Earth and Planetary Science - Cugb€¦ · Geology, China University of Geosciences (CUGB). electron imaging detector (SE2) was to charac-topographic features, and an AsB was used

Earth and Planetary Science Letters 527 (2019) 115797

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Earth and Planetary Science Letters

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A pulse of oxygen increase in the early Mesoproterozoic ocean at ca. 1.57–1.56 Ga

Mohan Shang a,b, Dongjie Tang a,c,∗, Xiaoying Shi a,b,∗, Limin Zhou d, Xiqiang Zhou e, Huyue Song f, Ganqing Jiang g

a State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (Beijing), Beijing 100083, Chinab School of Earth Sciences and Resources, China University of Geosciences (Beijing), Beijing 100083, Chinac Institute of Earth Sciences, China University of Geosciences (Beijing), Beijing 100083, Chinad National Research Center of Geoanalysis, Beijing 100037, Chinae Key Lab of Petroleum Resources Research, Institute of Geology and Geophysics, Chinese Academy of Sciences, Beijing 100029, Chinaf State Key Laboratory of Biogeology and Environmental Geology, China University of Geosciences (Wuhan), Wuhan 430074, Chinag Department of Geoscience, University of Nevada, Las Vegas, NV 89154-4010, USA

a r t i c l e i n f o a b s t r a c t

Article history:Received 2 January 2019Received in revised form 25 July 2019Accepted 29 August 2019Available online xxxxEditor: I. Halevy

Keywords:boring billionoxygenation episodeGaoyuzhuang FormationI/(Ca+Mg)North China Platform

The relationship between oxygen and evolution of early eukaryotes including algae and primitive animals in geological history has been debated, partly due to the varying estimates of oxygen levels in the mid-Proterozoic (ca. 1.8–0.8 Ga) ocean and atmosphere. The upper part of the Gaoyuzhuang Formation (ca. 1.60–1.54 Ga) in North China hosts decimeter-scale multicellular eukaryotic fossils and is documented with a decrease in cerium anomaly indicative of ocean oxygenation. However, the atmospheric oxygen level across this interval and its subsequent oxidation state require further investigation using additional redox proxies. Here we report I/(Ca+Mg) ratios, carbonate/organic carbon isotopes (δ13Ccarb and δ13Corg), and phosphorous (P) contents across the ca. 1.57–1.56 Ga fossil-bearing interval in the North China Platform. High I/(Ca+Mg) ratios (≥2.6 μmol/mol; up to 3.8 μmol/mol) from shallow-water carbonates of the Gaoyuzhuang Formation suggest an episode of significant oxygen increase up to ≥4% PAL (present atmospheric level). The I/(Ca+Mg) ratios return back to ≤0.5 μmol/mol shortly after the peak values without evidence for increasing water depth or diagenetic alteration, implying a short-lived oxidation event. The increase of I/(Ca+Mg) ratios is associated with a −3.5� negative δ13Ccarb and δ13Corg anomaly and an increase in P/Al ratios that are best explained by oxidation of dissolved organic carbon (DOC) in the ocean. Oxygen consumption through oxidation of DOC may have quickly lowered marine and atmospheric O2 levels to the early mid-Proterozoic (1.8–1.4 Ga) background oxygen concentration of ≤0.1–1% PAL. Short-lived oxidation events in an overall anoxic mid-Proterozoic ocean and atmosphere best explain the existing geochemical data and evolutionary stasis of eukaryotes during the “Boring Billion”.

© 2019 Elsevier B.V. All rights reserved.

1. Introduction

The mid-Proterozoic (ca. 1.8–0.8 Ga) witnessed the emergence but slow diversification of eukaryotes (Butterfield, 2015; Knoll, 2014). Multi-proxy geochemical studies suggested that the mid-Proterozoic ocean was mostly ferruginous with euxinic wedges on shelf margins (e.g., Luo et al., 2014; Planavsky et al., 2011;Poulton and Canfield, 2011) and perhaps episodes of increased oxygen or spatially-limited oxygenation in some sedimentary

* Corresponding authors at: State Key Laboratory of Biogeology and Environmen-tal Geology, China University of Geosciences (Beijing), Beijing 100083, China.

E-mail addresses: [email protected] (D. Tang), [email protected] (X. Shi).

https://doi.org/10.1016/j.epsl.2019.1157970012-821X/© 2019 Elsevier B.V. All rights reserved.

basins at ca. 1.85 Ga (Planavsky et al., 2018a), ca. 1.57–1.54 Ga(Tang et al., 2016; Zhang et al., 2018), ca. 1.4 Ga (Cox et al., 2016; Hardisty et al., 2017; Mukherjee and Large, 2016;Sperling et al., 2014; Yang et al., 2017; Zhang et al., 2016), and ca. 1.1 Ga (Gilleaudeau et al., 2016). The overall weakly oxy-genated condition was possibly caused by low primary produc-tion and low organic carbon burial flux (Crockford et al., 2018;Ozaki et al., 2019). Estimates of mid-Proterozoic atmospheric O2

levels also vary significantly. Earlier assessments on the basis of iron retention in paleosols suggested atmospheric O2 con-centrations of ≥1–3% PAL [present atmospheric levels] (Rye and Holland, 1998; Zbinden et al., 1988; but see Planavsky et al.,
2018b for a different view). Paleoenvironment and diagenetic mod-

2 M. Shang et al. / Earth and Planetary Science Letters 527 (2019) 115797

els constructed from trace element concentrations, organic car-bon contents and biomarkers of the Mesoproterozoic Xiamaling Formation (ca. 1.40–1.35 Ga) in North China also suggested at-mospheric O2 levels of ≥4–8% PAL (Zhang et al., 2016, 2017). However, the lack of observable chromium (Cr) isotope fraction-ation in mid-Proterozoic marine ironstones and shales, and Ce anomalies in carbonates implies a much lower O2 level that was likely <0.1–1% PAL (Bellefroid et al., 2018; Cole et al., 2016;Planavsky et al., 2014). In contrast, Cr isotopes from the 1.1–0.9 Ga marine carbonates show considerable fractionations and suggest atmospheric O2 concentrations of >0.1–1% PAL (Gilleaudeau et al., 2016). Similarly, high Cr isotopes from the Shennongjia Group in South China suggest atmospheric O2 levels of >1% PAL since ca. 1.33 Ga (Canfield et al., 2018).

The higher O2 estimates for mid-Proterozoic surface environ-ments would argue against oxygen limitation as an evolutionary barrier for eukaryotes (particularly animals) because many early stem-group animals and sponges have a minimum oxygen require-ment of 0.5–4% PAL (e.g., Mills et al., 2014; Sperling et al., 2013;Zhang et al., 2016). However, most geochemical data used for O2 estimation were obtained from short stratigraphic intervals in geographically distinct sedimentary basins, which leads to uncer-tainties about the stability of O2 levels in the mid-Proterozoic. Although many have assumed that atmospheric O2 concentrations remained relatively stable in the mid-Proterozoic, either at lower or higher levels (e.g., Lyons et al., 2014; Planavsky et al., 2014;Zhang et al., 2016), there exists a possibility that O2 concentra-tions fluctuated significantly during the mid-Proterozoic.

A particularly intriguing interval to test the oxygen stability is the ca. 1.57–1.56 Ga oxidation event documented from the Gaoyuzhuang Formation of the North China Platform (Zhang et al., 2018). This interval hosts decimeter-scale multicellular eu-karyotes (Zhu et al., 2016) that have been linked to oxygen increase in the early Mesoproterozoic. The associated decrease of cerium (Ce) anomaly from ∼1.0 to ∼0.8 was interpreted as evidence for significant ocean oxygenation (Tang et al., 2016;Zhang et al., 2018), but whether it represents a pulse of oxy-gen increase or the initiation of a permanently more oxygenated Mesoproterozoic ocean requires further investigation. In this pa-per, we report iodine-to-calcium-magnesium [I/(Ca+Mg)] ratios of carbonate rocks, carbonate/organic carbon isotopes (δ13Ccarb and δ13Corg), and phosphorus contents across the ca. 1.57–1.56 Ga in-terval of the Gaoyuzhuang Formation and discuss the potential temporal changes in atmospheric oxygen concentration during the early Mesoproterozoic.

2. Geological setting

Our study focuses on the carbonate rocks from Member II and III of the Gaoyuzhuang Formation in the Mesoproterozoic North China Platform (Fig. 1). The Gaoyuzhuang Formation is the basal unit of the Jixian Group that overlies the Dahongyu Formation of the Changcheng Group and underlies the Yangzhuang Forma-tion of the Jixian Group (Fig. 1D). The Gaoyuzhuang Formation is commonly interpreted as being deposited in a shelf environment of the North China Platform (Wang et al., 1985). Various zircon U-Pb ages have been reported from the Mesoproterozoic succes-sion in North China (Fig. 1D; Li et al., 2010, 2014; Su et al., 2010;Tian et al., 2015; Zhang et al., 2015). Based on the existing radio-metric ages and stratigraphic relationships, the base and top of the Gaoyuzhuang Formation are approximately assigned at ca. 1.60 Ga and ca. 1.54 Ga, respectively (Gao et al., 2009, 2010; Li et al., 2010).

The Gaoyuzhuang Formation is dominated by carbonates and can be subdivided into four members in the Yanqing area (Figs. 1D and 2). Member I consists of cross-bedded sandstones in the low-
ermost part and stromatolitic dolostones in the upper part, which
have been interpreted as deposits from shallow subtidal to in-tertidal environments (Mei, 2008). Member II consists of man-ganiferous dolostone in the lower part and medium- to thick-bedded dolostone in the upper part. The lower Member II con-tains dark shale interbeds with manganese concretions and has been interpreted as subtidal deposits (Mei, 2008). The upper Mem-ber II contains microbially induced sedimentary structures and mud cracks indicative of deposition from intertidal to suprati-dal environments (Fig. 2B). The lower to middle part of Mem-ber III is composed of thin-bedded muddy dolostone (Fig. 2C and D), calcareous mudstone with cm- to dm-sized carbonate concre-tions (Fig. 2E) and some thrombolite layers (Fig. 2F). The gen-eral lack of wave- and tide-agitated sedimentary structures in this part suggests deposition in low-energy environments proba-bly close to storm wave base (Guo et al., 2013; Luo et al., 2014;Mei, 2008), while the layered thrombolites were likely formed in subtidal environments (Tang et al., 2013). The upper part of Member III consists of laminated microbial dolostones with flaser bedding (Fig. 2G) and interference ripple marks, suggestive of de-position from shallow subtidal to intertidal environments. Member IV of the Gaoyuzhuang Formation is characterized by massive mi-crobial reef dolostones with a total thickness up to 450 m (Mei, 2008). Conical stromatolites (Conophyton-like) over 2 m high and 35 cm wide and various other microbialites (cf. Bartley et al., 2015) are common in this member, suggesting subtidal depositional en-vironments at least episodically below fair-weather wave base. Pre-vious studies on carbonate fabrics and mineral compositions of the Gaoyuzhuang Formation suggested fabric-retentive early dolomiti-zation (e.g., Zhang W. et al., 2016; Tang et al., 2017a), which may have helped preserve primary geochemical signatures.

3. Materials and methods

In this study, 234 samples from Member II and III of the Gaoyuzhuang Formation in the Gan’gou section were analyzed. Decimeter-sized multicellular eukaryotic fossils were not reported in this section, but the lithological and carbon isotope correlation indicates that the studied interval covers the fossil-bearing strata of the Qianxi and Kuancheng sections (Zhu et al., 2016). The sam-pling section is well exposed along the fresh road cuts at Gan’gou (40◦39′35.05′′N, 116◦14′35.63′′E), north of Beijing, China (Fig. 1). Collected samples were cut into chips and only the central parts of the samples were used for geochemical analyses. Fresh sample chips were cleaned, dried, and then grounded into powders (∼200 meshes) in an agate mortar avoiding any metal contact.

Macroscopic features were observed in the field. Microfabrics were examined on thin sections with a Stereo Discovery V20 mi-croscope for large scope and a Zeiss Axio Scope A1 microscope for high magnification. Ultrastructures were investigated using a Zeiss Supra 55 field emission scanning electron microscope (FE-SEM) under 20 kV accelerating voltage with a working distance of ∼15 mm, at the State Key Laboratory of Biogeology and Environ-mental Geology, China University of Geosciences (Beijing) (CUGB). Secondary electron imaging detector (SE2) was used to charac-terize topographic features, and an AsB detector was used to re-veal compositional difference (backscattered electron, BSE, image). Samples were coated with 6-nm-thick platinum for electric con-duction before analysis. A Gatan ChromaCL2 cathodoluminescence (CL) detector connected to the FESEM was used to obtain CL im-ages under 8 kV accelerating voltage with ∼30 min scanning time for each image. Element concentrations of micron-sized spots were quantitatively analyzed by an Oxford X-act energy dispersive X-ray spectrometer (EDS) connected to the FESEM, operated at 20 kV with a working distance of ∼15 mm and a beam diameter of
∼2 μm. Minerals as well as synthetic phases (MINM25-53) were

M. Shang et al. / Earth and Planetary Science Letters 527 (2019) 115797 3

Fig. 1. Geological background of the study area. (A) Major tectonic subdivisions of China. The box shows the area illustrated in panel B. (B) Generalized Mesoproterozoic paleogeographic map of the central North China Platform (modified after Wang et al., 1985). (C) Simplified geological map of the study area, showing locality of the studied section (modified after the 1:200,000 Geological Map of China, The China Geological Survey, 2013). (D) Stratigraphic columns of the Jixian Group and Gaoyuzhuang Formation in the Yanqing area, northern suburb of Beijing. Samples were collected from the upper part of Member II and Member III of the Gaoyuzhuang Formation (Li et al., 2010;Tian et al., 2015) that covers the intervals correlated to the horizons yielding large eukaryotic fossils (Zhu et al., 2016), Grypania (Niu, 1998; Sun et al., 2006) and Chuaria(Sun et al., 2006). The ages in panel D are adopted from Li et al. (2010, 2014), Su et al. (2010), Tian et al. (2015), Zhang et al. (2015) and from those summarized in Zhu et al. (2016) and Tang et al. (2017a). (For interpretation of the colors in the figure(s), the reader is referred to the web version of this article.)


Fig. 2. (A) Stratigraphic column and sedimentary facies of the Gaoyuzhuang (GYZ) Formation in the Gan’gou section, Yanqing. (B) Medium-bedded muddy dolostone with mud-cracks and ripple marks on bedding surface, indicative of shallow subtidal to intertidal facies; upper Member II of the GYZ Formation. (C) Medium- to thinly-bedded muddy dolostone, suggestive of deep subtidal to outer shelf facies around the storm wave base; Member III of the GYZ Formation. (D) Thinly-bedded, dark-gray muddy dolostone without wave-agitated structures, indicative of deep subtidal facies below fair-weather wave base; Member III of the GYZ Formation. (E) Centimeter- to decimeter-sized carbonate concretions in calcareous mudstone, suggestive of outer shelf facies close to the storm wave base; Member-III of the GYZ Formation. (F) Thrombolites with millimeter-scale mesoclots, indicative of shallow subtidal facies; Member III of the GYZ Formation. (G) Microbially-laminated muddy dolostones with chert bands, indicative of shallow subtidal to intertidal facies; Member III of the GYZ Formation. (H) Schematic diagram showing facies subdivisions used in panels A–G.


used as reference standards. Duplicate analyses of individual points showed analytical error less than 2%.

For phosphorus (P), aluminum (Al) and strontium (Sr) content analyses, ∼5 g of sample powders were directly measured using a handheld energy dispersive XRF spectrometer (HHXRF) model Xsort with a Rh anode from Spectro, following the method de-scribed in Tang et al. (2017b). The analytical uncertainty moni-tored by international reference material BCR-2 (Basalt) and JDo-1 (dolostone) was ≤7% (Table S1). For calcium (Ca) and magnesium (Mg) analyses, ∼5 g of sample powders were rinsed 4 times with 18.25 M� Milli-Q (MQ) water to remove clay minerals (Tang et al., 2017b) and any soluble salts. After drying, the samples were grounded again into smaller and more homogenized powders in an agate mortar. Finally, ∼5 mg rinsed dry powders were weighted and dissolved with 3% HNO3 and then diluted to 1:50,000 with 2% HNO3 before analyses. Concentrations were measured using a PerkinElmer NexION 300Q Inductively Coupled Plasma Mass Spec-trometry (ICP-MS) at the National Research Center for Geoanalysis, Beijing. A certified reference material JDo-1 (dolostone) was mea-sured after every nine samples and the analytical uncertainties monitored by JDo-1 (n = 16) were <3% for Mg and <2% for Ca (Table S2).

For iodine analyses, ∼5 mg of MQ water rinsed, dry powders were weighed. Nitric acid (3%) was added for dissolution and then centrifuged to obtain supernatant. To stabilize iodine, 3% tertiary amine solution was added to the supernatant, and then diluted to 0.5% with MQ water (Hardisty et al., 2017; Lu et al., 2010). The iodine content was measured within 48 hours to avoid any iodine loss (Lu et al., 2010), using a Sector Field Inductively Cou-pled Plasma Mass Spectrometry (SF-ICP-MS; Element XR, Thermo Fisher Scientific, Germany) at the State Key Laboratory of Geo-logical Processes and Mineral Resources, China University of Geo-science (Wuhan) (CUGW). The sensitivity of iodine was tuned to ∼110 kcps for a 1 ppb standard in the SF-ICP-MS. The rinse solu-tion used for each individual analysis contains 0.5% HNO3, 0.5% ter-tiary amine, and 50 μg/g Ca, and the typical rinse time is ∼1 min. Analytical uncertainties for 127I monitored by the internal standard GSR 12 and duplicate samples are ≤6% (1σ ) (Table S3). The long term accuracy is checked by repeated analyses of the reference ma-terial GSR 12 (Table S3). The detection limit of I/(Ca+Mg) is on the order of 0.1 μmol/mol. Considering the significance of samples with high iodine concentrations, we re-analyzed these samples us-ing MC-ICP-MS (Neptune Plus, Thermo Fisher Scientific, Germany) at the National Research Center of Geoanalysis, Beijing, in order to obtain more precise results. The sensitivity of iodine was tuned to ∼1,500 kcps for a 1 ppb standard in the MC-ICP-MS, and the analytical uncertainties of 127I monitored by the internal standard GSR 12 and duplicate samples are ≤7% (1σ ; Table S3). The repro-ducibility of most I contents and I/(Ca+Mg) ratios measured by the two ICP-MS setups is better than 90% (Table S4).

Carbonate carbon isotopes (δ13Ccarb) were analyzed at the LVIS Lab of the University of Nevada Las Vegas. About 50–200 μg of carbonate powders were reacted with orthophosphoric acid for 10 min at 70◦C in a Kiel IV device connected to a Finnigan Delta Plus dual-inlet mass spectrometer. The precision monitored by NBS-19 and an internal standard (USC-1; δ13Ccarb = 2.09�; δ18Ocarb =−2.08�) is better than 0.08� for both C and O isotopes. Organic carbon isotopes (δ13Corg) were analyzed at the State Key Labora-tory of Biogeology and Environmental Geology, CUGW. About 2 g of sample powders were decarbonated using 10% HCl at room tem-perature for 48 hours. The carbonate-free residue was then rinsed with deionized water repeatedly until the pH reached nearly neu-tral, centrifuged and dried at 45◦C. Dried samples were powdered, weighed (5–20 mg), and wrapped in tin capsules and combusted at 960◦C in an Elemental Analyzer (Thermo Fisher). The CO2 released
from organic matter was analyzed for δ13Corg by a Delta V Advan-
tage isotope ratio mass spectrometer (IRMS). Analytical uncertainty was less than 0.2� (1σ ) for δ13Corg based on replicate analyses of two international standards (USGS40, δ13Corg = −26.39�; IVA-Urea, δ13Corg = −37.32�).

The total organic carbon (TOC) contents were analyzed at CUGB. About 10 g sample powders were weighed, decarbonated with 10% HCl, rinsed by deionized water and dried. About 200 mg decar-bonated residues were weighed and wrapped in tin capsules and analyzed using an Elementer Macro Cube element analyzer. The uncertainty monitored by replicate analyses of an internal standard was ≤0.1%.

4. Results

All 234 samples were petrographically examined and 12 rep-resentative samples were selected for CL and EDS analyses. Pet-rographic observations show that calcite and dolomite grains are mainly ∼20 μm in size and rarely larger than 50 μm (e.g., Figs. 3and 4). CL image shows that most calcite and dolomite grains are non-luminescent, with some dolomite grains in the middle part of the studied section displaying a non-luminescent core and dull luminescent rim (Fig. 3). This core-rim structure can also be iden-tified in BSE image, with dark cores and relatively light outer rims (Fig. 4A). Quantitative EDS analyses show that the core is of low Mn and Fe contents while the light rim is slightly enriched in Mn and Fe (Fig. 4B; Table S5). The calcite has low Mn and Fe contents (Fig. 4B; Table S5). In addition, authigenic apatite grains (commonly <10 μm in diameter) were identified in Member III samples, with euhedral to anhedral morphology (Fig. 4C–H).

The most prominent feature of the study interval is the pos-itive shift in I/(Ca+Mg) ratios from Member II to Member III of the Gaoyuzhuang Formation (Fig. 5; Table S6). The variation of I/(Ca+Mg) ratios can be divided into four stages (Fig. 5). Stage I (0–60 m) has I/(Ca+Mg) values between 0.0 μmol/mol and 0.5 μmol/mol. The I/(Ca+Mg) ratios increase from ∼0.5 μmol/mol to ∼2.0 μmol/mol in stage II (60–147.9 m). In stage III (147.9–220 m), I/(Ca+Mg) ratios reach the highest value of 3.8 μmol/mol, with 30% of data points (9 of 30) higher than 2.6 μmol/mol. Almost all the I/(Ca+Mg) ratios return back to ≤0.5 μmol/mol in stage IV (220–274.5 m).

Estimates from previously published data suggest that more than 95% of the Proterozoic samples (n = 466) have I/(Ca+Mg) ra-tios of <0.5 μmol/mol excluding the samples from intervals with an oxygen rise (Fig. 6; Hardisty et al., 2017; Lu et al., 2017). Thus, the value of 0.5 μmol/mol has been taken as the baseline for I/(Ca+Mg) ratios in Precambrian carbonates (Lu et al., 2017). The I/(Ca+Mg) ratios in stage III of the study interval (Fig. 5) are higher than the maximum values reported from the mid-Proterozoic sam-ples up to date, but similar to those of the diagenetically altered carbonates found in sediments from the Clino core in the Ba-hamas (Fig. 6; Hardisty et al., 2017). They are also higher than those reported from the carbonates across the Great Oxidation Event (GOE) at ca. 2.3–2.4 Ga (∼2 μmol/mol; Hardisty et al., 2014), but are markedly lower than the highest I/(Ca+Mg) values across the Bitter Springs δ13C anomaly (∼8 μmol/mol; 810–800 Ma; Lu et al., 2017; Wörndle et al., 2019) and the Ediacaran Shuram δ13C excursion (∼8 μmol/mol; ca. 580 Ma; Hardisty et al., 2017;Wei et al., 2019).

The positive shift in I/(Ca+Mg) is accompanied by an increase of phosphorus (P) contents and P/Al ratios (Fig. 5; Table S6). Dur-ing stage I, P concentrations were stable and close to 0.05 wt%, comparable to the mean value of pre-Cryogenian shales (0.051 ±0.003 wt%; Reinhard et al., 2017). During stage II, P contents were highly variable (0.00–0.13 wt%), with peak values at the top part that are comparable to the mean value of post-Tonian shale sam-
ples (0.209 ± 0.023 wt%; Reinhard et al., 2017). The P contents


Fig. 3. Examples of Scanning Electron Microscope (SEM) and Cathodoluminescence (CL) images showing the textural and diagenetic features of carbonates from the study section. (A) Secondary electron (SE) image of a dolomite (Dol) from stage I in Fig. 5. (B) CL image of the same area in panel A, showing non- to dull (orange red) luminescence. (C) SE image of a dolomitic limestone showing calcite (Cal) and dolomite (Dol, arrows) from stage II in Fig. 5. (D) CL image of the same area in panel C, showing that calcite has non-luminescence and dolomite has the non-luminescent core and dull (orange red) luminescent rim (arrows). (E) SE image showing calcite (Cal) and dolomite (Dol, arrows) of a dolomitic limestone from stage III in Fig. 5. (F) CL image of the same area in panel E, showing non-luminescent calcite, and the non-luminescent core and dull (red) luminescent rim (arrows) of dolomite. (G) SE image showing calcite (Cal) and dolomite (Dol, arrows) of a dolomitic limestone from stage IV in Fig. 5. (H) CL image of the same area as in panel G, showing mainly non-luminescent calcite and dolomite. The magnification of all the panels are the same as in panel A.


Fig. 4. SEM image and EDS analysis of the Gaoyuzhuang carbonates. (A) Backscattered electron (BSE) image showing the core-rim structure of dolomite grains (arrows). (B) Cross plot of the Mn and Fe contents of calcite and the core and rim of dolomite (Table S5). Both calcite and the core of dolomite have low Mn and Fe contents, but the rim of dolomite has relatively high Mn and Fe contents. (C) SE image showing an euhedral apatite. (D) SE image showing subhedral to euhedral apatite aggregates. (E) Close view of the apatite aggregates in panel D. (F) BSE image showing anhedral apatite. (G) BSE image showing globular apatite aggregates. (H) EDS spectrum of the apatite in panel G. All the apatite grains were from Member III of the Gaoyuzhuang Formation.


Fig. 5. Profiles of I/(Ca+Mg), P, P/Al, δ13Ccarb and δ13Corg of Member II and III of the Gaoyuzhuang Formation at Gan’gou, Yanqing County, north of Beijing. The smoothed thick lines represent the LOESS curves. In the I/(Ca+Mg) profile, the vertical red dashed line at 0.5 μmol/mol marks the I/(Ca+Mg) baseline—the maximum [O2] level of 10 μM that could be produced by local primary production, and the red dashed line at 2.6 μmol/mol represents the estimated [O2] of >20–70 μM. In the profile P and P/Al, the vertical red dashed line at 0.05 wt% marks the mean value of P concentrations in pre-Cryogenian shales (0.051 ± 0.003 wt%; Reinhard et al., 2017). Positive shift in
I/(Ca+Mg) coincides with the increase in P and P/Al, and is accompanied with a negative shift in δ13C and δ13C .
range from 0.00 to 0.09 wt% during stage III and fall back to near zero at stage IV (Fig. 5). Consistent with the high P contents, au-thigenic apatite grains were identified in the carbonate samples of stage II and III (Fig. 4). The presence of authigenic apatite (Fig. 4) and co-varying P/Al and P content (Fig. 5) suggest that the pos-itive shift in P content was not caused by siliciclastic input. For samples with TOC contents >0.2%, the P/TOC ratios also show an increase from stage I to stage III and a decrease in stage IV (Table S6), suggesting that the P change was not caused by TOC variations in these carbonates. However, some low TOC samples (TOC <0.2%), particularly those from stage I, have higher P/TOC ratios. Consider-ing the 0.1% uncertainty related to TOC measurement, these high P/TOC ratios may have been exaggerated by the low TOC rather than high P contents (Table S6).

Both δ13Ccarb and δ13Corg show a negative excursion (Table S6) that accompanies the positive shift in I/(Ca+Mg) ratios (Fig. 5). The δ13Ccarb values range from 0% to −2� in stages I and II, and reach the minimum of −3.5� at stage III (Fig. 5). This negative δ13Ccarbexcursion has also been reported from the sections in the Ming Tombs (Li et al., 2003), Jixian (Zhang et al., 2018), and Pingquan (Guo et al., 2013) of North China (Fig. 7), suggesting that it is at least a regionally consistent δ13C excursion. The δ13Corg val-ues show some variations between −29� and −32� in stage I, and a decline to the minimum values around −34� in stage III (Fig. 5). There are some variations between the negative δ13Ccarband δ13Corg excursions, and the increase in I/(Ca+Mg) ratios oc-
curs stratigraphically ∼10 m below the decrease in δ13Ccarb, but
carb org

overall both δ13Ccarb and δ13Corg reach their minimum values at the interval with highest I/(Ca+Mg) ratios.

5. Discussion

5.1. Primary vs. diagenetic signals

Carbonate I/(Ca+Mg) ratios are susceptible to diagenetic alter-ation (Hardisty et al., 2017; Lu et al., 2010; Zhou et al., 2015;Wörndle et al., 2019). During diagenesis, carbonate-associated io-date (IO−

3 ) is reduced to iodide (I−) and is excluded from carbonate lattices, thus lowering the I/(Ca+Mg) ratios of carbonates (Hardisty et al., 2017; Lu et al., 2010). Since iodide cannot enter the lat-tice of carbonate minerals, likely due to its large ion radius (Lu et al., 2010), most diagenetic processes would decrease rather than increase the I/(Ca+Mg) ratios (Hardisty et al., 2017). In this re-gard, the prominent increase in I/(Ca+Mg) from ∼0.5 μmol/mol to >2.6 μmol/mol in stage II and III of the Gaoyuzhuang For-mation (Fig. 5) is unlikely an artifact of diagenesis. In fact, the I/(Ca+Mg) values may have been initially higher; i.e., the observed peak should be regarded as a minimum estimate of local seawater iodate levels.

Liberation of iodine from organic matter during chemical anal-ysis may artificially increase the I/(Ca+Mg) ratio, but iodine in organic matter is tightly bonded and hard to be liberated (Zhou et al., 2015, 2017). During the sectioning and chemical analyses, we have avoided samples that have dark color (potentially high
TOC) and used diluted acid (3% nitric acid; Lu et al., 2010) to min-


Fig. 6. Secular variations in I/(Ca+Mg) through time (modified from Hardisty et al., 2017; Lu et al., 2017, 2018). Vertical blue bars mark the intervals where the I/(Ca+Mg) data are consistent with possible oxygenation events (Gilleaudeau et al., 2016; Kendall et al., 2009; Planavsky et al., 2014; Zhang et al., 2016) relative to the long term atmospheric p[O2] curve (Lyons et al., 2014; Planavsky et al., 2014; Sperling et al., 2015) and biological production of O2 (Cox et al., 2018). The vertical orange bars mark the multi-cellularization of eukaryotes in the Mesoproterozoic (e.g., Zhu et al., 2016) and the diversification of eukaryotes in the Neoproterozoic (Mills et al., 2014), respectively.
The grey dashed line at 0.5 μmol/mol marks the Precambrian I/(Ca+Mg) baseline and the grey dashed line at 2.6 μmol/mol represents the threshold of [O ] >20–70 μM.
imize the iodine from noncarbonate phases (cf. Zhou et al., 2017;Wörndle et al., 2019). In addition, sample powder was rinsed 4 times to remove loosely absorbed iodine on the surface of min-erals. The lack of correlation between I/(Ca+Mg) and (CaO+MgO) (Fig. 8A), I/(Ca+Mg) and TOC (Fig. 8B), I/(Ca+Mg) and Al (Fig. 8C), but a strong correlation (R2 = 0.97) between I/(Ca+Mg) and I (Fig. 8D) suggest that the I/(Ca+Mg) is controlled by the iodine content in carbonates, rather than the carbonate and TOC contents of the samples. Therefore, the iodine from noncarbonate phases in-cluding organic matter may not be a significant contributor to the positive shift in I/(Ca+Mg) (Fig. 5).

The highest I/(Ca+Mg) values are present in partially dolomi-tized samples (Fig. 8E) and their distribution in the cross plot
of I/(Ca+Mg) vs. Mg/Ca is similar to that of the other Protero-
2

zoic samples (Fig. 8F). Most high I/(Ca+Mg) values are from sam-ples that have relatively high δ18O and Sr contents (Fig. 8G and H), suggesting that meteoric diagenesis may not be a major con-trolling factor of the I/(Ca+Mg) ratios (cf. Wörndle et al., 2019). The CL observation shows that calcite and dolomite display non-to dull luminescence (Fig. 3) indicative of low Mn and Fe con-tents that also suggest limited influence from meteoric diagen-esis. Some dolomite grains show core-rim structure in both CL and BSE images. The core is non-luminescent (CL) or dark (BSE) with low to undetectable Mn and Fe contents, while the rim is dull (CL) and light (BSE) with relatively high Mn and Fe con-tents (Figs. 3 and 4). These features suggest that the rims of dolomite grains were likely formed in the Mn reduction zone,
while the nuclei were possibly primary precipitates or formed in


Fig. 7. Correlation of the negative δ13Ccarb excursion and fossil-bearing interval in Member III of the Gaoyuzhuang Formation of the North China Platform at Gan’gou, Yanqing, northern suburb of Beijing (this study), the Ming Tombs, ∼35 km north of Beijing (Li et al., 2003), Jixian, Tianjin, ∼100 km east to Beijing (Zhang et al., 2018), and Pingquan, Hebei Province (Guo et al., 2013). In addition, macroscopic fossils are found in the Qianxi and Kuancheng sections, Hebei Province (Zhu et al., 2016). The interval with Ce
anomaly data from the Jixian section (Zhang et al., 2018) is also marked.
the nitrate-reduction zone (Canfield and Thamdrup, 2009; Liu et al., 2019). Since nitrate and Mn reduction zones are commonly very shallow and the reduction potential of Mn-oxide, iodate, and nitrate is very close to each other (Canfield and Thamdrup, 2009;Lu et al., 2010), partial dolomitization within the Mn and nitrate reduction zones must have happened during early diagenesis near the seawater/sediment contact where porewater was exchangeable with the overlying seawater. During early diagenesis, seawater or seawater-derived fluids may progressively transfer IO−

3 (and other ions and cations) into authigenic carbonates precipitated from porewater (e.g., Ahm et al., 2018; Higgins et al., 2018). However, because IO−

3 is the only iodine species that can be incorporated into carbonate minerals (Lu et al., 2010), the resulting IO−

3 con-tent and I/(Ca+Mg) ratio would not exceed that of the overlying seawater. In the samples we have analyzed, large (>50–100 μm) euhedral dolomites were rarely observed, suggesting that subse-quent burial diagenesis had limited effects on mineral textures and potentially, chemical compositions.

The low but non-zero (mostly ≤0.5) I/(Ca+Mg) values in stage I and IV (Fig. 5) fall into the I/(Ca+Mg) range of diagenetically al-tered samples (Hardisty et al., 2017), raising the possibility that diagenesis may have significantly lowered the I/(Ca+Mg) ratios in these intervals. These values, however, are also within the range of I/(Ca+Mg) ratios obtained from modern and ancient carbon-ates that were deposited in suboxic to anoxic environments (Glock et al., 2014; Hardisty et al., 2014, 2017; Lu et al., 2010, 2016; Zhou et al., 2015). Particularly, they are within the range of I/(Ca+Mg) values of Proterozoic samples from different deposi-tional environments (Fig. 6; Hardisty et al., 2017). Without addi-tional analyses such as Ca and Mg isotopes (e.g., Ahm et al., 2018;
Higgins et al., 2018), it is difficult to separate the potential diage-
netic overprint from the redox-controlled water-column signal in those samples; however, petrographic and CL observations did not show significant mineralogical and textural differences between the samples that have low (stage I and IV) and high (stage II and III) I/(Ca+Mg) values (Fig. 5). Thus, we consider that the low I/(Ca+Mg) values of stage I and IV record more reduced water col-umn conditions.

5.2. Carbonate I/(Ca+Mg) as a semi-quantitative redox proxy

Iodine-to-calcium-magnesium [I/(Ca+Mg)] ratios of carbonate rocks represent one of the few geochemical proxies that track the redox condition of shallow oceans (e.g., Hardisty et al., 2017; Lu et al., 2010, 2016). Dissolved iodine in seawater has two ma-jor thermodynamically stable forms: the oxidized species iodate (IO−

3 ) and reduced species iodide (I−). Iodate is the dominant form in oxic waters and is quantitatively reduced to iodide to-wards the core of the oxygen minimum zone (OMZ) or in anoxic waters (Emerson et al., 1979; Lu et al., 2010; Luther and Camp-bell, 1991). Since IO−

3 is the only iodine species that can incor-porate into the lattice of carbonate minerals, carbonates precipi-tated from oxic waters would have high I/(Ca+Mg) ratios, while those precipitated closer to the OMZ or anoxic waters record lower I/(Ca+Mg) ratios (Lu et al., 2010, 2016). Diagenetic alteration may lower the I/(Ca+Mg) of primary carbonates, but it would not in-crease the I/(Ca+Mg) in carbonate rocks (Hardisty et al., 2017;Wörndle et al., 2019). Our petrographic observations suggest that dolomitization of the Gaoyuzhuang Formation carbonates occurred during early diagenesis and the I/(Ca+Mg) ratios may record the
minimum estimate of water column [IO−
3 ] (and free O2) avail-


Fig. 8. Cross plots of the Gaoyuzhuang geochemical data. (A) I/(Ca+Mg) vs. (CaO+MgO), (B) I/(Ca+Mg) vs. TOC, (C) I/(Ca+Mg) vs. Al, (D) I/(Ca+Mg) vs. I, (E) I/(Ca+Mg) vs. Mg/Ca of the Gaoyuzhuang samples. The positive correlation between I/(Ca+Mg) and I indicates that the high I/(Ca+Mg) ratios relate to high iodine concentrations in carbonates. (F) Comparison of the I/(Ca+Mg) vs. Mg/Ca of the Gaoyuzhuang samples with those of the modern and ancient carbonates (Hardisty et al., 2017; Lu et al., 2017).
(G) I/(Ca+Mg) vs. δ18O of the Gaoyuzhuang samples. (H) I/(Ca+Mg) vs. Sr of the Gaoyuzhuang samples.
ability in the depositional environment (Feng and Redfern, 2018;Hardisty et al., 2017; Lu et al., 2010).

Previous studies suggested that a minimum of 1–3 μM [O2] is required for marine IO−

3 accumulation and for the presence of carbonate-bound iodine (Hardisty et al., 2014, 2017). This require-ment is supported by observations in the OMZs along the eastern coast of the north Pacific (e.g., Rue et al., 1997), the Arabian Sea (e.g., Farrenkopf and Luther, 2002; Farrenkopf et al., 1997), and in other modern anoxic basins (Luther and Campbell, 1991). In depositional environments where the OMZ is present, oxygen con-centrations vary from ∼225 μM near the ocean surface to 1–3 μM in the center of the OMZ, with IO−

3 concentrations decreas-ing from >0.25 μM to ∼0.01 μM (Rue et al., 1997). Consequently, I/(Ca+Mg) > 0.1 μmol/mol was suggested as an indicator for
water-column O2 concentration of ≥1–3 μM (Table 1; Hardisty et
al., 2014, 2017). Since diagenetic alternation of carbonates com-monly proceeds in suboxic to anoxic conditions, I/(Ca+Mg) values of carbonates should be lower than that of the primary carbon-ate sediments (Hardisty et al., 2017). In rare cases, the I/(Ca+Mg) values of carbonates could be higher than that of the primary pre-cipitates if the diagenetic fluids were highly oxic. However, this is unlikely for the Gaoyuzhuang Formation because (1) none of the adjacent Mesoproterozoic stratigraphic units produced higher (e.g., >4 μmol/mol) I/(Ca+Mg) ratios (Fig. 6; Hardisty et al., 2017) and (2) field, petrographic, and mineral observations exclude the pos-sibility that the Gaoyuzhuang Member III served as a conduit of highly oxic, Phanerozoic diagenetic fluids.

In modern oceans, due to biotic absorption of iodine and tem-porospatial redox variations (e.g., the presence/absence of OMZ),
seawater IO−
3 and O2 concentrations commonly do not show a


Table 1Semi-quantitative estimation of surface ocean O2 using I/(Ca+Mg) ratios of carbon-ates.

I/(Ca+Mg) (μmol/mol)

[IO−3 ]

(μM)[O2] (μM)

References

∼0.0 <0.01 <1–3 Hardisty et al., 2014, 20170.0–0.5 ≤0.05 ≤10 cf. Olson et al., 2013;

Reinhard et al., 2016;Lu et al., 2017

0.5–2.6 ≤0.25 ≤20–70 Lu et al., 2016>2.6 >0.25 >20–70 Lu et al., 2016

perfect linear relationship (e.g., Lu et al., 2016; Zhou et al., 2014,2015). Lu et al. (2016) compiled the IO−

3 , I/Ca, and O2 concentra-tion data from modern anoxic basins and ocean water columns with the presence of an OMZ. The results show that the I/Ca values obtained from the modern and late Holocene planktonic foraminifera have I/Ca > ∼2.5 μmol/mol corresponding to IO−

3concentrations of >0.25 μM and O2 concentrations >20–70 μM in the water column (Table 1; Lu et al., 2016). The incorporation of iodate into biogenic and abiogenic carbonates could be slightly dif-ferent: IO−

3 may be preferentially incorporated into three naturally-occurring polymorphs of calcium carbonate in the order of va-terite > calcite > aragonite (Feng and Redfern, 2018). Considering this effect, Hardisty et al. (2017) adjusted the I/(Ca+Mg) ratio of >2.6 μmol/mol as an indicator of water-column O2 concentrations of >20–70 μM. The oxygen concentrations corresponding to the I/(Ca+Mg) values of 0.0–2.6 μmol/mol are less well constrained. However, existing data from anoxic basins and from the ocean wa-ter column with the presence of an OMZ (e.g., Glock et al., 2014;Lu et al., 2016; Zhou et al., 2014, 2015) demonstrated that when I/(Ca+Mg) <0.5 μmol/mol, the corresponding [IO−

3 ] levels are <0.05 μM and the O2 concentrations are mostly <10 μM (Table 1).

The Precambrian iodine reservoir might be slightly larger than or similar to that of the Phanerozoic due to less efficient up-take and burial of iodine by primary producers, particularly al-gae (Hardisty et al., 2017). However, because I/(Ca+Mg) tracks IO−

3 rather than I− , the I/(Ca+Mg) ratio was mainly controlled by water-column redox conditions, not the iodine reservoir size (Hardisty et al., 2017). Therefore, the relationship between [O2] and carbonate I/(Ca+Mg) ratio may have been similar in Precam-brian and modern surface oceans. If this reasoning is correct, the I/(Ca+Mg) ratios from the modern environments (Table 1) may be used to semi-quantitatively estimate the temporal O2 changes across the study interval (Fig. 5).

5.3. A pulse of oxygen increase at ca. 1.57–1.56 Ga

The atmospheric oxygen level during the mid-Proterozoic has been debated. Some studies proposed that the mid-Proterozoic oxygen concentration was <0.1–1% PAL (Planavsky et al., 2014;Cole et al., 2016), while others argued that the O2 concentra-tion was much higher (>1% PAL or >4% PAL; Gilleaudeau et al., 2016; Zhang et al., 2016, 2017; Canfield et al., 2018). To date, mid-Proterozoic intervals marked with high atmospheric oxygen are only identified at 1.4 Ga (e.g., Zhang et al., 2016, 2017; Hardisty et al., 2017) and ∼1.3–0.9 Ga (e.g., Gilleaudeau et al., 2016;Canfield et al., 2018). There is no evidence for persistently high O2 concentrations during the period of 1.8–1.4 Ga.

The increase of I/(Ca+Mg) ratios from <0.5 μmol/mol to peak values of ≥2.6 μmol/mol in the Gaoyuzhuang Formation (Fig. 5) is unlikely resulted from the shoaling of the carbonate platform which may have pushed the oxycline to a greater water depth. Most carbonates of the Gaoyuzhuang Formation in the Yanqing area were deposited from shallow-water environments above the
storm wave base (Tang et al., 2016) and the highest I/(Ca+Mg)
Fig. 9. Schematic interpretation for the pulse of oxygen increase in the upper Gaoyuzhuang Formation (referring to the temporal change of data in Fig. 5). In stage I, the increase of primary productivity and organic carbon burial led oxygen increase and a relatively heavy δ13C. In stage II and III, accumulation of dissolved O2 resulted in higher IO−

3 concentrations in the surface ocean and higher carbonate I/(Ca+Mg) ratios. Oxidation of DOC led to the negative δ13C excursion and P enrichment. In stage IV, oxidation of DOC consumed oxygen and returned back to low O2 level.

values are from the interval with relatively greater depositional water depth (Figs. 2 and 5). There is no temporal relationship be-tween I/(Ca+Mg) and facies changes. Therefore, we interpret that the positive I/(Ca+Mg) anomaly in the Gaoyuzhuang Formation may record a pulse of oxygen increase in the surface ocean and
atmosphere (Fig. 9).


Taking the semi-quantitative correlation between I/(Ca+Mg) and minimum O2 levels obtained from seawaters adjacent to OMZs and in anoxic basins as a reference (Table 1), the I/(Ca+Mg) val-ues of <0.5 μmol/mol (but non-zero) in stage I (Fig. 5) would imply water column O2 concentrations between 1–3 μM and 10 μM, which are lower than the maximum O2 concentration (<10 μM) that can be produced by local oxygenic primary produc-tion in the surface ocean (e.g., Kasting, 1991; Olson et al., 2013;Reinhard et al., 2016). This estimate is consistent with many other studies that show generally low O2 levels in the Mesoprotero-zoic surface ocean and atmosphere (e.g., Planavsky et al., 2014;Reinhard et al., 2016). The I/(Ca+Mg) values between 0.5 μmol/mol and 2.6 μmol/mol in stage II (Fig. 5) may record seawater O2concentrations between 10 μM and 20–70 μM, which is higher than the maximum O2 concentration attainable through local oxy-genic primary production (e.g., Reinhard et al., 2016). Therefore, the increase of I/(Ca+Mg) from 0.5 μmol/mol to 2.6 μmol/mol may record the initial increase of atmospheric oxygen at this time. The peak I/(Ca+Mg) values of >2.6 μmol/mol (up to 3.4 μmol/mol) in stage III likely suggest seawater O2 concentrations ≥20–70 μM (Table 1). During this stage, even if oxygenic primary production resulted in a 10 μM O2 concentration in the surface ocean (e.g., Kasting, 1991; Olson et al., 2013; Reinhard et al., 2016), it would still require a significant increase (≥10–60 μM) in surface ocean O2 contributed from the atmosphere. Taking the average O2 con-centration of 250 μM as a reference for the modern surface ocean (Garcia et al., 2013), a ≥10 μM increase in surface ocean O2 would imply that the atmospheric oxygen concentration was ≥4% PAL at the time. There is an apparent decrease of I/(Ca+Mg) values in stage IV (Fig. 5). Similar to those seen in stage I, most samples from stage IV have I/(Ca+Mg) values of ≤0.5 μmol/mol, suggest-ing a surface ocean O2 level of ≤10 μM. If these I/(Ca+Mg) values were not significantly lowered by diagenetic alteration, it implies that the pulse of oxygen increase did not last long, perhaps on the order of 10 million years from ca. 1577 ± 12 to 1560 ± 5 Ma, al-though the large uncertainty of available radiometric ages prevents a more precise estimation.

A pulse of oxygen increase is consistent with the coupled nega-tive δ13Ccarb and δ13Corg anomalies and the increase of P/Al ratios at the same stratigraphic interval (Fig. 5). Negative δ13C anoma-lies could be a result of oxidation of dissolved organic carbon (DOC) reservoir in the ocean (e.g., Rothman et al., 2003) or other forms of reduced carbon such as methane (e.g., Bjerrum and Can-field, 2011), terrestrial organic matter (e.g., Kaufman et al., 2007) and petroleum (e.g., Kroeger and Funnell, 2012), authigenic car-bonate precipitation (e.g., Schrag et al., 2013; Higgins et al., 2018), and/or low primary production (e.g., Kump, 1991), and/or diagene-sis (Swart and Eberli, 2005; Swart, 2008; Derry, 2010; Oehlert and Swart, 2014). The negative δ13C anomaly is accompanied by a pos-itive shift in I/(Ca+Mg), which was considered as evidence against diagenetic alteration in reduced environments (e.g., Hardisty et al., 2017). The interval with minimum δ13C values has higher TOC contents (Table S6), excluding the possibility of a sudden decrease or shutdown of primary productivity. Although other possibili-ties may also exist, the most parsimonious interpretation for this negative δ13C anomaly would be oxidation of the oceanic DOC reservoir, which has been suggested to be much larger (10–100 times) than that of modern ocean (Rothman et al., 2003). Oxida-tion of organic matter releases phosphorous (e.g., März et al., 2008;Reinhard et al., 2017), which explains the increase of P content and P/Al ratio across the δ13C anomaly. Although the magnitude of the δ13C anomaly (−3.5�) is smaller than that of the late Neo-proterozoic δ13C excursions, it is comparable with or even larger than those of the Phanerozoic negative δ13C excursions. Oxida-tion of organic matter consumes oxygen (and oxidants) and may
significantly lower the marine and atmospheric O2 concentrations
(e.g., Bristow and Kennedy, 2008). A simple mass balance calcula-tion indicates that a −3.5� δ13C excursion formed through oxi-dation of oceanic DOC (δ13C ≈ −25�) would require ∼ 5.6 × 1017

moles of O2 (cf. Dickens et al., 1995), which amounts to ∼1.9% PAL (∼ 3.9 × 1019 moles or 21% atmospheric O2) or 2.5 times of the total dissolved O2 in the modern ocean (∼ 2.3 × 1017 moles of O2). Considering that the duration of the δ13C anomaly might be longer than the carbon residence time (105 years), the amount of O2 (oxidants) requirement would be higher than the amount sug-gested by the mass balance calculation (cf. Bristow and Kennedy, 2008). Of course, the Mesoproterozoic dissolved inorganic carbon (DIC) pool could also be different from that of the modern ocean (Rothman et al., 2003) and the increase of P may facilitate pri-mary production and oxygen increase, which complicates the O2estimation. Nonetheless, the consumption of O2 (oxidants) by oxi-dation or remineralization of DOC remains the simplest interpreta-tion for the “short-lived” pulse of oxygen increase implied by the I/(Ca+Mg) ratios from the Gaoyuzhuang Formation (Figs. 5 and 9).

The increase of oxygen during ca. 1.57–1.56 Ga is also sup-ported by the negative shift in cerium (Ce) anomaly from the Jixian Section in North China (Zhang et al., 2018). To date, how-ever, there is no documentation of this event from other con-tinents globally, possibly due to the lack of precise ages and appropriate stratigraphic record that warrants future investiga-tion. In contrast, the rise of oxygen at ca. 1.4 Ga has been documented from many sections in multiple continents includ-ing North China (Cox et al., 2016; Hardisty et al., 2017; Zhang et al., 2016, 2017), North Australia (Mukherjee and Large, 2016;Yang et al., 2017) and West Siberia (Sperling et al., 2014). There is uncertainty on whether the increase of oxygen started at ca. 1.57 Ga and continued through ca. 1.4 Ga, but the lack of evidence for persistent oxidation between ca. 1.57 Ga and ca. 1.4 Ga sug-gests that they represent two distinct episodes of Mesoproterozoic oxygenation. The surface ocean and atmospheric O2 level (≥4% PAL) estimated from the I/(Ca+Mg) ratios of the Gaoyuzhuang For-mation is significantly higher than the 0.1%–1% PAL based on Cr isotopes (Cole et al., 2016; Crowe et al., 2013; Planavsky et al., 2014). This discrepancy can be reconciled by pulsed oxygenation under the background of long-lasting low oxygen levels. Our data show that, after the peak values in stage III, I/(Ca+Mg) values fall back to Proterozoic baseline values (Fig. 5). Similar trends were also found in other Proterozoic intervals that have high I/(Ca+Mg) values, including the interpreted oxidation events at ca. 1.4 Ga, 0.8 Ga, and 0.58 Ga (Fig. 6; Hardisty et al., 2017; Lu et al., 2017;Wörndle et al., 2019). The secular variations in I/(Ca+Mg) ra-tios (Fig. 6) and many other geochemical data seem to support pulsed oxygenation events in an overall anoxic low productiv-ity mid-Proterozoic ocean (Reinhard et al., 2016; Crockford et al., 2018, 2019; Ozaki et al., 2019).

An unresolved issue relates to the cause of oxygenation dur-ing ca. 1.57–1.56 Ga and other Proterozoic intervals. Increase of primary productivity and organic carbon burial is the most effi-cient way for an oxygen increase in atmosphere and ocean (e.g., Berner and Canfield, 1989; Berner et al., 2007). Increase of pri-mary productivity and organic carbon burial is commonly accom-panied by a positive δ13C excursion. So far there is no prominent positive δ13Ccarb excursion documented from the lower part of the Gaoyuzhuang Formation prior to the interpreted oxygenation event, although an up to 5� positive shift in δ13Corg is present be-low this interval (Luo et al., 2014). Most of the existing δ13C anal-yses have been concentrated at the middle-upper Gaoyuzhuang Formation and the underlying isotope record has low stratigraphic resolution (e.g., Guo et al., 2013; Li et al., 2003; Luo et al., 2014;Zhang et al., 2018). A more detailed isotope analysis for the in-
tervals below the interpreted oxidation events at ca. 1.57–1.56 Ga


and other Proterozoic intervals remains to be investigated in future studies.

5.4. Implication for eukaryotic evolution

Eukaryotes need oxygen for respiration and their minimum oxygen requirements have been estimated as >0.1% PAL for some simple forms and higher for more complex species (e.g., Mills et al., 2014; Sperling et al., 2013). Based on the oxygen diffusion model, the size of eukaryotes scales with the minimum oxygen requirement. Particularly, the minimum O2 requirements for the Mesoproterozoic macrofossil Grypania and Chuaria were estimated as higher than 1% PAL (Runnegar, 1991). The minimum O2 re-quirements for sponges, which are one of the earliest metazoan groups, are in the range of 0.5–4% PAL (Mills et al., 2014). The esti-mated atmospheric oxygen levels of >4% PAL during ca. 1.57–1.56 Ga are apparently sufficient to support the respiration of Grypa-nia, Chuaria, and perhaps many of the early animals. Coincidentally, decimeter-scale multicellular eukaryotes (Zhu et al., 2016), Grypa-nia (Niu, 1998; Sun et al., 2006), and Chuaria (e.g., Sun et al., 2006) are present at the interval with the highest I/(Ca+Mg) values in the Gaoyuzhuang Formation (Fig. 5). Therefore, it is reasonable to assume that the increase of oxygen provided a more hospitable environment for the multi-cellularization of early eukaryotes.

The pathway that oxygen influenced the evolution of eukary-otes is not well understood. Oxygen could directly influence the evolution of eukaryotes by meeting their respiration needs (e.g., Planavsky et al., 2014) and/or indirectly influence the evolution of eukaryotes through modulating the nutrient cycle (e.g., N and P; Zhao et al., 2018). The bioavailability of N and P in the oceans is controlled by the redox conditions. More oxic oceans could maintain a stable bioavailable N (mainly NO−

3 ) reservoir (e.g., Koehler et al., 2017; Wang et al., 2018), deepen the iron redox chemocline, and reduce the P loss through Fe-oxide precipitation (e.g., Hemmingsson et al., 2018). During the pulse of oxygena-tion identified in the Gaoyuzhuang Formation, oxygen concentra-tion was likely higher than the minimum requirements of those Gaoyuzhuang eukaryotes, and the nutrient elements, such as N and P, may have been sufficient due to the expansion of more oxic con-ditions and deepened iron-redox chemocline (Fig. 9). Therefore, the evolution of eukaryotes could be accelerated during the pulse of oxygen increase.

There is no evidence, however, for continued evolution and di-versification of eukaryotes after ca. 1.56 Ga, and the first appear-ance of metazoans happened almost one billion year later in the late Neoproterozoic. The simplest interpretation could be that un-stable or fluctuating O2 levels in the mid-Proterozoic may have served as a barrier for continuous evolution of eukaryotes (cf. Cox et al., 2018; Hardisty et al., 2017; Johnston et al., 2012). The exis-tence of a potentially large DOC reservoir and reduced iron pool in the ocean (ferruginous ocean) may have been the driver for the overall anoxic surface environments in the Proterozoic—only when the extra DOC and reduced iron in the ocean were largely removed through many pulses of oxidation, the surface ocean and atmospheric O2 levels could become stable enough to support a diversified eukaryotic ecosystem.

6. Conclusions

High I/(Ca+Mg) ratios (≥2.6 μmol/mol; up to 3.4 μmol/mol) from the ca. 1.57–1.56 Ga shallow-water carbonates of theGaoyuzhuang Formation in North China suggest a pulse of oxy-gen increase to ≥4% PAL. The increase of I/(Ca+Mg) ratios is accompanied by a −3.5� δ13Ccarb and δ13Corg anomaly and an
increase in P/Al ratios that are best explained by the oxidation
of oceanic DOC. Oxidation of DOC consumed oxygen (and ox-idants) and subsequently lowered the marine and atmospheric O2 levels to the background Mesoproterozoic oxygen concentra-tion of ≤0.1–1% PAL (Planavsky et al., 2014; Cole et al., 2016). Multiple episodes of oxygen increase may have happened dur-ing the mid-Proterozoic, including those at ca. 1.85 Ga (Planavsky et al., 2018a), ca. 1.57–1.56 Ga (Tang et al., 2016; Zhang et al., 2018), ca. 1.4 Ga (Cox et al., 2016; Hardisty et al., 2017;Mukherjee and Large, 2016; Sperling et al., 2014; Yang et al., 2017;Zhang et al., 2016), ca. 1.33–1.08 Ga (Canfield et al., 2018), and ca. 1.1 Ga (Gilleaudeau et al., 2016), and these oxidation events may have been short-lived due to the consumption of oxygen through oxidation of DOC and other reductants (such as reduced iron in the ocean). Only when oceanic DOC and reduced iron in the ocean were largely consumed and the redoxcline was pushed down to the deeper ocean, stable oxygenated surface environments could be maintained to support complex eukaryotic ecosystems. Temporally isolated oxidation events in an overall anoxic mid-Proterozoic ocean explain the increase of multicellular eukaryotes at ca. 1.57–1.56 Ga (Zhu et al., 2016) and ca. 1.4 Ga (Mukherjee et al., 2018), but there is no evidence for continuous diversification of eukaryotes between the oxidation events.

Acknowledgements

Thanks are given to the editor and three anonymous review-ers for their constructive suggestions and comments, which greatly improved the paper. Thanks also go to Wenhao Zhang, Jinjian Wu, Xianghe Li, Yun Liu and Yang Li for their assistance in the field. We are grateful to Hongmei Yang, Chongpeng Liu, Zhaochu Hu, Tao He, Yong Du, Chao Wei and Xiang Li for their helps in sample pretreatments and analyses. The study was supported by the Na-tional Natural Science Foundation of China (No. 41672336) and by the Fundamental Research Funds for the Central Universities (Nos. 2652018005 and 2652017050).

Appendix A. Supplementary material

Supplementary material related to this article can be found on-line at https://doi .org /10 .1016 /j .epsl .2019 .115797.

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