even variability seen in raw elemental concentrations, and to facilitate comparison with the average continental weathering input to the hydrosphere. Variations within the shale-normalized REE + Y pattern are commonly expressed either as anomalies relative to the adjacent elements (e.g. for Ce and Eu peaks), or as ratios between the elements (e.g. the Y/Ho ratio). The degree of light REE (LREE), middle REE (MREE) and heavy REE (HREE) enrichment is assessed through the ratios of PAAS-normalized Pr, Sm and Yb values.

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The silica phases of the D2 subunit have ∑REE concentrations from 4.8 to 6.4 ppm (Appendix Table 2). The PAAS-normalized REE + Y patterns (Fig. 8A) are characterized by: (1) LREE depletions, with [Pr/Yb]N

from 0.14 to 0.19 and [Sm/Yb]N from 0.17 to 0.23; (2) minor positive La anomalies, with [La/La*]N of 1.03 to 1.17, except for one sample with a [La/La*]N of 0.92; (3) slight negative Ce anomalies with [Ce/Ce*]N of


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Fig. 5. SEM images of chert nodules after HCl treatment. (A) Silica shows typical polyhedral blocks with pyramidal termination in sample S01. The euhedral habit of the quartz crystals and the nature of the intercrystalline porosity which was probably originally filled with calcite. Amorphous and sub-microcrystalline quartz is shown in (B) and (C) of sample S10, and (D) of sample S26. The carbonate components have been dissolved. Note the tight aggregation of silica minerals in (B) and loose aggregation in (C) and (D). Si = silica.

0.84 to 0.97; (4) positive Eu anomalies with [Eu/Eu*]N of 1.25 to 1.62; (5) slight negative Gd anomalies, with [Gd/Gd*]N of 0.84 to 0.92; and (6) Y/Ho ratios from 35.5 to 43.0. The Th/U ratios range from 0.37 to 0.73. The silica phases of the D3 subunit have ∑ REE concentrations from 2.3 to 27.3 ppm (Appendix Table 2). The PAAS-normalized REE + Y patterns include (Fig. 8B): (1) LREE depletions, with [Pr/ Yb]N from 0.08 to 0.29 and [Sm/Yb] N from 0.04 to 0.29; (2) minor positive La anomalies with [La/La*]N of 1.03 to 1.38; (3) slight negative Ce anomalies, with [Ce/Ce*] N of 0.76 to 0.97; (4) positive Eu anomalies with [Eu/Eu*]N of 1.37 to 2.72; (5) negative Gd anomalies with [Gd/Gd*]N of 0.58 to 0.93; and (6) Y/Ho ratios are from 26.3 to 47.5. The Th/U ratios range from 0.22 to 1.61, except for one sample with a value of 2.59. The silica phases of in the D4 subunit have ∑ REE concentrations ranging from 2.8 to 10.9 ppm (Appendix Table 2). The PAAS-normalized REE + Y patterns (Fig. 8C) typically show: (1) obvious LREE depletions, with [Pr/Yb]N from 0.11 to 0.24 and [Sm/Yb]N from 0.07 to 0.25; (2) minor positive La anomalies, with [La/La*]N of 1.15 to 1.41, except for one sample with a [La/La*]N of 0.97; (3) negative Ce anomalies with [Ce/Ce*]N of 0.75 to 0.89, excepting one sample with a [Ce/Ce*]N of 1.09; (4) positive Eu anomalies with [Eu/Eu*]N of 1.01 to 1.72; (5) negative Gd anomalies with [Gd/Gd*]N of 0.68 to 0.93; and (6) the Y/Ho ratios ranging from 32 to 37.8. Th/U ratios range from 0.52 to 1.98, except a single sample.

6. Discussion 6.1. δ18O values of Ediacaran surface seawater The oxygen isotopic compositions of carbonate and chert are controlled by both the δ18O of water (δ18OW) and the formation temperature (T) (Hoefs, 2009; Cunningham et al., 2012), allowing the δ18O value of ancient seawater to be calculated if the depositional temperature could be constrained. This technique assumes that the mineral precipitated directly from surface seawater, was in isotopic equilibrium with that seawater, and preserved the primary geochemical signals. The following calculation for quartz (Knauth and Epstein, 1976), calcite (O'Neil et al., 1969) and dolomite (Vasconcelos et al., 2005) will use δ18O values relative to the VSMOW standard: 1000 lnαQz−W ¼ 3:09 106 T−2 −3:29

ð1Þ

1000 lnαCal−W ¼ 2:78 106 T−2 −2:89

ð2Þ

1000 lnαDol−W ¼ 2:73 106 T−2 þ 0:26

ð3Þ

where T is the formation temperature of quartz, calcite and dolomite. The δ18O values of the dolostone matrix in the Doushantuo Formation range from +19.67 to +31.34‰, with most values between +23

Fig. 4. Thin section photomicrographs: (A) Two chert nodules composed of microcrystalline quartz (S1) and megaquartz (S3) in sample S14. Note deformation of micro-laminations around the nodule (yellow arrows) following compaction. Note how the onlap and wedge-shaped structures typical of truncation on an uneven surface. (B) Magnified view of chert nodules in the yellow square in panel (A). Note that there is no continuous micro-lamination running into the nodule. (C) Microcrystalline quartz (S2) disseminated in the dolomite and microcrystalline quartz (S4) and replacing calcite (C1) from the edges (yellow arrows) in sample N32. (D) Microcrystalline quartz (S4) replacing the calcite cement (C3) in sample N29. (E) Chert nodules (S1), calcite nodules (C2), calcite within chert nodules (C4), and micropeloidal calcite cements (C5, stained red) in sample S27. (F) Megaquartz (S3) in sample N34. Note the calcite vein (C6, yellow arrows) cross-cutting the megaquartz. (G) Chert nodule with microcrystalline quartz (S1) and calcite (C4) in sample S08. Note the deformation of micro-laminations around the nodule due to the compaction. (H) Photomicrograph of N23, showing the radial spherulitic quartz or chalcedony (S5) growing within the calcite cement (C2).

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Fig. 7. Comparison of δ18O values (VSMOW) of subsample C2, C3 and C4 to values from the dolostone matrix (C1). The δ18O values (VSMOW) from the various subsamples are consistently lower than those from the dolostone matrix.

Fig. 6. X-ray diffractograms for (A) Sample N32 (subunit D2) and Samples (B) N04, (C) S01 and (D) S08 (subunit D3). Note that only quartz is detectable in the acid-insoluble silica phases.

and +30‰ (Appendix Table 1). Neoproterozoic seawater temperatures are estimated to have been between 10 and 32 °C (Jaffrés et al., 2007). Ling et al. (2004) estimated the depositional temperature of the Doushantuo Formation from 32.2 to 34.0 °C, based on the δ18O values of dolomite and phosphate. Meng et al. (2011) measured the maximum homogenization temperatures of primary fluid inclusions in Ediacaran evaporites, estimating that the highest tropical seawater temperatures during Ediacaran Period would have reached 39.4 ± 1.0 °C. Bonifacie et al. (2008) suggested that the maximum temperature of the Ediacaran ocean was 28 °C, based on the clumped isotope carbonate paleothermometry, while Derkowski et al. (2013) estimated that the dolomite in the D2 subunit was deposited at temperatures b 30 °C. Thus, we can reasonably constrain the temperature of Ediacaran seawater at XFH to between 28 and 39 °C.

Ediacaran ocean δ18O values can be calculated based on Eq. (3) and the δ18O values of dolostone matrix. We use the relative frequency of calculated ocean δ18O values to constrain the most credible value for Ediacaran surface seawater (Fig. 9). If the seawater temperature was 28 °C, the calculated oceanic δ18O values would range from − 6 to −2‰ for the D2 subunit (Fig. 9A), from −4 to −1‰ for the D3 subunit (Fig. 9B) and from −2 to 0‰ for the D4 subunit (Fig. 9C). It appears that the δ18O slowly increased from the D2 through D4 subunit. If the seawater temperature is assumed to be 39 °C, the calculated oceanic δ18O values would be from − 4 to 0‰ for the D2 subunit (Fig. 9D), from − 2 to + 1‰ for the D3 subunit (Fig. 9E), and from 0 to + 2‰ for the D4 subunit (Fig. 9F). Taking the full range of possible water temperatures into account, we can constrain oceanic δ18O from −6 to 0‰ for the D2 subunit, from −4 to +1‰ for the D3 subunit, and from −2 to +2‰ for the D4 subunit. Although the oxygen isotopic composition of the Precambrian oceans remains controversial, several previous studies have suggested that the δ18O could have been lower than − 8.4‰, possibly as low as − 10‰ (Veizer et al., 1997; Veizer, 1999; Wallmann, 2001; Shields and Veizer, 2002; Jaffrés et al., 2007). Wallmann (2001) modeled the geological water cycle and determined that the Precambrian oceans would have had a δ18O of around + 3‰. Zhao and Zheng (2010) found Ediacaran carbonates with unusually low δ18O values of +4.52 to +11.74‰ in southern Anhui Province of South China, and suggested that the carbonate precipitated from a mixture of continental deglacial meltwater and seawater (Wang et al., 2014a, 2014b). Other studies have suggested that oceanic δ18O values were buffered at a value of 0 ± 2‰ (Land and Leo Lynch, 1996; Muehlenbachs, 1998; Lécuyer and Allemand, 1999; Joachimski et al., 2004). The δ18O values of Edicaran carbonates in South China are likely influenced by many factors, but there is no reason to expect Ediacaran


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probably represent the maximum range of δ18O variations for Ediacaran surface seawater. In fact, the δ18O values of the Ediacaran surface seawater would likely be relatively stable (Land and Leo Lynch, 1996; Muehlenbachs, 1998; Lécuyer and Allemand, 1999; Joachimski et al., 2004). 6.2. Diagenetic temperatures Oxygen isotope analyses of quartz-carbonate pairs can be used to estimate depositional or early diagenetic paleotemperatures. The fractionation equations below show that if quartz is in isotopic equilibrium with calcite/dolomite, the quartz will have higher δ18O:

Fig. 8. Diagrams of PAAS-normalized REE patterns for the silica phases of (A) subunit D2, (B) subunit D3 and (C) subunit D4 of the Doushantuo Formation overlying the cap carbonates.

seawater to have had a δ18O value N 0‰ typical of modern seawater (Veizer et al., 1997; Veizer, 1999; Wallmann, 2001; Shields and Veizer, 2002; Jaffrés et al., 2007). If we discount estimates N0‰, we can constrain the Ediacaran surface seawater δ18O values at XFH to between − 6 and 0‰ for the D2 subunit, − 4 and 0‰ for the D3subunit, and − 2 and 0‰ for the D4 subunit. It should be noted that these values

Δ18 OQz−Cal ¼ 0:31 106 T−2 −0:4

ð4Þ

Δ18 OQz−Dol ¼ 0:36 106 T−2 −3:55

ð5Þ

where Δ18OQz-Cal is the equilibrium oxygen isotope fractionation between quartz and calcite, and Δ18OQz-Dol is the equilibrium oxygen isotope fractionation between quartz and dolomite. For bulk samples at XFH (Fig. 10A), the oxygen isotopic compositions of both dolomite (δ18ODol) and silica phases (δ18OQz) vary widely, from +16.67 to +31.34‰ and from +19.0 to +30.98‰, respectively. Fractionation (Δ18OQz-Dol) values range from − 8.56 to + 3.35‰ (Appendix Table1). Those quartz-dolostone pairs with Δ18OQz-Dol values b 0‰ are clearly not in isotopic equilibrium (Fig. 10A), suggesting that the silica phases precipitated after the matrix dolostone when the lower δ18OQz values could result from higher temperature or lower δ18O values in pore fluid. Those quartz-dolostone pairs with Δ18OQz-Dol values N 0‰ most likely are isotopic equilibrium. For those pairs with δ18O values near the 0 °C isotherm fractionation line, quartz possibly precipitated directly from the surface seawater along with the dolostone matrix. No positive correlations are observed between δ18ODol and δ18OQz in subunits D2 and D3. However, there is a positive correlation in δ18ODol and δ18OQz of the D4 subunit (Fig. 10A), suggesting that meteoric water was involved in early diagenesis. A low-δ18O meteoric water input would have contributed to the observed 18O-depletion in chert and coadjacent carbonate (Abruzzese et al., 2005; Hohl et al., 2015). In chert nodules and bands (Fig. 10B), δ18O values range from + 20.12 to + 28.75‰ and those of the coadjacent calcite range from + 17.06 to + 31.34‰ (Appendix Table 1). It should be noted that the range of variability of δ18O in both chert and calcite is greater in the D3 and D4 subunits than in the D2 subunit, suggesting larger variations in temperatures and/or fluid δ18O values in the D3 and D4 subunits. The fractionation (Δ18OQz-Cal) values range from −4.83 to +9.75‰ (Appendix Table 1). Again, those quartz-calcite pairs with Δ18OQz-Cal values lower than 0‰ are almost certainly not in isotopic equilibrium (Fig. 10B). However, many Δ18OQz-Cal values are higher than 0‰, with δ18O values falling above the 0 °C isotherm fractionation line, which suggests that these cherts formed at lower temperatures or in more 18O– enriched fluids than the coadjacent calcite. This interpretation seems to contradict the petrographic evidence, which indicates that the silica of the S1–S5 phases precipitated after the C2, C3 and C4 carbonate phases, since one would expect temperature to increase with burial depth. The δ18O values are also inconsistent with a high-temperature hydrothermal origin. Thus, the cherts most likely precipitated from fluids with high δ18O values relative to those from which calcite phases precipitated, implying that the diagenetic fluids were derived from seawater rather than meteoric water. The δ18O values of the fluids which precipitated chert was likely similar to, or slightly lower than, that of the overlying surface seawater (Behl and Smith, 1992; Raiswell and Fisher, 2000; Loyd et al., 2012; Zhao and Zheng, 2013, 2015). Using this constraints, we can calculate diagenetic temperatures ranging from 23 to 74 °C (with the

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Fig. 9. Frequency distribution of calculated surface seawater δ18O values of (A) subunit D2, (B) subunit D3 and (C) subunit D4 at a temperature of 28 °C, and (D) subunit D2, (E) subunit D3 and (F) subunit D4 at a temperature of 39 °C.

most frequent calculated temperatures (MFTs) ranging from 27 to 60 °C) for the D2 subunit (Fig. 11A and B), 19 to 90 °C (with MFTs from 53 to 60 °C, and only one estimate of 90 °C) for the D3 subunit (Fig. 11C and D), and 29 to 40 °C for the D4 subunit (Fig. 11E and F). Thus, we can constrain the maximum range of credible diagenetic temperatures of the D2, D3 and D4 subunits to between 27 and 60 °C (Fig. 11). Numerous studies have reported estimates of the diagenetic temperature of chert formation. Loyd et al. (2012) calculated a diagenetic temperature of 17 to 35 °C for concretion formation in the Miocene Monterey Formation, California. Behl and Smith (1992) estimated a formation temperature of 7 to 26 °C for carbonate-replacing quartz cherts, 22 to 25 °C for bedded cherts, and 32 to 34 °C for hydrothermal cherts from Ocean Drilling Program Leg 129. Silica in a Mesozoic mud mound is reported to have precipitated at a temperature of 25 to 30 °C in a shallow burial environment (Floquet et al., 2012). The δ18O values of chert and coadjacent carbonate in the Doushantuo Formation overlying the cap carbonate at XFH indicate that there is no hydrothermal influence. Therefore, the calculated diagenetic temperature of up to 60 °C would be the highest diagenetic temperature experienced by the dolostone matrix in the D2, D3 and D4 subunits of the Doushantuo Formation. The diagenetic temperatures calculated for the dolostone matrix, ranging from 27 to 60 °C, correspond to vitrinite reflectance values of around 0.2 to 0.3 (Machel et al., 1995), and are consistent with shallow burial in the nitrate and iron reduction zones, rather than at depths experiencing sulfate reduction and methanogenesis (Claypool and Kaplan, 1974; Wehrmann et al., 2014). Thermochemical sulfate reduction (TSR) to sulfide is sufficiently slow at such low temperatures as to be insignificant, even over tens of millions of years. This indicates that the sulfur isotopic compositions of pyrite in the host carbonates (δ34SPy) were likely not altered at the low temperatures they experienced.

The carbon isotopic composition of carbonates is most strongly controlled by the δ13C of dissolved inorganic carbon (DIC), and only slightly affected by temperature (less than −0.1‰/°C; Emrich and Vogel, 1970; Romanek et al., 1992). Thus, if the δ13C values of dolostone (δ13CCarb) reflect the original DIC pool, one would expect to see little systematic variability due to temperature. This inference is supported by the lack of an apparent relationship in δ13C values between the host dolostone matrix and drilled subsamples from the same sample (Appendix Table 1). Additionally, REE + Y patterns have been reported to survive severe metamorphism (Banner et al., 1988). Therefore, the geochemical compositions of dolostone matrix at XFH, including δ13CCarb values, δ34SPy values and REE + Y patterns, are unlikely to have been significantly altered at the diagenetic temperatures they experienced. 6.3. Depositional and diagenetic environments inferred from REE systematics Rare earth element (REE) chemistry is a valuable tool for investigating early sedimentary environments because the relative abundances of REE in hydrogenous sediments, such as carbonate, phosphate and chert, vary systematically depending on the influence of hydrothermal, terrestrial or marine inputs into the depositional basin (Van Kranendonk et al., 2003; Jiang et al., 2007; Chang et al., 2008; Huang et al., 2009; Allwood et al., 2010; Vernhet and Reijmer, 2010; Eker et al., 2012; Wang et al., 2012). Murray (1994) demonstrated that REEs patterns in chert can distinguish between continental margin, proximal ridge and pelagic deposits. Typical marine sediments reflect a seawater REE + Y patterns with uniform LREE depletion, positive La anomalies, distinctively high Y/Ho ratios, and negative Ce anomalies in oxic seawater (Zhang and Nozaki, 1996; Webb and Kamber, 2000; Van Kranendonk et al., 2003; Bolhar and Van Kranendonk, 2007; Van den Boorn et al., 2010). Hydrothermal fluids display very different REE + Y patterns,


Fig. 10. Scatter plot of (A) δ18O values of dolostone matrix vs. coadjacent quartz phases for bulk rocks and (B) δ18O values of quartz and coadjacent calcite for chert nodules or beds of the D2, D3 and D4 subunits of the Doushantuo Formation. Isothermal oxygen isotopic fractionation lines for different temperatures are based on the theoretical calibrations of Knauth and Epstein (1976) and Vasconcelos et al. (2005) with Δ18OQz-Dol = 0.36 × 106 (T−2) − 3.55.

which have distinct Eu enrichment in an otherwise smooth, LREEenriched distribution (Michard et al., 1983; Michard and Albarede, 1986). In contrast, riverine REE + Y patterns are diverse, displaying slight depletions in LREEs or enrichments in MREEs, with no conspicuous single-element anomalies (Lawrence et al., 2006). These patterns have been identified in lacustrine and estuarine carbonates (e.g. Nothdurft et al., 2004; Bolhar and Van Kranendonk, 2007). Sedimentary carbonates associated with upwelling deep-ocean brine are typically characterized by enrichments in Fe and trace metals, with positive Eu anomalies (Meyer et al., 2012; Wang et al., 2014a, 2014b). Therefore, these systematic differences in pristine REE patterns make them a useful tool for exploring variation in mineral-precipitating fluids and environments at thermodynamic equilibrium (Owen et al., 1999a; Bolhar et al., 2004; Bolhar and Van Kranendonk, 2007; Friend et al., 2008; Eker et al., 2012). REEs in ancient sedimentary rocks can be readily contaminated by continental silicate detritus, and such contamination must be ruled out before REE data can be interpreted. For the silica phases at XFH, no correlation is observed between (LREE/HREE)N and Zr/Hf ratios (Fig. 12A), suggesting that no continental silicate contamination has occurred. However, a positive correlation between [Pr/Yb]N and ∑ REE (Fig. 12B) suggests that the REE + Y patterns of certain silica phases in the D3 and D4 subunits, with high ∑ REE concentrations, could have been contaminated by continental silicate material (Kamber et al., 2004; Frimmel, 2009; Ling et al., 2013). Owing to the relatively

161

high concentration of REEs in silicate minerals, even a small (e.g. 1– 2%) REE contribution from silicates could greatly reduce the La and Ce anomalies, the degree of LREE depletions, and the Y/Ho ratios of silica phases (Nothdurft et al., 2004; Frimmel, 2009; Ling et al., 2013). Thus, only those REE + Y patterns of silica phases with relatively low ∑REE concentrations are likely to reflect the REE + Y pattern of Ediacaran seawater. Primary REE + Y distributions in chert can withstand high water-torock ratio diagenetic processes (e.g. recrystallization, dolomitization), due to the typically low concentrations of REEs in formation fluids, as well as the simple chemical composition of chert and its low tolerance for substitutional and interstitial lattice imperfections (Murray, 1994; Owen et al., 1999a, 1999b; Knauth and Lowe, 2003; Abruzzese et al., 2005; Shen et al., 2011). The low diagenetic temperatures of the Doushantuo Formation would further facilitate the preservation of primary signals. Therefore, the geochemical signals preserved in cherts are unlikely to have changed substantially during diagenetic alteration. In modern sedimentary systems, early silicification can be explained via the redistribution of biogenic silica (opal A) within the sediment due to the chemical instability of biogenic opal. However, silica-secreting organisms were not present in the Precambrian ocean (Hesse, 1989; Siever, 1992; Maliva et al., 2005; Shen et al., 2011). Other sources of silica include the dissolution of detrital silicates (Peterson and Von der Borch, 1965), the release of silica during clay transformations (Mahran, 1999), and silica introduced via springs and seeps (Smith and Mason, 1991). The PAAS-normalized REE + Y patterns of those silica phases at XFH with relatively low Zr concentrations are typically characterized by LREE depletions, positive La and Eu anomalies, minor negative Ce anomalies, and slightly super-chondritic Y/Ho ratios (Fig. 8). Of these features, LREE-depletions, positive La anomalies, and superchondritic Y/Ho ratios are characteristic of modern seawater, as well as the chemical and biogenic sediments derived from seawater origin (Van Kranendonk et al., 2003; Nothdurft et al., 2004; Shields and Webb, 2004; Bolhar and Van Kranendonk, 2007; Friend et al., 2008; Van den Boorn et al., 2010; Ling et al., 2013; Wen et al., 2016). Cherts typically have higher Y/Ho ratios in the D2 (N 43) and D3 (N 42) subunits than in the D4 (N37.8) subunit, indicating that the silica phases in the D4 subunit probably reflect a greater freshwater input. This is consistent with the positive between δ18ODol and δ18OQz in bulk rocks (Huang et al., 2013; Hohl et al., 2015). The silica phases of the Doushantuo Formation at XFH show positive Eu anomalies. There is also a positive correlation between [Eu/Eu*]N and Ba concentrations in the D2 subunit (Fig. 12D), indicating that the positive Eu anomalies seen in the D2 subunit are probably influenced by Ba. No relationship between [Eu/Eu*]N and Ba concentrations is observed in the D3 and D4 subunits (Fig. 12D). In general, the positive Eu anomaly is a clear indication of either high-temperature (N 250 °C) hydrothermal alteration or extremely anoxic conditions (Bau and Alexander, 2009; Pi et al., 2013; Zhu et al., 2014). However, REE + Y patterns of silica phases throughout the XFH section are generally not consistent with those of hydrothermal effluents, which typically show LREE enrichment and a large positive Eu anomaly. In addition, the δ18O values of chert nodules and coadjacent calcite are not consistent with hydrothermal fluids, allowing us to rule out the possibility that silica phases were precipitated from hydrothermal fluid flux (Wang et al., 2012). Under intensely reducing conditions, Eu3+ can be reduced to Eu2+, which has a smaller ionic radius, allowing for easier substitution into mineral lattices. However, the relative abundance of Eu in silica phases is probably not a reflection of redox conditions, since no evidence of anoxic is observed in the Doushantuo Formation. A more likely explanation is an elevated supply of Eu to the depositional basin, via river, dust, and/or hydrothermal sources (Kamber et al., 2004), which is consistent with the elevated levels of CO2 in the atmosphere (Bolhar and Van Kranendonk, 2007), although this hypothesis needs further testing. Cerium can undergo oxidation in seawater, from soluble Ce3 + to highly insoluble Ce4 +, resulting in the partitioning of Ce from other

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Fig. 11. Frequency distribution of the calculated temperatures (T) of chert for the D2, D3, D4 subunits under their corresponding diagenetic fluid conditions. T range is based on the calculated temperatures from all δ18O values of chert. MFT is the most frequent calculated temperature.

REEs. Ce anomalies in chert and carbonate are mainly controlled by redox conditions, and have been used as a tracer to distinguish between oxic and anoxic depositional environments in the geological record (German and Elderfield, 1990; Guo et al., 2007; Ling et al., 2013; Xin et al., 2015, 2016). Negative Ce anomalies are observed in the silica

phases at XFH. This indicates substantial levels of dissolved oxygen even in marine porewater, which further suggests a high enough level of atmospheric oxygen to sustain aerobic organisms (Komiya et al., 2008; Frimmel, 2010; Ling et al., 2013; Fan et al., 2014; Wen et al., 2016).


163

Fig. 12. Cross plots showing correlations between (A) [LREE/HREE]N and Zr/Hf ratios, (B) Zr concentrations and [Pr/Yb]N ratios, (C) Zr concentrations and Y/Ho ratios, (D) Ba concentrations and [Eu/Eu*]N ratios for the D2, D3 and D4 subunits in the Doushantuo Formation.

6.4. Formation processes The concentration of dissolved silica in ocean was likely higher during the Precambrian (Hesse, 1989; Siever, 1992; Maliva et al., 2005; Shen et al., 2011). Bekker et al. (2010) suggested that organisms may have drawn down the concentration of silica in seawater to sufficiently low levels that silica was not absorbed into precipitating iron oxyhydroxides. Robert and Chaussidon (2006) noted that the δ30Si values of cherts increased from the Archean until approximately 600 Ma, then began to decrease toward the low modern value, consistent with the onset of biological control of the silica cycle at the beginning of the Phanerozoic. The chert nodules in the Doushantuo Formation, which frequently contain large microbial mat fragments, are consistent with silica mineralization around organic nuclei (McFadden et al., 2009; Xiao et al., 2010). Petrographic features suggest that the cherts in the dolostone of the Doushantuo Formation formed after initial cementation, but prior to consolidation and compaction. Silica replacement of calcite cements (Fig. 4C, D and E) implies that the pH of porewater changed during early diagenesis (Holdaway and Clayton, 1982; Floquet et al., 2012). Δ18OQz-Cal values above the 0 °C isotherm fractionation line and REE + Y patterns suggest that the cherts probably precipitated postdepositionally from seawater or a pore fluid with similar composition. The formation and preservation of silica within carbonates is favored by variations in pH (Siever, 1992; Bustillo, 2010; Perry and Lefticariu, 2007), which may have occurred due to anaerobic oxidation of organic matter through bacterial sulfate reduction (Siever, 1992; McFadden et al., 2008; Xiao et al., 2010, 2012; Léonide et al., 2014; Muscente et al.,

2015). The minor negative Ce anomalies observed in cherts likely indicate that the post-depositional fluids were at least slightly oxic (Komiya et al., 2008; Ling et al., 2013). Therefore, the cherts in the Doushantuo Formation most likely precipitated directly from seawater, or replaced the calcite cements in an early diagenetic porewater environment composed of primarily oxic seawater and influenced by the oxidation of multicellular acritarchs organisms (Xiao et al., 2002; Wang and Wang, 2006; Li et al., 2010). 7. Conclusions Several types of silica phases (S1 to S5) are observed in the dolostones of the Doushantuo Formation overlying the cap carbonate at XFH section, in the Three Gorges area of South China. The five types of silica phases include microcrystalline quartz in chert nodules and disseminated in the host dolostone (S1 and S2), megaquartz (S3), and microcrystalline quartz and fibrous silica replacing calcite cement (S4 and S5). Various types of carbonate are also present, including the host dolostone matrix (C1) and five types of calcite cements (C2 to C6) as observed in the thin sections. The sedimentary textures of the cherts indicate that silica precipitation postdated the calcite cements but predated burial compaction. The δ18O values show in most cases the carbonate and coadjacent chert are not in oxygen isotopic equilibrium. Based on the carbonate δ18O values, we estimated Ediacaran seawater to have had δ18O values ranging from −6 to 0‰, and surface seawater to have had temperatures ranging from 28 to 39 °C. Diagenetic temperatures calculated from the δ18O values of chert nodules or bands, range from 27 and 60 °C. Under such diagenetic conditions, primary geochemical

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signatures, such as δ13CCarb values, δ34SPy values, and REE + Y patterns may have been preserved in the host dolostones. The REE patterns of cherts are characterized by minor light REE depletions, positive La, Eu, and Gd anomalies, negative Ce anomalies, and slightly super-chondritic Y/Ho ratios, which indicate that REEs in the chert were derived from ambient seawater. Ce anomalies indicate that the bottom water or early diagenetic porewater during chert formation in the Doushantuo Formation were likely oxic. Acknowledgments This study was supported by funds from the Natural Science Foundation of China (41273008 and 41230102) and the National Basic Research Program of China (2013CB835003). We would like to express our gratitude to Pi Daohui and Liu Guangxin for their assistance during field sampling, and to Zha Xiangping and Li Weiping for their assistance during the stable isotope and element concentration analyses. Discussion with Dr. Huang Jing has greatly improved this manuscript. We would also like to thank Prof. Thomas Algeo, Prof. Kathleen McFadden and two anonymous reviewers for their comments on this manuscript, which greatly improved the presentation. Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.palaeo.2016.10.004. References Abruzzese, M.J., Waldbauer, J.R., Chamberlain, C.P., 2005. Oxygen and hydrogen isotope ratios in freshwater chert as indicators of ancient climate and hydrologic regime. Geochim. Cosmochim. Acta 69 (6), 1377–1390. http://dx.doi.org/10.1016/j.gca. 2004.08.036. Ader, M., Macouin, M., Trindade, R.I.F., Hadrien, M.H., Yang, Z., Sun, Z., Besse, J., 2009. A multilayered water column in the Ediacaran Yangtze platform? Insights from carbonate and organic matter paired δ13C. Earth Planet. Sci. Lett. 288 (1–2), 213–227. http:// dx.doi.org/10.1016/j.epsl.2009.09.024. Allwood, A.C., Kamber, B.S., Walter, M.R., Burch, I.W., Kanik, I., 2010. Trace elements record depositional history of an Early Archean stromatolitic carbonate platform. Chem. Geol. 270, 148–163. http://dx.doi.org/10.1016/j.chemgeo.2009.11.013. Banner, J.L., Hanson, G.N., Meyers, W.J., 1988. Rare earth element and Nd isotopic variations in regionally extensive dolomites from the Burlington-Keokuk Formation (Mississippian): implications for REE mobility during carbonate diagenesis. J. Sed. Petr. 58, 415–432. http://dx.doi.org/10.1306/212F8DAA-2B24-11D7-8648000102C1865D. Bau, M., Alexander, B.W., 2009. Distribution of high field strength elements (Y, Zr, REE, Hf, Ta, Th, U) in adjacent magnetite and chert bands and in reference standards FeR-3 and FeR-4 from the Temagami iron-formation, Canada, and the redox level of the Neoarchean ocean. Precambr. Res. 174 (3–4), 337–346. http://dx.doi.org/10.1016/j. precamres.2009.08.007. Behl, R.J., Smith, B.M., 1992. Silicification of deep-sea sediments and the oxygen isotope composition of diagenetic siliceous rocks from the Western Pacific, Pigafetta and East Mariana Basin, Leg 129. In: Larson, R.L., Lancelot, Y. (Eds.), Proceeding of the Ocean Drilling Program. Scientific Results, pp. 81–116. Bekker, A., Slack, J.F., Planavsky, N., 2010. Iron formation: the sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes. Econom. Geol. 105, 467–508. http://dx.doi.org/10.2113/econgeo.107.2.377. Bolhar, R., Van Kranendonk, M.J., 2007. A non-marine depositional setting for the northern Fortescue Group, Pilbara Craton, inferred from trace element geochemistry of stromatolitic carbonates. Precambr. Res. 155, 229–250. http://dx.doi.org/10.1016/j. precamres.2007.02.002. Bolhar, R., Kambera, B.S., Moorbathb, S., Fedoc, C.M., Whitehoused, M.J., 2004. Characterisation of early Archaean chemical sediments by trace element signatures. Earth Planet. Sci. Lett. 222, 43–60. http://dx.doi.org/10.1016/j.epsl.2004.02.016. Bonifacie, M., Eiler, J., Fike, D.A., 2008. Temperature and Oxygen Isotope Composition of the Ediacaran Ocean: Constraints From Clumped Isotope Carbonate Thermometry. American Geophysical Union, pp. 21B–1420 Fall meeting. Bristow, T.F., Bonifacie, M., Derkowski, A., Eiler, J.M., Grotzinger, J.P., 2011. A hydrothermal origin for isotopically anomalous cap dolostone cements from south China. Nature 474 (7349), 68–71. http://dx.doi.org/10.1038/nature10096. Bustillo, M.Á., 2010. Silicification of continental carbonates. In: Alonso-Zarza, A.M., Tanner, L.H. (Eds.), Carbonates in Continental Settings: Processes, Facies and Applications. Elsevier, Oxford, pp. 153–174. Chang, H.J., Chu, X.L., Feng, L.J., Huang, J., Zhang, Q.R., 2008. REE geochemistry of the Liuchapo chert in Anhua, Hunan. Geol. Chin. 35 (5), 879–887. Chen, J.B., Algeo, T.J., Zhao, L.S., Chen, Z.Q., Zhang, L., Yang, L., 2015. Diagenetic uptake of rare earth elements by bioapatite, with an example from Lower Triassic conodonts

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