Mesoproterozoic oxygenated deep seawater recorded by early diagenetic carbonate concretions from the Member IV of the Xiamaling Formation, North China

Journal Pre-proofs Mesoproterozoic oxygenated deep seawater recorded by early diagenetic car‐ bonate concretions from the Member IV of the Xiamaling Formation, North China Anqi Liu, Dongjie Tang, Xiaoying Shi, Xiqiang Zhou, Limin Zhou, Mohan Shang, Yang Li, Hao Fang PII: DOI: Reference:

S0301-9268(19)30362-6 https://doi.org/10.1016/j.precamres.2020.105667 PRECAM 105667

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Precambrian Research

Received Date: Revised Date: Accepted Date:

25 June 2019 1 February 2020 12 February 2020

Please cite this article as: A. Liu, D. Tang, X. Shi, X. Zhou, L. Zhou, M. Shang, Y. Li, H. Fang, Mesoproterozoic oxygenated deep seawater recorded by early diagenetic carbonate concretions from the Member IV of the Xiamaling Formation, North China, Precambrian Research (2020), doi: https://doi.org/10.1016/j.precamres. 2020.105667

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1

Mesoproterozoic oxygenated deep seawater recorded by early diagenetic

2

carbonate concretions from the Member IV of the Xiamaling Formation, North

3

China

4 5

Anqi Liua, b, Dongjie Tang*a, b, Xiaoying Shia, c, Xiqiang Zhoud, Limin Zhoue, Mohan

6

Shangc, Yang Lic, Hao Fangc

7 8

aState

9

of Geosciences (Beijing), Beijing 100083, China

Key Laboratory of Biogeology and Environmental Geology, China University

10

bInstitute

11

100083, China

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cSchool

13

Beijing 100083, China

14

dKey

15

Geophysics, Chinese Academy of Sciences, Beijing 100029, China

16

eNational

of Earth Sciences, China University of Geosciences (Beijing), Beijing

of Earth Sciences and Resources, China University of Geosciences (Beijing),

Laboratory of Petroleum Resources Research, Institute of Geology and

Research Center for Geoanalysis, Beijing 100037, China

17 18

*Corresponding author. E-mail: [email protected] (D. Tang), Tel.: +86 10

19

82323199.

20 21

Abstract

22

Oxygen availability is crucial for the evolution of eukaryotes in geological

23

history. However, the evolution of mid-Proterozoic oceanic–atmospheric redox

24

conditions remains heavily debated, e.g., [O2] <0.1–1% PAL vs. >4–8% PAL (present

25

atmospheric level). In order to further constrain the surface oxygen levels during

26

Mesoproterozoic, an investigation on the I/(Ca+Mg) ratios of deep-water (>100 m)

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early diagenetic carbonate concretions from the shale-predominated Member IV of

28

the Xiamaling Formation (~1.4 Ga) in three sections of the North China Platform was

29

conducted. The results show that more than half (36/47) of the I/(Ca+Mg) values

30

obtained from the concretions are higher than 0.5 μmol/mol (avg. 0.68), indicating

31

non-negligible iodate retained in the porewaters where the concretions were formed.

32

Compared with the iodine redox cycle and oxygen transportation pathway from

33

bottom seawater to surface sediments in modern oceans, the oxygen concentration of

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the bottom seawater was estimated to be higher than 16–22 μM, and the minimal

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atmospheric oxygen level should be higher than 6–9% PAL. This result differs from

36

the previous estimation that the atmospheric oxygen level is lower than 0.1–1% PAL,

37

but is consistent with the estimation of >4–8% PAL at ~1.4 Ga.

38 39 40

Keywords: Boring Billion; Oxygen concentration; Carbonate concretion; Early diagenesis; Eukaryote

41 42

1. Introduction

43

It remains highly debated whether the early diversification of eukaryotes and

44

appearance of animals in geological history are directly linked to the increase in

45

environmental oxygen (O2) levels or the result of genetic and/or developmental

46

innovations independent of any environmental control (e.g., Butterfield, 2009; Erwin

47

et al., 2011; Sperling et al., 2013; Mills et al., 2014; Planavsky et al., 2014). This

48

controversy largely derives from the fragmental and controversial constraints on

49

atmospheric and surface ocean oxygen levels during mid-Proterozoic (or “Boring

50

Billion”, ~1.8–0.8 Ga). These constraints indicate that the oxygen levels were low

51

(Stüeken, 2013; Planavsky et al., 2014; Cole et al., 2016; Koehler et al., 2017) or

52

relatively high (Sperling et al., 2014; Gilleaudeau et al., 2016; Mukherjee and large,

53

2016; Hardisty et al., 2017; Wang et al., 2017; Yang et al., 2017; Diamond et al.,

54

2018; Zhang et al., 2016, 2017, 2018).

55

Traditionally, the atmospheric oxygen levels of mid-Proterozoic were variably

56

constrained between ~1% and 40% PAL based on iron retention in paleosols (e.g.,

57

Zbinden et al., 1988; Lyons et al., 2014), and ocean anoxia and a steady state ocean

58

ventilation model (e.g., Canfield, 1998). However, well-preserved mid-Proterozoic

59

paleosols were characterized by loss of iron (Fe) and manganese (Mn), which were

60

similar to Archean paleosols but in contrast to Paleoproterozoic and Phanerozoic

61

records, suggesting a low (<1% PAL) pO2 (Zbinden et al., 1988; Mitchell and

62

Sheldon, 2009; Lyons et al., 2014).

63

Recently, some studies on Cr isotopes in marine ironstones and shales suggested

64

that the atmospheric oxygen level of mid-Proterozoic was <0.1–1% PAL (Planavsky

65

et al., 2014; Cole et al., 2016), lower than the minimum requirement for primitive

66

animals. In contrast, other studies suggested that there might exist sufficient oxygen

67

for animal respiration, at least during some intervals of mid-Proterozoic. For example,

68

Cr isotopes in a number of 1.33–1.08 Ga marine shales and 1.1–0.9 Ga marine

69

carbonates showed high fractionation, consistent with the atmospheric oxygen

70

concentrations of >1% PAL, higher than the minimum requirement of primitive

71

animals (Gilleaudeau et al., 2016; Canfield et al., 2018a). Moreover, the atmospheric

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oxygen level was estimated as high as >4% PAL, based on a simple ocean

73

water-column carbon-cycle model for the unit 3 (i.e., Member II) of the Xiamaling

74

Formation (~1.40–1.35 Ga) in the North China Platform (Zhang et al., 2016). More

75

recently, the atmospheric oxygen level at ~1.4 Ga was further estimated as >4% PAL

76

or even up to >8% PAL, according to the aerobic diagenesis identified in the unit 1

77

(i.e., Member IV) of the Xiamaling Formation based on trace metal systematics, iron

78

speciation and organic geochemistry (Zhang et al., 2017). In addition, some other

79

proxies, such as I/(Ca+Mg) ratios in shallow marine carbonates (Hardisty et al.,

80

2017), trace elements in pyrites (Mukherjee and Large, 2016), U and Mo isotopes in

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black shales of similar ages (Yang et al., 2017; Diamond et al., 2018) from different

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continents also point to a possibility of elevated atmospheric oxygen level at ~1.4 Ga.

83

Recent studies suggest that the Xiamaling Formation in North China can serve as

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an important window into the study of mid-Proterozoic redox conditions (e.g., Cole et

85

al., 2016; Planavsky et al., 2016; Zhang et al., 2016, 2017, 2019; Diamond et al.,

86

2018). However, viewpoints concerning the atmospheric and oceanic oxygen levels

87

during this stage are highly controversial; therefore, further constraints from multiple

88

geochemical proxies are urgently required. In recent years, iodine recorded in

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shallow-water carbonate was proposed as a robust redox proxy for paleoceanography

90

and has been widely used for many periods of geological records (e.g., Lu et al., 2010,

91

2016, 2018; Hardisty et al., 2014, 2017; Zhou et al., 2014, 2015, 2017; Edwards et al.,

92

2018; Shang et al., 2018). The Xiamaling Formation is dominated by dark shales with

93

some early diagenetic limestone concretions in its upper part (Liu et al., 2019). In this

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context, the carbonate concretions, which were mainly formed in the nitrate- to

95

manganese-reduction zones immediately below seawater/sediment interface (Liu et

96

al., 2019), may provide a precious opportunity for further tracing the redox conditions

97

of porewater and bottom seawaters based on iodine geochemistry, although this is not

98

straightforward as that of primary carbonates. In this study, I/(Ca+Mg) data from the

99

carbonate concretions and V/Al ratios of the host shales in the Member IV of the

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Xiamaling Formation in the Zhaojiashan, Jizhentun and Huangtugang sections, North

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China (Fig. 1) were measured, and the porewater and atmospheric oxygen levels were

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further constrained. Our new results, thus, can provided an insight into the evaluation

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of ocean-atmosphere oxygen levels at ~1.4 Ga.

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2. Geological setting and the occurrence of carbonate concretions

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The geological setting of the study area during the Mesoproterozoic has been

107

thoroughly introduced in Tang et al. (2017, 2018) and Liu et al. (2019), and only a

108

brief introduction will be provided herein. The studied carbonate concretions were all

109

collected from the Member IV of the Xiamaling Formation in the North China

110

Platform (Fig. 1). This formation lies disconformably between the underlying Tieling

111

Formation of the Jixian Group and the overlying Changlongshan Formation of the

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Qingbaikou Group (Fig. 2). The duration of the Xiamaling Formation has been

113

constrained

114

zircon/baddeleyite ages (Zhang et al., 2009, 2015). The studied interval is located in

115

middle part of the Member IV (Fig. 2) and can be approximately estimated as ~1.36

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Ga in age based on the two high-precision zircon/baddeleyite ages (Zhang et al.,

117

2015) and the average deposition rate. Paleomagnetic studies suggested that the North

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China Platform was most likely located in between 10°N and 30°N during the

119

deposition of the Xiamaling Formation (Evans and Mitchell, 2011; Zhang et al.,

120

2012). Organic matter preserved in the Xiamaling Formation is ranked as immature to

121

early thermal mature, with burial temperatures of ≤90 °C (Zhang et al. 2015; Tang et

122

al., 2017, 2018).

between

~1.40

Ga

and

~1.35

Ga

based

on

high-precision

123

In the study area, the Xiamaling Formation is dominated by dark siltstone and

124

black shales, and can be subdivided into four members (Member I to IV) in ascending

125

order, which constitutes a large transgressive-regressive cycle with its maximum

126

depositional water-depth in the Member III (Fig. 2; Zhang et al., 2016; Tang et al.,

127

2017, 2018). The lower part of Member IV consists of alternating black and greenish

128

shales, with a few of carbonate concretion interbeds (Fig. 3A and B). The upper part

129

of Member IV consists of a regressive succession from silty shale to yellowish

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siltstone in its upper portion, and is unconformably overlain by cross bedding-bearing

131

medium to coarsely grained quartz sandstone of the Changlongshan Formation. As

132

shown in Fig. 2, the Member IV targeted in this study is roughly equivalent to the unit

133

1 referred in Wang et al. (2017) and Zhang et al. (2017), which are stratigraphically

134

higher than the units 2 and 3 studied by Zhang et al. (2016) and Diamond et al.

135

(2018).

136

Notably in the middle part of the Member IV, there is a ~6 m-thick

137

limestone-bearing interval, which is unique for the Xiamaling Formation (Fig. 2), and

138

will be the focus of this study. The limestone-bearing interval is characterized by

139

densely or sparsely distributed concretions (Fig. 3B and C). Concretions are

140

commonly ellipsoidal in shape, 3–7 cm in height and 5–20 cm in width. Some of the

141

concretions are wrapped by bended laminations of surrounding shales (Fig. 3D and

142

E). The lateral transition zone from concretion carbonate to shale laminae is typically

143

wedge-shaped; this zone tapers from concretion margin to the more compacted host

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shale (Fig. 3E). Within the concretions, some well-preserved depositional laminations

145

can be observed, but a brown zone, likely deriving from late diagenetic alternation,

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commonly occurs at the margins as well (Fig. 3F).

147

The genesis of the carbonate concretions has been thoroughly studied by Liu et

148

al. (2019), and will be briefly introduced herein. These concretions have been

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considered of early diagenetic origin prior to the significant compaction of clay

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minerals, based on their macro- to microscopic fabrics, including deformed shale

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laminae surrounding the concretions (Fig. 3D and E), “cardhouse” structures of clay

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minerals (Fig. 5c and d in Liu et al. 2019) and calcite geodes (Fig. 6c and e in Liu et

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al. 2019) in the concretions. These concretions were not likely formed at the

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seawater/sediment interface directly in contact with bottom seawater, since vertically

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they were isolated and trapped by shale laminae and laterally the shale laminae were

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expanded in the concretion by diagenetic carbonate growth (Fig. 3E). The

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bicarbonates required for the concretion formation were mainly sourced from

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seawater by diffusion rather than produced by methanogenesis or anoxic oxidation of

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methane (AOM), because the carbon isotope compositions of the concretions (−1.7‰

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to +1.5‰; Table S1) are stable and close to or slightly lower than that of the

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contemporaneous seawater (Guo et al., 2013). Bacterial sulfate reduction (BSR) did

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not play a significant role in their formation either, because of the rare occurrence of

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authigenic pyrite grains in the concretions. Almost all the calcite grains in the

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concretions have low Mn–Fe contents in their nuclei but high Mn–Fe contents in their

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rims with average Mn/Fe ratio close to 3.3 (Table S2). The calcite shows positive Ce

166

anomalies (avg. 1.43) and low Y/Ho ratios (avg. 31) (Table S3). This evidence

167

suggests that Mn reduction is the dominant process responsible for the formation of

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calcite rims while nitrate reduction probably triggered the precipitation of calcite

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nuclei (cf. Tostevin et al., 2016; Tang et al., 2018).

170

A deep-water setting (below storm wave base) has been suggested for most of

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the Member IV, except for its upper part where hummocky and small-scale

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cross-bedding were locally observed, indicative of storm wave influences (Wang et

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al., 2017; Zhang et al., 2017). The studied interval (marked with red bar in Fig. 2)

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consists mainly of green to gray silty shale with some centimeter-scaled black shale

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interbeds and several layers of carbonate concretions (Fig. 3). No evidence of

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agitation from currents or waves has been recognized in this interval, likely indicating

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a depositional water-depth below storm-wave base (>100 m, Zhang et al., 2017).

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3. Materials and methods

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Samples analyzed in this study were collected from the Member IV of the

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Xiamaling Formation, along road cuts at the Zhaojiashan (40°28′4.76″N,

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115°24′5.74″E), Jizhentun (40°27′43.10″N, 115°16′29.48″E), and Huangtugang

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sections (40°27′8.62″N, 115°12′55.74″E), Hebei Province, North China (Fig. 1).

184

Collected limestone samples were cut into chips and only central parts of the samples

185

were used for geochemical analyses. Fresh sample chips were cleaned, dried, and then

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drilled for powders, avoiding weathered surfaces and recrystallized areas.

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For I/(Ca+Mg) analyses, ~5 mg of sample powders below 200 mesh were rinsed

188

4 times with 18.25 MΩ Milli-Q (MQ) water to remove absorbed clay minerals (Tang

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et al., 2017) and any potential soluble salts. After drying, the samples were grounded

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again into smaller and more homogenized powders in an agate mortar, and then

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weighted. Nitric acid (3%) was added for dissolution in 40 min and then centrifuged

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to obtain supernatant. The dilute nitric acid, short time of dissolution, and low content

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of clay minerals in this study could guarantee the negligible iodine from clay

194

minerals. For calcium (Ca) and magnesium (Mg) analyses, 0.2 mL supernatant was

195

used and then diluted to 1:50,000 with 3% HNO3 before analyses. Ca and Mg

196

concentrations were measured using a PerkinElmer NexION 300Q Inductively

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Coupled Plasma Mass Spectrometry (ICP-MS) at the National Research Center for

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Geoanalysis, Beijing. A certified reference material JDo-1 (dolostone) was measured

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after every nine samples and the analytical uncertainties monitored by JDo-1 were

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<3% for Mg and <2% for Ca. For iodine analysis, 1 mL supernatant was used, and 3%

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tertiary amine solution was added to the supernatant, and then diluted to 0.5% with

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MQ water to stabilize iodine (Lu et al., 2010; Hardisty et al., 2017). The iodine

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content was measured within 48 hours to avoid any iodine loss (Lu et al., 2010), using

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a MC-ICP-MS (Neptune Plus, Thermo Fisher Scientific, Germany) at the National

205

Research Center of Geoanalysis, Beijing. The sensitivity of iodine was tuned to

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~1,500 kcps for a 1 ppb standard in the MC-ICP-MS. The rinse solution used for each

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individual analysis contains 0.5% HNO3, 0.5% tertiary amine, and 50 μg/g Ca, and

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the typical rinse time is ~1 min. Analytical uncertainties for

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standard GSR 12 and duplicate samples are ≤ 6% (1σ), and the long term accuracy is

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checked by repeated analyses of the reference material GSR 12 (Shang et al., 2019).

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The detection limit of I/(Ca+Mg) is ~0.1 μmol/mol.

127I

monitored by the

212

The major and trace element concentrations of host shales were measured using a

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handheld energy dispersive XRF spectrometer (EDXRF) model Xsort with Rh anode

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from Spectro following the method described in Tang et al. (2017). Certified

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reference materials (GBW 07107 for shale) were measured after every five samples.

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Compared with the known values for shale standard, the accuracy (percent) of

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analyzed elements was generally <10%. Total organic carbon (TOC) and total sulfur

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(TS) contents of the shale samples were analyzed at the State Key Laboratory of

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Biogeology and Environmental Geology, China University of Geosciences (Beijing)

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following the method described in Liu et al. (2019).

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4. Results

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Altogether, I/(Ca+Mg) values of 47 limestone-concretion samples from the

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Zhaojiashan, Jizhentun, and Huangtugang sections were measured, and the results

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were presented in Table 1 and shown in Figs. 4 and 5. Values above 0.5 μmol/mol are

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commonly regarded as relatively high for the Proterozoic carbonates (Lu et al., 2017).

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In our study, 36 samples have their I/(Ca+Mg) values higher than 0.5 μmol/mol, and

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11 samples have values slightly lower than 0.5 μmol/mol but obviously higher than 0

229

μmol/mol (Table 1). No significant differences in the I/(Ca+Mg) values were

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observed among the samples from the three different sections, nor different parts of

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one concretion (Table 1, Fig. 3F). I/(Ca+Mg) values range from 0.32 to 1.59

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μmol/mol (mean: 0.71±0.30 μmol/mol, n = 22), from 0.29 to 1.20 μmol/mol (mean:

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0.63±0.21 μmol/mol, n = 16), and from 0.42 to 1.14 μmol/mol (mean: 0.73±0.24

234

μmol/mol, n = 9), in the Zhaojiashan, Jizhentun, and Huangtugang sections,

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respectively. In a concretion, the I/(Ca+Mg) values are stable in the vertical section,

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and from bottom to up the values are 0.71, 0.57, 0.78, 0.78, and 0.65 μmol/mol,

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respectively (Fig. 3F).

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The major and trace elements of 11 shale samples adjacent to the limestone from

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the Zhaojiashan section were analyzed using EDXRF, and the results were presented

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in Table 2 and shown in Fig. 5. Compared with post-Archaean Australian shales

241

(PAAS), vanadium (V) in this interval shows no enrichment, with V concentration

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(54–116 μg/g, avg. 78 μg/g) and V/Al (ppm/wt%) ratio (8–16, avg. 11) slightly

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lower than those of PAAS (140 μg/g and 14, respectively; McLennan, 2001). The

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lower V concentration differs significantly from that of the Member II (62–491 μg/g,

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avg. 353 μg/g) of the Xiamaling Formation that was explained as deposition in

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anoxic environment (Tang et al., 2017, 2018). The TOC contents of carbonate

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concretion hosting shales are commonly less than 0.2 wt% (0.12±0.06 wt%, n = 8);

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the TS contents of the shale are less than 0.06 wt% (Table 2).

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5. Discussion

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5.1. Oxygenated deep bottom seawater

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In recent years, I/(Ca+Mg) ratio has been proposed as a robust proxy for

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seawater redox conditions and widely used in the study of ancient carbonates,

254

including dolostone and limestone older than 2.5 Ga (e.g., Lu et al., 2010, 2016, 2018;

255

Hardisty et al., 2014, 2017; Zhou et al., 2014, 2015, 2017; Edwards et al., 2018; Wei

256

et al., 2019; Shang et al., 2019). In modern oceans, iodine composition mainly occurs

257

in two states: oxidized-state iodate (IO3-) and reduced-state iodide (I-). With the

258

decrease of oxygen concentrations in seawater (such as in oxygen minimum zone,

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OMZ), iodate would be reduced into iodide, displaying a positively correlation with

260

oxygen concentration (Wong and Brewer, 1977; Emerson et al., 1979; Wong et al.,

261

1985; Luther III and Campbell, 1991). Laboratory experiments demonstrated that only

262

IO3- could be incorporated into the lattices of carbonate minerals with a fixed

263

distribution coefficient, but I- would be excluded (Lu et al., 2010; Feng and Redfern,

264

2018). Thus, the iodine concentration in carbonates [expressed as I/(Ca+Mg)] can be

265

used as a robust proxy for redox conditions of seawater in which the carbonate was

266

deposited (e.g., Lu et al., 2010, 2016, 2018; Hardisty et al., 2014, 2017; Zhou et al.,

267

2014, 2015).

268

Iodate could not be retained but would be effectively reduced to iodide in

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anoxic seawater (Wong and Brewer, 1977; Emerson et al., 1979; Wong et al., 1985;

270

Luther III and Campbell, 1991) or porewater (Szecsody et al., 2017), resulting in a

271

zero I/(Ca+Mg) value in carbonates or carbonate concretions. For example, the

272

reduction rate of iodate in anoxic porewater could be up to ~13 pM/h, with a half-life

273

less than 20 hours (Szecsody et al., 2017). Thus, the non-zero I/(Ca+Mg) values (avg.

274

0.68 μmol/mol) indicates that there was non-negligible amount of iodate retained in

275

the porewater during the diagenetic formation of these concretions (cf. Hardisty et al.,

276

2014, 2017). In addition, the retaining of iodate in the porewater should indicate the

277

presence of molecular oxygen rather than retention of iodate in anoxic porewater. In a

278

typical diagenetic condition, oxygen would be sharply decreased from bottom

279

seawater to the surface sediments, and fully consumed in the Mn- to Fe-reduction

280

zone several mm to cm below the seawater/sediment interface, mainly owing to the

281

aerobic respiration of organic matter (Canfield and Thamdrup, 2009). In this study,

282

the stable I/(Ca+Mg) values revealed in a concretion from its bottom to up (Fig. 3F)

283

requires the vertical growth model of the concretions, keeping the carbonate

284

precipitation front in the nitrate- to Mn- reduction zone during the formation of

285

concretions (Fig. 8 in Liu et al., 2019). Thus, the I/(Ca+Mg) values preserved in these

286

Xiamaling concretions may have recorded the signals of porewater in the nitrate- to

287

Mn- reduction zone in a shallow depth below the sediment surface. Since both the

288

iodate and oxygen in the porewater were mainly diffused from bottom seawater,

289

therefore, the bottom seawater should also contain certain amount of iodate and

290

oxygen (e.g., Jørgensen and Revsbech, 1985; Gundersen and Jørgensen, 1990; Glud et

291

al., 2007; Glud, 2008).

292

The moderately oxygenated bottom seawater during deposition of the Member

293

IV of the Xiamaling Formation is also supported by hydrogen index and iron

294

speciation data reported previously from this interval (Zhang et al., 2017), and by the

295

low V/Al ratios in the surrounding shale revealed in this study (Fig. 5). Under anoxic

296

environments, high V enrichment would be expected due to the reduction of vanadyl

297

species (Emerson and Huested, 1991; Tribovillard et al., 2006; Piper and Calvert,

298

2009; Zhang et al., 2016). Under oxic ([O2] >63 μM) to hypoxic ([O2] <63 μM but >0

299

μM) environments (cf. Middelburg and Levin, 2009), vanadate is transported to

300

sediments as the vanadate ion [H2V(VI)O4−] adsorbed onto Mn oxides. The vanadate

301

ion is released as the Mn oxides are reduced to Mn2+ (Emerson and Huested, 1991),

302

and, where oxygen is limiting, Mn oxides do not readily reform at the sediment

303

surface (Emerson and Huested, 1991; Morford and Emerson, 1999; Morford et al.,

304

2005), allowing vanadate to escape to the overlying water and resulting low V/Al in

305

sediments (Zhang et al., 2016). The low V/Al ratios in the black shales from the

306

Member III of the Xiamaling Formation were ascribed to highly restricted conditions

307

as well, mainly based on the high TOC/TS ratios (Diamond et al., 2018), while

308

significantly high V/Al ratios were found in the shale and siltstone from the lower part

309

of the Xiamaling Member II (Tang et al., 2017, 2018). Therefore, we suggested that in

310

the studied interval, the lower V/Al ratios (slightly lower than that of PAAS) are

311

likely related to Mn reduction in porewater and could indicate a hypoxic–oxic bottom

312

seawater conditions, rather than caused by basin restriction (Zhang et al., 2019).

313 314

5.2. Estimation of oxygen concentration

315

The similarity of iodine reservoir and redox cycle in Proterozoic and modern

316

oceans (Hardisty et al., 2017) permits an estimation of oxygen concentrations in

317

Proterozoic seawater based on the relationship between I/(Ca+Mg) ratios and oxygen

318

concentrations in modern seawaters. In modern oceans, the iodate and oxygen

319

concentrations are commonly positively correlated (e.g., Rue et al., 1997; Lu et al.,

320

2010), although iodate concentrations may be lowered by microbial absorption (Zhou

321

et al., 2014) or by the existence of neighboring seawaters with low oxygen

322

concentration (Lu et al., 2016). Iodate as a micronutrient is used by marine organisms,

323

and the higher primary productivity may cause more iodate loss in surface ocean

324

waters through microbial absorption (Lu et al., 2014 and references cited therein). The

325

relationship between oxygen and iodate concentrations in porewater is rarely reported

326

in sediments from modern oceans, but no factors that could change this

327

disproportionately have been proposed so far. We tentatively deduce that the

328

positively correlation between oxygen and iodate concentrations shown in seawater

329

column could also be roughly retained in porewater.

330

Previous studies have confirmed that the minimum [O2] requirement was at least

331

1–3 μM for marine IO3- accumulation and for the presence of carbonate-bound iodine

332

carbonates (Rue et al., 1997; Hardisty et al., 2014, 2017). Consequently, I/(Ca+Mg)

333

>0 μmol/mol was suggested as an indicator for seawater [O2] >1–3 μM (Hardisty et

334

al., 2014, 2017). In our study, all the analyzed samples from the Member IV of the

335

Xiamaling Formation show I/(Ca+Mg) values obviously higher than 0 μmol/mol (Fig.

336

5), likely indicating a condition of porewater [O2] at least higher than 1–3 μM

337

(Hardisty et al., 2014, 2017). The studies on the correlation between oxygen and

338

iodate concentrations in modern oxygen minimal zones are limited but important for

339

further constraining the porewater oxygen concentration. We have plotted our data on

340

the crossplot between oxygen and iodate concentrations in surface waters from

341

Atlantic (Lu et al., 2010), where the relationship between iodate and oxygen

342

concentrations has been well studied. The results show that our I/(Ca+Mg) data likely

343

reflect the porewater oxygen concentrations ranging from 17–70 μM (avg. 38 μM)

344

(Fig. 6A). It has been proposed that the seawater [IO3-] is mainly controlled by in situ

345

[O2], and is also strongly influenced by the redox condition of neighboring seawater

346

(Lu et al., 2016; Hardisty et al., 2017). The sharp water column gradient in [IO3-] is

347

mainly caused by fast change in oxygen levels and influenced by the variation in

348

primary productivity (Zhou et al., 2014). In Fig. 6B, although our data indicate the in

349

situ [O2] only higher than 1 μM, the sharp gradient in [IO3-] likely indicates an

350

obviously higher oxygen concentration in the slightly shallower seawater. It can be

351

found that the [O2] and [IO3-] show obvious covariation in Fig. 6B, but the decrease

352

trend of [IO3-] is at a position ~30 m deeper than that of [O2] in shallow seawater.

353

Thus, our I/(Ca+Mg) data likely indicate the existence of overlying oxygenated

354

seawater with [O2] ≥30 μM (Fig. 6B). The reason for the decreasing trend of [IO3-] at

355

~30 m deeper than that of [O2] is possibly caused by the reduced primary productivity

356

or the fast diffusion of IO3- from shallower seawater than that of O2. Considering the

357

estimation uncertainty caused by limited reference data of modern seawater (Fig. 6A)

358

and unsynchronized variation in [O2] and [IO3-], we only used the lower limit of the

359

oxygen concentration to make a conservative estimation. As a result, the porewater

360

[O2] is conservatively estimated higher than 10 μM, although more comparisons with

361

modern seawaters are required to confirm this estimation.

362

The oxygen concentration could also be further estimated through a comparison

363

with typical I/(Ca+Mg) values of Precambrian carbonates. A statistical analysis shows

364

that, except for the samples from the intervals with elevated oxygen contents, more

365

than 95% of data (n = 466) in previously reported Proterozoic samples have

366

I/(Ca+Mg) values of <0.5 μmol/mol (Hardisty et al., 2017; Lu et al., 2017; Shang et

367

al., 2019). Thus, a value of 0.5 μmol/mol has been suggested as the baseline for

368

I/(Ca+Mg) concentrations in Precambrian carbonates (Lu et al., 2017). A numerical

369

modeling shows that when atmospheric oxygen concentrations were <2.5% PAL, the

370

oxygen concentrations of surface seawater were fundamentally controlled by the

371

primary production and a disharmonious oxygenation would exist between the ocean

372

and atmosphere (Reinhard et al., 2016). In addition, simulation study also indicates

373

that the maximum oxygen concentration in the surface ocean caused by the primary

374

production should be 10 μM (e.g., Kasting, 1991; Olson et al., 2013; Reinhard et al.,

375

2016). Thus, it is likely that the oxygen concentrations caused by primary production

376

in the surface ocean during Proterozoic (Reinhard et al., 2016) could only result in

377

carbonate I/(Ca+Mg) ratios less than 0.5 μmol/mol. In our study, more than half of the

378

carbonate concretions (36/47) show I/(Ca+Mg) values higher than 0.5 μmol/mol,

379

likely indicating that the porewater oxygen concentration was higher than 10 μM, the

380

highest oxygen concentration that could be produced by Proterozoic primary

381

production. A small portion of the carbonate concretions (11/47) show I/(Ca+Mg)

382

values slightly lower than 0.5 μmol/mol, which is likely caused by stronger diagenetic

383

reduction of iodate in deeper depth below seawater/sediment interface. Thus, this

384

estimation is in good consistence with the estimation based on the relationship

385

between iodate and oxygen concentrations in modern oceans.

386

Generally, the oxygen concentration of bottom seawater is obviously higher than

387

that in the sediment porewater below seawater/sediment interface, because molecular

388

diffusion of oxygen and aerobic degradation of organic matter would sharply decrease

389

the oxygen concentration. It is commonly accepted that in continental margin

390

sediments, oxygen is supplied to the sediment surface and, subsequently, into the

391

sediments via molecular diffusion, through a 400–800 μm-thick viscous diffusive

392

boundary layer immediately above the seawater-sediment interface (e.g., Jørgensen

393

and Revsbech, 1985; Gundersen and Jørgensen, 1990; Boudreau, 2001; Glud et al.,

394

2007; Glud, 2008). Even if the [O2] is close to 0 μM (but >0 μM) in the porewater

395

below the viscous boundary layer, a minimum [O2] of bottom seawater between 6 μM

396

(with a 400 μm boundary layer) and 11 μM (with an 800 μm boundary layer) would

397

be required to maintain the oxygen concentration of >0 μM in the porewater (cf.

398

Zhang et al., 2017). Therefore, the minimum bottom-water oxygen concentration

399

should be >16–22 μM at that time in order to keep the porewater oxygen

400

concentration higher than 10 μM (Fig. 6C and D), and the atmospheric oxygen level

401

should be higher than 6–9% PAL (modern surface seawater [O2] = ~250 μM, Garcia

402

et al., 2013). It should also be pointed out that this calculation on seawater oxygen

403

concentration does not account for any reduction in bottom-water and porewater

404

oxygen concentration that might have been caused by the respiration of settling

405

organic particles through the oxic water column and in porewater. Accommodating

406

this oxygen loss would further increase the estimated values of the atmospheric

407

oxygen level.

408

Increasing evidence from North China and other continents seems to indicate

409

that a relatively high oxygen level in surface ocean and atmosphere at ~1.4 Ga is

410

likely a global phenomenon, although the redox conditions are dynamic and shallow

411

anoxic seawaters were locally identified (e.g., Tang et al., 2017, 2018; Canfield et al.,

412

2018b). In North China, the high redox-sensitive metal concentrations from the

413

~1.40–1.35 Ga Xiamaling Formation (Member II or unit 3 in Fig. 2) and a

414

carbon-oxygen cycle model were used to argue that the atmospheric O2 levels were

415

>4% PAL (Zhang et al., 2016). Based on the aerobic diagenesis identified in the

416

Xiamaling Formation (Member IV or unit 1 in Fig. 2), the trace metal systematics,

417

iron speciation and organic-geochemistry, the atmospheric oxygen level of the time

418

was estimated to be >4% PAL (Zhang et al., 2016) or even higher (>8% PAL) (Zhang

419

et al., 2017). The high Mo concentration (51 g/g) and a δ98Mo value of +1.7‰

420

obtained in the black shale of the Xiamaling Formation (Member III or unit 2 in Fig.

421

2) also point to the possibility that the Earth’s surface environments were transiently

422

more oxygenated at ~1.4 Ga compared to preceding times (Diamond et al., 2018). In

423

addition, the high I/(Ca+Mg) ratios (up to 2.7 μmol/mol) obtained in carbonates from

424

the ~1.44–1.40 Ga Tieling Formation also suggest relatively high atmospheric O2

425

levels (Hardisty et al., 2017). In North Australia, a set of geochemical data from the

426

upper Velkerri Formation (the McArthur Basin) have been used to argue an episode

427

of increased ocean oxygenation at ~1.4 Ga, including elevated redox-sensitive metal

428

enrichments in black shales (Cox et al., 2016; Mukherjee and Large, 2016). An

429

uranium isotope mass-balance modelling suggests that <25% of the seafloor was

430

anoxic at 1.36 Ga (Yang et al., 2017). In West Siberia, Fe speciation, organic

431

geochemistry, pyrite S isotope compositions, redox-sensitive metal concentrations,

432

and clearly identified eukaryotic microfossils from the basinal sedimentary rocks in

433

the Arlan Member (Kaltasy Formation) all seem to support that O2-bearing deep

434

marine waters were present at places at ~1.4 Ga (Sperling et al., 2014). These

435

continents were far apart during Mesoproterozoic according to relevant paleomagnetic

436

reconstruction (e.g., Zhang et al., 2012), thus, it is likely that ~1.4 Ga was a stage with

437

relatively high oxygen level both in surface ocean and atmosphere globally, though its

438

precise timing and causes need to be further constrained and investigated in details.

439 440

6. Conclusions

441

The middle part of the Member IV of the Xiamaling Formation contains valuable

442

carbonate concretions that may have recorded the signals of redox conditions in the

443

early diagenetic porewater and could be used to reflect the redox conditions of

444

overlying seawater indirectly. I/(Ca+Mg) values from these concretions are obviously

445

higher than 0 μmol/mol (avg. 0.68 μmol/mol), likely indicating the oxygen

446

concentration in porewater higher than 10 μM during their formation. Furthermore,

447

the minimum bottom-seawater oxygen concentration estimated from the oxygen

448

molecular diffusion model is higher than 16–22 μM (6–9% PAL). This relatively high

449

oxygen concentration is also supported by lower V/Al ratios in surrounding shales.

450

The relatively high oxygen concentration recorded in the ~1.4 Ga Xiamaling

451

Formation, in combination with various redox-based geochemical evidence identified

452

in other continents, likely indicates that ~1.4 Ga was a stage witnessing high

453

atmospheric oxygen concentration globally.

454 455

Acknowledgments The study was supported by the National Natural Science

456

Foundation of China (Nos. 41930320 and 41972028), the Fundamental Research

457

Funds for the Central Universities (No. 2652018005), and the Key Research Program

458

of the Institute of Geology & Geophysics, CAS (No. IGGCAS-201905). Thanks are

459

given to Haoming Wei, Zhipeng Wang, and Tong Wu for their assistance in field

460

work.

461 462

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734 735

Figure and Table Captions

736

Fig. 1. (A) Major tectonic subdivisions of China; the box showing the area illustrated

737

in panel B. (B) Simplified map showing the location of the studied area. (C)

738

Simplified geological map of the study area (modified after the 1:200,000 Geological

739

Map of China, The China Geological Survey, 2013).

740 741

Fig. 2. Correlation of different subdivision schemes for the Xiamaling Formation in

742

Zhaojiashan, Jizhentun (Wang et al., 2017), and Huangtugang sections. The interval

743

studied in this paper is in the Member IV or unit 1.

744 745

Fig. 3. Lithology and facies features in the Member IV of the studied sections. (A)

746

Alternation of black and greenish shales in the Member IV, Zhaojiashan section. (B)

747

Studied interval in middle part of the Member IV of the Xiamaling Formation,

748

showing carbonate concretions (arrows) in the black–green silty shales, Zhaojiashan

749

section. (C) Overview of the carbonate concretion layers in green shales, Jizhentun

750

section. (D) A carbonate concretion wrapped by green shale, Huangtugang section.

751

(E) A concretion with lateral shale laminae (dash lines and arrows), and the laminae

752

thickness in concretions is expanded by an order of magnitude relative to the laminae

753

in adjacent shales, Jizhentun section. (F) Polished slab of a concretion, showing

754

depositional lamination and late diagenetic brown rim preserved in the concretion,

755

Jizhentun section. The I/(Ca+Mg) values (μmol/mol) of the white circle areas are: a =

756

0.65, b = 0.78, c = 0.78, d = 0.57, e = 0.71, respectively.

757 758

Fig. 4. Secular variations in I/(Ca+Mg) through time (modified from Hardisty et al.,

759

2017; Lu et al., 2017, 2018). Vertical blue bars mark the intervals where the

760

I/(Ca+Mg) data are consistent with possible oxygenation events (Kendall et al., 2009;

761

Planavsky et al., 2014; Gilleaudeau et al., 2016; Zhang et al., 2016) relative to the

762

long term atmospheric p[O2] curve (Lyons et al., 2014; Planavsky et al., 2014;

763

Sperling et al., 2015) and biological production of O2 (Cox et al., 2018). The vertical

764

orange bars mark the multi-cellularization of eukaryotes in the Mesoproterozoic (e.g.,

765

Zhu et al., 2016) and the diversification of eukaryotes in the Neoproterozoic (Mills et

766

al., 2014), respectively. The grey dashed line at 0.5 μmol/mol marks the Precambrian

767

I/(Ca+Mg) baseline and the grey dashed line at 2.6 μmol/mol represents the threshold

768

of [O2] >20–70 μM.

769 770

Fig. 5. Stratigraphic distributions of I/(Ca+Mg) and V/Al in the study interval of the

771

Xiamaling Formation at Zhaojiashan, Jizhentun and Huangtugang sections, Hebei

772

Province (please refer to Fig. 2 for legends). The red dash line marks the baseline for

773

Precambrian carbonates (Lu et al., 2017), while the black dash line for the V/Al value

774

of PAAS; the red circles represent carbonate concretions, while the black circles

775

represent shales.

776 777

Fig. 6. Estimation the oxygen concentrations of porewater and bottom seawater. (A)

778

Estimation of porewater [O2] based on the relationship between [IO3-] and [O2] in the

779

Atlantic Ocean (Lu et al., 2010). The green area shows the range of [IO3-] and

780

estimated [O2], and the red dash line is the mean value of [IO3-] and estimated [O2].

781

The [IO3-] = [I/(Ca+Mg)]/9.74, [O2] = 561 × [IO3-] (cf. Lu et al., 2010). (B)

782

Estimation of porewater [O2] based on the relationship between [IO3-] and [O2] in the

783

Peruvian oxygen minimal zone (Rue et al., 1997). The green area shows the range of

784

[IO3-] and estimated in situ [O2], while the blue area shows the range of overlying

785

seawater [O2]. (C) and (D) Schematic diagrams showing the viscous sublayer, the

786

diffusive boundary layer (DBL) and the effective DBL as derived from the O2

787

concentration profile in modern ocean (Boudreau, 2001; Glud, 2008). In continental

788

margin sediments, oxygen is commonly supplied to the sediment surface and,

789

subsequently, into the sediments via molecular diffusion, through a 400–800 μm-thick

790

viscous diffusive boundary layer immediately above the seawater-sediment interface

791

(e.g., Jørgensen and Revsbech, 1985; Gundersen and Jørgensen, 1990; Boudreau,

792

2001; Glud et al., 2007; Glud, 2008). This process would significantly decrease the

793

oxygen concentration in porewater than that of the bottom seawater above the viscous

794

DBL. When the thickness of DBL is 0.4 mm or 0.8 mm and the porewater [O2] is 10

795

μM, the bottom seawater [O2] could be estimated as 16 μM or 22 μM, respectively

796

(cf., Glud, 2008; Zhang et al., 2017).

797 798

Table 1. I/(Ca+Mg) values of the carbonate concretions from the Member IV of the

799

Xiamaling Formation, North China.

800

801

Table 2. Major and trace element concentrations of the shales from the Member IV of

802

the Xiamaling Formation, Zhaojiashan section.

803 804 805

Table 1 I/(Ca+Mg) values of the carbonate concretions from the Member IV of the Xiamaling Formation, North China

806 Height

I

Ca

I/(Ca+Mg)

(wt%)

(μmol/mol)

0.52

30.89

0.35

0.30

0.36

32.48

0.48

1.23

0.42

0.39

37.20

0.35

Zhaojiashan

1.23

0.36

0.38

34.42

0.32

1710ZJS-24-3-1c

Zhaojiashan

1.23

1.18

0.31

31.62

1.16

No. 006

1710ZJS-24-7-1a

Zhaojiashan

2.32

0.54

0.34

32.03

0.52

No. 007

1710ZJS-24-7-1b

Zhaojiashan

2.32

0.95

0.50

30.36

0.96

No. 008

1710ZJS-24-9-1a

Zhaojiashan

2.45

0.83

0.47

36.29

0.71

Serial No.

Sample ID

Section

No. 001

1710ZJS-24-1-1a

Zhaojiashan

0.09

0.35

No. 002

1710ZJS-24-1-1b

Zhaojiashan

0.09

No. 003

1710ZJS-24-3-1a

Zhaojiashan

No. 004

1710ZJS-24-3-1b

No. 005

(m)

Mg

(μg/g) (wt%)

No. 009

1710ZJS-24-9-1b

Zhaojiashan

2.45

0.66

0.39

39.26

0.52

No. 010

1710ZJS-24-11-1

Zhaojiashan

2.55

0.67

0.44

39.21

0.53

No. 011

1710ZJS-24-13-1

Zhaojiashan

2.8

0.77

0.50

33.00

0.72

No. 012

1710ZJS-24-15-1

Zhaojiashan

3.12

1.11

0.34

38.64

0.90

No. 013

1710ZJS-24-15-2

Zhaojiashan

3.13

1.76

0.43

34.23

1.57

No. 014

1710ZJS-24-17-1

Zhaojiashan

3.25

0.87

0.48

35.34

0.75

No. 015

1710ZJS-24-17-2a Zhaojiashan

3.3

1.04

0.39

38.46

0.84

No. 016

1710ZJS-24-17-2b Zhaojiashan

3.3

0.91

0.66

34.71

0.80

No. 017

1710ZJS-24-19-1

Zhaojiashan

3.65

1.06

0.62

31.16

1.04

No. 018

1710ZJS-24-21-1a Zhaojiashan

3.82

0.48

0.40

27.78

0.53

No. 019

1710ZJS-24-21-1b Zhaojiashan

3.82

0.50

0.48

41.65

0.37

No. 020

1710ZJS-24-23-1

Zhaojiashan

4.38

0.55

0.22

32.92

0.52

No. 021

1710ZJS-24-25-1

Zhaojiashan

4.53

0.65

0.49

31.21

0.64

No. 022

1710ZJS-24-J

Zhaojiashan

5.53

1.16

0.66

37.69

0.94

No. 023

1809JZT-18-0.5

Jizhentun

0.5

0.78

0.27

33.11

0.73

No. 024

1809JZT-18-1.1

Jizhentun

1.1

1.35

0.32

34.91

1.20

No. 025

1809JZT-18-1.8-a

Jizhentun

1.8

0.79

0.51

37.29

0.65

No. 026

1809JZT-18-1.8-b

Jizhentun

1.8

0.86

0.60

33.92

0.78

No. 027

1809JZT-18-1.8-c

Jizhentun

1.8

0.89

0.58

34.90

0.78

No. 028

1809JZT-18-1.8-d

Jizhentun

1.8

0.64

0.50

34.58

0.57

No. 029

1809JZT-18-1.8-e

Jizhentun

1.8

0.81

0.56

34.89

0.71

No. 030

1809JZT-18-1.9

Jizhentun

1.9

0.58

0.39

31.74

0.56

No. 031 1809JZT-18-2.2(2)-c Jizhentun

2.2

0.57

1.44

29.04

0.78

No. 032 1809JZT-18-2.2(2)-k Jizhentun

2.2

0.43

0.35

34.73

0.39

No. 033

1809JZT18-2.9-a

Jizhentun

2.9

0.44

0.42

32.39

0.42

No. 034

1809JZT-18-2.9-c

Jizhentun

2.9

0.92

1.53

31.91

0.74

No. 035

1809JZT-18-3.0

Jizhentun

3

0.34

0.30

35.57

0.29

No. 036

1809JZT-18-3.6

Jizhentun

3.6

0.68

0.41

37.30

0.57

No. 037 1809JZT-18-4.2(2)-c Jizhentun

4.2

0.64

0.33

35.53

0.55

No. 038

1809JZT-18-J

Jizhentun

4.3

0.38

0.22

34.12

0.35

No. 039

1809HTG-01

Huangtugang

0.4

0.88

0.48

36.78

0.74

No. 040

1809HTG-02

Huangtugang

0.5

1.41

0.44

39.73

1.10

No. 041

1809HTG-03

Huangtugang

0.9

0.53

0.26

27.99

0.59

No. 042

1809HTG-04

Huangtugang

1.1

0.68

0.54

31.39

0.67

No. 043

1809HTG-05-c

Huangtugang

1.4

0.92

0.59

35.04

0.80

No. 044

1809HTG-06

Huangtugang

1.8

1.14

0.53

30.57

1.14

No. 045

1809HTG-07-c

Huangtugang

2.2

0.60

0.45

41.59

0.45

No. 046

1809HTG-08

Huangtugang

2.9

0.75

0.67

34.12

0.67

No. 047

1809HTG-09

Huangtugang

3.7

0.53

0.63

38.71

0.42

807 808 809

Table 2 Major and trace element concentrations of the shales from the Member IV of the Xiamaling Formation, Zhaojiashan section

Height Sample ID

MgO

Al2O3

SiO2

P2O5

K2O

CaO

Ti

MnO

Fe2O3

V

Lithology

Fe/Al (m)

(wt%)

(wt%)

(wt%)

(wt%)

(wt%)

(wt%)

(wt%)

(wt%)

(wt%)

(μg/g)

(wt%)

C

0.6

9.6

12.4

60.3

0.2

3.4

3.6

0.4

0.1

5.2

54

0.6

8.2

0.06

0.01

5

ZJS-24-2-2

BS

0.8

8.4

12.3

64.9

0.2

3.4

2.6

0.3

0.0

5.1

83

0.5

12.7

0.96

0.01

74

ZJS-24-4-1

BS

1.5

10.7

13.6

61.3

0.2

3.7

1.8

0.4

0.0

5.5

116

0.5

16.1

1.14

0.02

52

ZJS-24-4-2

GS

1.7

8.1

13.6

61.9

0.2

3.6

2.3

0.4

0.0

5.3

71

0.5

9.8

0.18

0.03

6

ZJS-24-6-1

GS

2.0

10.1

13.8

62.3

0.2

3.8

1.7

0.4

0.0

6.0

76

0.6

10.4

0.09

0.03

3

ZJS-24-6-2

BS

2.2

12.5

13.6

62.9

0.2

3.5

2.3

0.4

0.1

6.2

94

0.6

13.1

1.12

0.03

33

ZJS-24-14-1

GS

3.1

8.5

13.8

57.6

0.2

3.9

1.4

0.4

0.0

6.7

72

0.6

9.9

0.09

0.04

2

ZJS-24-18-1

GS

3.5

5.2

12.5

56.9

0.1

3.4

1.9

0.4

0.3

10.3

71

1.1

10.8

0.12

0.05

2

ZJS-24-20-1

GS

3.8

8.5

15.8

68.4

0.2

4.0

1.4

0.4

0.1

5.6

82

0.5

9.8

0.15

0.06

2

ZJS-24-22-1

GS

4.1

8.7

11.8

55.9

0.2

3.4

3.2

0.4

0.2

7.5

78

0.8

12.5

0.21

0.04

5

ZJS-24-26-1

GS

5.0

9.2

12.5

59.0

0.2

3.3

1.4

0.4

0.1

9.4

56

1.0

8.6

0.06

0.04

1

811

812

Highlights

813

 Mesoproterozoic Xiamaling carbonate concretions were formed in early diagenesis.

815

 Non-zero I/(Ca+Mg) values of the concretions suggest oxygenated porewater.

816

 Atmospheric oxygen level was estimated >6–9% of the present atmospheric

818

(wt%)

GS

810

817

TS

ZJS-24-2-1

Note: GS = green shale, BS = black shale.

814

TOC V/Al

level.