Marine biogeochemical cycling during the early Cambrian constrained by a nitrogen and organic carbon isotope study of the Xiaotan section, South China

Precambrian Research 225 (2013) 148–165

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Marine biogeochemical cycling during the early Cambrian constrained by a nitrogen and organic carbon isotope study of the Xiaotan section, South China Lorenzo Cremonese a,b,∗ , Graham Shields-Zhou a,c , Ulrich Struck b , Hong-Fei Ling d , Lawrence Och a , Xi Chen d , Da Li d a

UCL–Gower Street, WC1E 6BT London, UK Museum für Naturkunde, Invalidenstrasse 43, 10115, Berlin, Germany c NIGPAS, Nanjing, 39 East Beijing Road, Nanjing 210008, China d Nanjing University, Nanjing 210008, China b

a r t i c l e

i n f o

Article history: Received 24 June 2011 Received in revised form 5 December 2011 Accepted 6 December 2011 Available online 18 December 2011 Keywords: Nitrogen isotopes Precambrian–Cambrian boundary Yangtze Platform Nutrient cycles Chemocline

a b s t r a c t The Precambrian–Cambrian boundary is claimed as one of the most crucial biological breakthroughs on our planet, when changes in chemical and physical conditions, together with key biological innovations, helped to trigger a biodiversity “explosion”. The Yangtze Platform (South China), mainly characterized by continuous and unaltered successions of this age, is ideally suited to high-resolution, palaeomarine investigations of this event. In this study, ␦15 N and ␦13 Corg records from Xiaotan Section (Yunnan, SW China) were investigated in order to provide insight into variations in primary productivity, ecological developments and marine environment. The Ediacaran-early Cambrian Xiaotan section is characterized by relatively high sedimentation rates and variable lithologies (carbonates, cherts, phosphorites, sandstones and siltstones) that alternate through nearly 600 meters thickness. Organic carbon isotope values vary between −36‰ and −21‰, tightly following and helping to complete published ␦13 Ccarb trends. The base of the Cambrian is characterized in this section by a significant drop from −25‰ to −35‰, mirroring the ␦13 Ccarb fall observed at this and other sections worldwide. Higher in the section in the Dahai Member, values increase to −20‰, again demonstrating communication between DOC and DIC pools during the early Cambrian on 105 –106 year time scales. From the base of the overlying Shiyantou Formation, ␦13 Corg values remain around −30‰ until the end of the section, testifying to an interval of more modest change in DIC ␦13 C that would be consistent with biostratigraphic correlations. Nitrogen isotope values vary independently from carbon isotope trends, exhibiting several major ␦15 N cycles, with values fluctuating between +9‰ and −1‰. Robust trends in ␦15 N within the Dahai Member testify to changes in the ratio between nitrogen fixation and denitrification, possibly reflecting fluctuations in the water column chemocline as ␦13 C values reached their global acme. Up section, nitrogen isotope values show dampened cyclicity with values remaining between 0‰ and +5‰, indicating establishment of an equilibrium state in marine biogeochemistry and nutrient cycles. The overall variability in ␦15 N closely resembles that of recent marine sediments, while the observed cyclicity is defined by several samples in all cases. These observations argue for good preservation of original isotopic signatures which bodes well for future N isotope studies of the Precambrian-Cambrian boundary. Crown Copyright © 2011 Published by Elsevier B.V. All rights reserved.

1. Introduction The Precambrian–Cambrian boundary records a substantial change in the interactions between geological and biological cycles

∗ Corresponding author at: 2nd floor Pearson Building, UCL, Gower Street, WC1E 6BT London, UK. Tel.: +44 20 7679 7821, fax: +44 20 7679 2433. E-mail addresses: [email protected] (L. Cremonese), [email protected] (G. Shields-Zhou), [email protected] (U. Struck), hfl[email protected] (H.-F. Ling), [email protected] (L. Och), [email protected] (X. Chen), [email protected] (D. Li).

due to evolutionary diversification amid considerable plate tectonic reconfiguration. During earlier geological times, organisms populating the biosphere were considerably less diverse, largely immobile and soft-bodied. The so-called Cambrian “explosion” of life marks a fundamental threshold in Earth history and represents a step change in the active participation of biochemical processes in planetary scale changes. The influence of biological activity on the whole Earth system changed crucially after the C transition, developing more sophisticated feedbacks (Falkowski et al., 2008; Saltzman, 2005; Logan et al., 1995). In fact, different from purely inorganic reactions driven by tectonic and atmospheric photochemical processes, biologically-controlled transformations

0301-9268/$ – see front matter. Crown Copyright © 2011 Published by Elsevier B.V. All rights reserved. doi:10.1016/j.precamres.2011.12.004

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are generally characterized by great efficiency through biological catalysis (active enzymes) and life’s capability to adapt to changing ecological conditions through advantageous gene mutations. The development of life on Earth added new vigorous biological activities to the already existent geological cycles, contributing to a greater complexity and novel solutions to buffer periodic imbalances in the Earth system. New cause-effect systems characterize the Phanerozoic, during which biological activity played and continues to play a crucial role in driving and forcing Earth system feedbacks (Falkowski et al., 2008). Many aspects related to the PcC boundary remain unsolved, stimulating researchers’ interest in further investigation into the timing and cause-effect mechanisms at the beginning of the Phanerozoic Eon. The Precambrian–Cambrian boundary defines the passage between the Proterozoic and the Phanerozoic eras, bridging the Cryogenian (850–635 Ma) and the newly-defined Ediacaran (ca. 635–542 Ma) periods with the Cambrian (542–493 Ma) Period. The Ediacaran Period, in particular, is characterized by concurrent profound changes in global tectonics (Li et al., 2010; Dalziel, 1997), possibly the largest climatic deviation since the Palaeoproterozoic popularized by the “Snowball Earth” concept (Fairchild and Kennedy, 2007; Hoffman et al., 1998), the first evidence leading up to the possibly “explosive” but indisputably largest evolutionary radiation event of the biosphere (Cohen et al., 2009; Grotzinger et al., 1995; Shen et al., 2008), and a significant increase of free atmospheric and oceanic oxygen (Fike et al., 2006; McFadden et al., 2008; Johnston et al., 2009). Younger Cambrian strata record the gradual establishment of a newly acquired ecological stability (Logan et al., 1995; Hu et al., 2007; Weber et al., 2007; Payne et al., 2009; Vannier et al., 2007), the onset of the conquest of Earth‘s surface environment by Metazoa (Maloof et al., 2010; Steiner et al., 2007; Li et al., 2007), and a profound rearrangement in the chemical composition of Earth’s atmosphere and oceans (Wille et al., 2008; Canfield et al., 2007; Hough et al., 2006; Chen et al., 2009; Anbar and Knoll, 2002). Despite the lack of direct observations on seawater chemistry, the study of nitrogen and organic carbon stable isotopes shows great potential for paleoenvironmental and biogeochemical reconstructions (Garvin et al., 2009; Godfrey and Falkowski, 2009; Owens, 1987; Peterson and Fry, 1987; Struck et al., 2001). Most abiotic and biotic reactions occurring in seawater are associated with isotopic fractionations to different extents and magnitude, particularly conspicuous in the case of nitrogen because of the numerous oxidation states, molecular diversity and microbial typologies involved (Canfield et al., 2010). It is then conceivable, through geochemical and biochemical investigation of marine sedimentary rocks, to delineate the ocean nitrogen cycle in greater detail. The total nitrogen fraction stored in ocean sediments derives from the dissolved nitrogen in seawater principally in the form of nitrate, dinitrogen or ammonia, which has been exported to the sea floor through the organic matter sink (export productivity). After processes like nitrogen fixation and assimilation by organisms, organic nitrogen as well as inorganic nitrogen incorporated in clay minerals is transported to the sediments and preserved through geological time. The isotopic values measured from these sediments are believed to represent the isotopic signal of nitrate in the ocean at that time, assuming negligible alteration due to diagenesis or other forms of post-depositional alteration (Altabet et al., 1999; Freudenthal et al., 2001; Lehmann et al., 2002; Holmes et al., 2002; Lourey et al., 2003; Struck et al., 2001). Therefore, the ␦15 N value measured in sediments can directly reflect the nitrogen biogeochemical cycle acting in seawater during the interval of sedimentation. Coupling isotopic study with sedimentological observation could help toward a better comprehension of depositional setting and water column depth.

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The use of nitrogen isotopes for seawater investigations has been debated in the literature for a long time because of a supposed strong susceptibility to diagenesis together with its complex biological cycle leading to its general rejection as a chemostratigraphic tool for paleoenvironmental reconstructions. In contrast to this opinion and in the light of our results, we aim to demonstrate its potential for paleoceanographic reconstruction reinforced by organic carbon isotope results. Since nitrogen is only stored in kerogen and clay mineral lattices in sedimentary rocks (mostly montmorillonite and illite; Godfrey and Falkowski, 2009; Krooss et al., 2006; Calvert, 2004) at minor or very low concentrations, nitrogen isotope investigations are often hampered by experimental limits, restricting its use to particularly organic-rich successions only. Although we faced problems of low nitrogen concentrations, we managed to fill gaps in isotopic trends by decalcifying samples in order to concentrate the N-bound fractions. Our outcomes show how important environmental information may be obtained by studying these two proxies together, both being strongly correlated with productivity, redox state and chemical stratification. The crucial Pc-C boundary, although well characterized in terms of biostratigraphy and sedimentology, still lacks detailed organic and biogeochemical study.

2. Geological setting According to palaeomagnetic records, the South China block was positioned at low to intermediate latitudes during the two global glaciations in Neoproterozoic time, i.e. at 40◦ latitude during the mid-Cryogenian ‘Sturtian’ glaciation (ca. 720 Ma; Fanning and Link, 2004 Rui and Piper, 1997) and near the equator during the late Cryogenian ‘Marinoan’ glaciation (ca. 590–610 Ma; Macouin et al., 2004; Hoffmann et al., 2004). The Yangtze block and the Cathaysia block together form the South China Block (SCB), which was generated during the early Neoproterozoic Sibao-Jinning Orogeny at ca. 1000 Ma or >900 Ma (Li et al., 1995; Wang, 2000; Li et al., 2002, 2003a, 2005). Today, the Yangtze platform is located between the Cathaysia suture to the southeast and the Qin-Lin orogenic belt to the North. The timing of collision between the Yangtze and Cathaysia blocks is controversial, as is its position toward the centre of the supercontinent Rodinia, because of geochemical evidence suggesting active continental margins around the Yangtze Block during the early Neoproterozoic that post-date the amalgamation of the Yangtze and Cathaysia blocks around ca. 800 Ma (Zhou et al., 2002a,b, 2003, 2004) and not related to the Grenvillian orogeny around 900 Ma as previously claimed (Li et al., 1995, 1999). Additional recent interpretations based on U-Pb dating support this late collision between the two blocks, claiming that South China was instead placed in a marginal position relative to the supercontinent Rodinia and dating the amalgamation of the Yangtze and Cathaysia blocks at ca. 830 Ma. Late Neoproterozoic strata (Nanhua basin) would have been unconformably deposited on top in the intracontinental Nanhua basin in response to backarc spreading and subsidence caused by the rollback of the subductive slab (Zhao et al., 2011). Some authors support a mantle plume cause for the breakup of Rodinia between ca. 830 Ma and 740 Ma on the western, northwestern, and southeastern margins of the Yangtze Block (Li, 1999; Ling et al., 2003; Li et al., 2003a,b), creating several rift basins as described by Wang and Li (2003). During the late Ediacaran, these shallow water platforms expanded and much of the northern Yangtze Block is characterized by shallow-water platform facies. After formation of sedimentary basins, the distribution of Neoproterozoic successions was strongly influenced by syndepositional tectonic activity, shaping and modifying constantly their geometries. During and following the breakup of Rodinia, the

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Yangtze platform moved away from the rest of the supercontinent, becoming isolated and unique in its Neoproterozoic and Paleozoic stratigraphy. The Ediacaran successions in South China were deposited after the termination of global glaciation at 635 Ma, not simply overlapping the rift successions in the basins on the margins of the Yangtze Block, but also covering the entire block. Thus, a variety of Ediacaran facies associations from shallow to deep water (former Nanhua rift basin) were developed across South China. Depositional settings and water conditions changed significantly in the Yangtze basin at the onset of the Cambrian Period. It seems likely that global tectonics, reorganisation of ocean currents and the rising influence of metazoans led to changes in ocean composition, sedimentary deposition and oxygen distribution in the basin. The Sichuan Massif and the Kangdian Massif developed as the earliest continental nuclei in East Sichuan/West Hubei and East Yunnan (Wang and Mo, 1985). These regions represented sedimentary highs during the Early Cambrian, from which sediment was delivered into the nearby shallow seas and the oceanic basins. Fine clastic sedimentation was predominant across vast regions of the Yangtze Platform, particularly during the Qiongzhusian and Canglangpuian (broadly equivalent to the Atdabanian and Botoman stages of Siberia, respectively) of the early Cambrian. 2.1. The Dengying Fm. The Dengying Fm. (Fig. 2g) overlies conformably the Doushantuo Fm. Its basinal equivalent is the Liuchapo Fm., composed of black cherts, silicified layers and black shales. The total thickness in the Yichang (Three Gorges area) area where it crops out in towering cliffs is only about 100 m (e.g. the Jiulongwan section), but can reach 800 or 900 m in the area of our study in northern Yunnan and eastern Sichuan (Steiner et al., 2005). The Dengying Formation is commonly composed of three parts (in ascending order): the Hamajing, Shibantan and Baimatuo members (Three Gorges Area) and Donglongtan, Jiucheng and Baiyanshao members (Yunnan Province). The uppermost member usually consists of 40–400 m thick massive micritic and recrystallized dolomite with an erosional surface at the top that is overlain by various Cambrian successions (Zhu et al., 2003). The tubular fossil Sinotubulites, which may represent the earliest shell-producing metazoan, was found in shelly beds in the lower part of the Baimatuo Mb. (Chen et al., 1981; Ding et al., 1992). The depositional environment of the Dengying Fm. is interpreted as a widespread prograding platform accompanied by sea-level fall. Oolitic textures and oncolites in the dolomitic Hamajing Mb. are characteristic of a high-energy environment following black shale deposition at the top of the Doushantuo Fm. Toward the southeast, the carbonate successions become gradually more condensed, ultimately changing into the slope and basinal facies of the corresponding Liuchapo Fm. In the Xiaotan section of NE Yunnan, only the upper Dengying Fm. is well exposed and so discussed here. 2.2. Early Cambrian stratigraphy in South China The Cambrian Period defines the beginning of metazoan expansion and eukaryotic diversification, with its lowest part stratigraphically correlated using assemblages of Small Shelly Fossils (SSF’s) (Steiner et al., 2007). In 2004, the International Subcommission on Cambrian Stratigraphy proposed to divide the Cambrian Period into four Series, each subdivided into two or three stages whose beginning would be defined by the First Appearance Datum (FAD) of a specific fossil biomarker. Although correlation across the Yangtze Platform is generally achievable using SSF assemblages, correlation further afield is complicated by the

endemic nature of many early Cambrian fossil taxa (Steiner et al., 2007). The Ediacaran–Cambrian boundary is disconformable in many places in South China, particularly in the shallow-water platform facies where in some cases several million years of deposition may not be recorded. The very base of Cambrian units appear often condensed or absent also in apparently deeper water environments, characterized by deposition of just a few cm of black shale and phosphatic nodules (Steiner, 2001; Steiner et al., 2007). On the northwestern part of the platform, the base of the Cambrian is represented by carbonate and phosphorite rocks, overlying the top of the Baimatuo Mb. or equivalent top-Dengying Fm. unit disconformably (Fig. 2c). The lower Cambrian in South China often contains a sulphidic metal-rich layer 10–30 cm thick, where Mo, Ni and other heavy metals are present in high concentration (Fig. 2b; Coveney et al., 1991; Wille et al., 2008; Pi et al., this volume). On the Yangtze platform, one of the best outcrops for the basal Cambrian is observed at the Xiaotan section, North Yunnan (Figs. 1 and 2 a). Here a relatively high sedimentation rate created one of the best examples of early Cambrian stratigraphy. At this section, the boundary between the Dengying and Zhujiaqing Fm. is observed as an undulating erosional disconformity, above which an apparently uninterrupted and well preserved sedimentary succession is clearly displayed from the very beginning of the Zhujiaqing Fm. (Daibu Member) until massive sandstones belonging to the Canglangpu Fm. (still Cambrian II). The thickness of the early Cambrian succession at Xiaotan is greater than at other locations in South China, especially from shelf and platform settings. 2.3. The Zhujiaqing Fm. 2.3.1. Daibu Mb The Daibu Member represents the first unit of the Zhujiaqing Fm. It consists of grey siliceous microsparite dolostones organized in centimetre to decimetre-bedded layers, with cherts in the form of nodules or intercalated laminae (Fig. 2d). Its total thickness is about 50 meters, delimited in its upper part by the first appearance of phosphorite layers. No fossils have been found in this member (Zhou et al., 1997). 2.3.2. Zhongyicun Mb. The second unit of the Zhujiaqing Formation, the Zhongyicun Member, is a dolomitic unit characterized by interlayers of sandstone, chert or phophorite in nodules or thin beds (Fig. 2e). The onset and end of this member are marked by thick phosphatic layers. The contact with the Daibu Mb. is transitional in this section. Dark cherts in well-defined decametric layers are present all through the member. About two meters of black shales are intercalated between the two main phosphorite levels and several sandy levels were also observed during sampling. It is in the Zhongyicun Mb. where skeletal microfossils such as Conotheca (Missarzhevsky, 1969), Turcutheca (Missarzhevsky, 1969), Protohertzina (Missarzhevsky, 1973), Siphogonuchites (Qian, 1977), and Hyolithellus (Billings, 1871) make their lowest appearance at Xiaotan (Li and Xiao, 2004). Using the standard classification system of the early Cambrian in China, this member has been referred to the SSF assemblages I and II (Anabarites trisulcatus–Protohertzina anabarica and Siphogonuchites triangularis–Paragloborilus subglobosus, respectively; Zhou et al., 1997). 2.3.3. Dahai Mb The last unit of the Zhujiaqing Formation (the Dahai Member) consists, at the Xiaotan section, of nearly 100 meters of fairly regular calcitic pelmicrites, often thickly bedded with abundant SSFs characteristic of the “Heraultipegma yunnanensis” (Watsonella crosbyii)

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Fig. 1. Paleogeographic map of South China and stratigraphic settings of the Xiaotan section. SSF assemblages are depicted after Zhou et al. (1997) and Steiner, M. (personal communication), together with Cambrian Series and Stages according to the International Subcommission on Cambrian Stratigraphy (ISCS). DGY = Dengying, CLP = Canglangpuian, SSFA = Small Shelly Fossil Assemblage.

assemblage (SSF assemblage III). Millimetric phosphorite nodules are found at several levels within this member.

The lower Shiyantou Fm. is devoid of SSFs. However, thin-bedded and lenticular limestones in the upper Shiyantou Fm. contain SSF’s of the Sinosachites flabelliformis–Tannuolina zhangwentangi assemblage (SSF assemblage IV; Zhou et al., 1997).

2.4. The Shiyantou Fm. The overlying Shiyantou Fm. is 250 m thick, and its lower parts contain several silty and sandy layers, followed by tens of meters of dark nodular) and lenticular carbonate beds with relatively high organic content (1–2%). The first 20 m of this formation are characterized by septarian carbonate concretions up to 40 cm in diameter (Fig. 2h); this was also observed at other areas in South China (e.g., Meishucun, Yunnan). About 30 m from the base of the formation, a long series of sandy-clay siltstones with high organic carbon content begins for a total thickness of nearly 150 m. The upper part of the Shiyantou Fm. is characterized by very finely laminated dark siltstones containing large carbonate concretions (up to 60 cm in diameter, Fig. 2f), overlain by greenish, grey/dark grey finely laminated siltstone with sporadic carbonate beds 10–20 cm thick. Close to its top, still abundant, but smaller, elongated, sometimes lenticular carbonate concretions are present.

2.5. The Yuanshan Fm The base of this formation is marked by renewed deposition of 50 m of black shales and carbonaceous siltstones, turning more thinly bedded (1 cm) up section. This unit is considerably darker than the top of the Shiyantou Fm., and the boundary with the underlying Shiyantou Fm. is marked by a 15 cm layer containing decimetric calcitic concretions. The second unit is principally a siltstone succession with rare limestone lenticles and interbeds (Fig. 2g) for a total thickness of 90 m. Close to the top part of the formation, layers turn massive and dark grey in color. The basal unit contains SSFs of the fifth assemblage in calcareous horizons. At about 19 meters from its base, the trilobite Wutingaspis first appears in silty layers, followed by the appearance of Eoredlichia (Luo et al., 1984).

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Fig. 2. (a) Panoramic view of Xiaotan section; (b) Calcareous nodules interbedded in black-shales, Yuanshan Fm.; (c) Dahai Mb.; (d) Chert nodules in Daibu Mb.; (e) Phosphorites in Zhongyicun Mb.; (f) Carbonate nodules at the top of Shiyantou Formation; (g) Dahai/Shiyantou boundary; (h) Septarian concretion in black-shales, basal Shiyantou.

3. The marine nitrogen cycle Nitrogen exists as two stable isotopes in nature: 14 N (99.63%) and 15 N (0.37%). The N cycle in marine environments consists of many exchange processes among living matter, water and sediments. The numerous oxidation states of inorganic nitrogen (−4, 0, 2, 3, 5) and the multiplicity of biological pathways interconnecting different pools often make it difficult to identify clearly the key processes controlling nitrogen through ecosystems, conferring to the whole cycle greater complexity than other bioelements’ cycles.

For convenience, isotope ratios in samples containing only slight 15 N enrichments or depletions, are reported on the ␦ scale: ␦15 N = 1000 × (Rsample − Rstandard )/Rstandard , where Rsample stays for ([15 N]/[14 N])sample . The standard is atmospheric N2 (Rstandard = 0.0036765, ␦15 N = 0‰). 3.1. The nitrogen biogeochemical cycle The N isotopic composition of nitrate, particulate matter and sediments is a diagnostic tool widely used to investigate N-cycle dynamics of the modern ocean, such as nitrate utilization, N2

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Fig. 3. Simplified structure of nitrogen biological cycle in marine environment: modern system. Standard ranges of ␦15 N in seawater for each species are reported in brackets and represent values in normal conditions. Modified after Beaumont and Robert (1999).

fixation and denitrification rates. A simplified model for the biological nitrogen cycle in the modern marine environment is shown in Fig. 3 in which biogeochemical pathways are depicted by arrows. Atmospheric N2 enters the biogeochemical nitrogen cycle via biological fixation (pathway 1 in Fig. 3) achieved by prokaryotes in the bacterial and archaeal domain. Most of these organisms are represented by oxic and anoxic phototrophic cyanobacteria, able to transform inorganic dinitrogen into organic-bounded ammonium through action of the nitrogenase enzyme (generally called “diazotrophs”). After the death of these organisms, nitrogen contained in their nucleic and amino acids returns to the environment via mineralization (pathway 2) which leads to the formation of NH4 + or NO2 − /NO3 − according to oxygen concentrations in seawater. Ammonium derived from organic matter degradation or from reduction of nitrate through dissimilatory nitrate reduction to ammonium (DNRA, pathway 3) can be assimilated by other prokaryotes (pathway 4). If oxygenated conditions are established, ammonium can also be back transformed to NO3 − by nitrification, mediated by specific groups of bacteria and archaea (pathway 5). An alternative route for ammonium is to be oxidized to dinitrogen by strictly anaerobic chemoautotrophic bacteria when water redox and chemical conditions support nitrite in the seawater. This process, called anammox (anaerobic ammonium oxidation, pathway 6), is driven by a group of bacteria called planctomycetes and seems to occur nowadays only in few areas like the Arabian Sea (Ward et al., 2009), the Eastern Tropical South Pacific OMZ (Lam et al., 2009; Thamdrup et al., 2006; Hamersley et al., 2007), and in anoxic waters such as the Golfo Dulce (Dalsgaard et al., 2003) and the Black Sea (Kuypers et al., 2003), although its recent discovery calls for further investigations (Thamdrup and Dalsgaard, 2002). On the other hand, nitrate ions can undergo canonical heterotrophic denitrification in the absence of oxygen, coupled to the anaerobic oxidation of organic carbon releasing N2 back into the atmosphere (pathway 7). Acting as a respiratory electron acceptor during oxidation of organic matter, they regenerate essential macronutrients such as carbon dioxide (CO2 ), ammonium (NH4 + ), and phosphate (PO4 3− ) necessary to sustain seawater primary productivity. Denitrifying organisms include representatives of more than 60 genera of Bacteria and Archaea, as well as some eukaryotes. Heterotrophic and autotrophic (anammox) denitrification represents the last step in the nitrogen cycle, returning it to the atmosphere. Organic nitrogen that reaches the seafloor and

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escapes mineralization (pathway 2) is incorporated into sediments and transformed to kerogen during sedimentation and subsequent diagenesis (pathway 8). Organic nitrogen can also pass higher in the food web through predation by zooplankton on phytoplankton. Nitrogen fixation is accompanied by little or no isotopic fractionation (there is no preference by these organisms for 14 N or 15 N), so any input of combined nitrogen to the ecosystem through this way will have a ␦15 N of nearly 0‰ (Altabet, 2007; Hoering and Ford, 1960; Wada et al., 1975; Wada and Hattori, 1976; Mariotti et al., 1980; Minagawa and Wada, 1986; Macko et al., 1987). It has been observed that Fe concentration during dinitrogen uptake may condition considerably the isotopic value, lowering it in specific cases to −3‰ (Zerkle et al., 2008). This represents the only mechanism for new nitrogen to enter the seawater biochemical cycle. Isotopic fractionation associated with ammonium assimilation and general metabolism of DIN by phytoplankton typically produces phytoplankton biomass depleted in 15 N relative to the substrate used for growth. The degree of isotopic discrimination associated with primary production varies between taxa and with the growth conditions (Montoya. et al., 1992). Moreover, a general increase in ␦15 N with trophic level is typically observed in marine ecosystems, with an average effect of ca. 3.5‰ for each level (e.g. Phytoplankton → Herbivores → Carnivores). Nitrate denitrification (NO3 − /NO2 − /NO/N2 O ⇒ N2 ) is responsible for the difference in nitrogen isotopic signature between atmospheric and organic nitrogen. During heterotrophic denitrification (the major biochemical process responsible for marine nitrogen loss) 14 N contained in oceanic nitrate returns preferentially to the atmosphere, leaving seawater NO3 − enriched in 15 N in seawater (Liu and Kaplan, 1988; Mariotti et al., 1980; Wada et al., 1975). This effect can only be detected in cases where the nitrate pool has been completely exhausted (Godfrey and Glass, 2011; Meyers et al., 2009), while denitrification occurring within sediment pore waters does not induce large variations (maximum 3‰) in the nitrate isotopic signature of seawater because of weak exchanges between the two reservoirs (Sigman et al., 2009; Galbraith et al., 2008). For this reason, marine organisms are systematically enriched in 15 N by ca. +6‰ relative to the atmosphere whenever conditions of “normal marine productivity” are established (Montoya et al., 2002).

3.2. ı15 N as a geochemical proxy Changes in the global rate of denitrification and nitrogen fixation/assimilation may significantly alter the isotopic composition of marine NO3 − , causing global shifts in the average ␦15 N of marine ecosystems. At a smaller scale, the ␦15 N of sedimentary nitrogen may help to characterize the relative extent of nutrient utilization in the water column (Francois et al., 1992; Altabet and Francois, 1994). In a system where nitrogen loss by denitrification or assimilation exceeds nitrogen supply in the pool, the remaining nitrate will become progressively enriched in 15 N, with the same fate for the produced biomass. Should this process consume the entire nitrate fraction, the ␦15 N of the integral curve area will as a whole coincide with the substrate ␦15 N (Rayleigh fractionation). The ␦15 N value of +5‰ of modern marine sediments (Altabet and Francois, 1994) confirms essentially complete nitrate utilization in todayˇıs ocean inasmuch as the ␦15 N value of dissolved nitrate averages 4.8‰ (Sigman et al., 1999). Beneath areas with extensive denitrification, variable but relatively high ␦15 N in sedimentary records reveal a history of the changing influence of denitrification in such environments. This in turn would affect the ␦15 N of marine plankton, since

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the isotopic composition of the zooplankton depends most on the ␦15 N of the available DIN, and consequently all the nitrogen pool. In modern oceanic environments, the O2 /H2 S interface is, in most cases, buried beneath the sediment-water interface except for meromictic lakes, fjords and some closed basins like the Black Sea and the Cariaco Trench. However, on the early Earth when atmospheric O2 levels were substantially lower than at present, the O2 /H2 S interface could have moved into shallower portions of the water column (Sinninghe Damstè et al., 1998). Samples with negative ␦15 N values (generally from 0‰ to −8‰) have been potentially influenced by anaerobic primary production, possibly by Green or Purple sulfur Bacteria (Pennock et al., 1996; Fogel and Cifuentes, 1993) as in the modern Black Sea (Fuchsman et al., 2008), Lake Kaiike (Ohkouchi et al., 2005) and in Eastern Mediterranean Sapropels (Passier et al., 1999; Struck et al., 2001). These organisms are able to fix dinitrogen, ammonia and/or nitrite in their tissues in the euphotic (sulfidic) anoxic zone with a fractionation ranging from 0 to −2‰ for bacteria using N2 as substrate and from −25‰ to −8‰ (depending on ammonia concentration) when using ammonia, e.g. Skeletolema Costatum (Ohkouchi et al., 2005; Pennock et al., 1996). Several studies have demonstrated how ␦15 N can be modified during diagenesis, overprinting the original signature. These variations typically range from +4‰ in oxic and low sedimentation rate regimes (Gaye-Haake et al., 2005; Smith et al., 2002; Altabet and Francois, 1994) to −2‰ in suboxic and anoxic domains (Lehmann et al., 2002; Freudenthal et al., 2001; Voss et al., 1997). Other investigations support good conservation of the original isotopic signature especially in areas characterized by high sedimentation rates, highlighting its utility as a geological proxy (Altabet, 1996; Altabet et al., 1999, 2002; Smith et al., 2002; Gaye-Haake et al., 2005; Galbraith et al., 2008; Godfrey and Falkowski, 2009; Thomazo et al., 2011 and references therein). In previous studies, it was generally not possible to assess “a priori” whether nitrogen isotopic signatures represented a robust biogeochemical proxy, but in our opinion each case will be different, depending on nitrogen concentration, C/N ratio, lithology and diagenetic conditions. In conclusion “normal” marine production is characterized by a state of equilibrium between nitrate assimilation, N2 fixation and denitrification, stabilizing values inside the range +2‰ to +6‰ (modern oceanic values). During periods of high denitrification rates accompanied by complete nitrate utilization the N isotope ratio of seawater increases to values above +6‰. When biochemical settings favor N2 fixation in the water column (formation of a stable pycnocline or fall in the N/P ratio), an isotope value averaging zero is recorded for the particulate organic matter (PON), or can be negative in cases where anoxic conditions reach the euphotic zone. It is worth noting that biogeochemical processes involving nitrogen occur at the same time, resulting in the settling of a mix of organic compounds bearing nitrogen with different isotopic values. The final value we measure is, then, a “summation” of different processes which means that cases of data-overlapping or even deletion of meaningful evidence is sometimes possible. The isotopic variability shown by this proxy in modern aquatic environments limits its use to local reconstructions and probably rules out the possibility of global correlations (Quan et al., 2008). The nitrogen isotopic composition in sedimentary organic matter can thus constitute a tracer of major changes in the nitrogen cycle (Wada and Hattori, 1991). A summary of possible interpretations of major biogeochemical processes occurring in the seawater, based on the nitrogen isotopic composition measured in coeval sediments, is provided in Table 1. Given the large range of processes that can occur in natural systems, these values must be considered indicative rather than diagnostic.

Table 1 ␦15 N

Dominant process in seawater

−4‰ to −1‰

Euphotic zone anoxia and purple sulfur bacteria Dinitrogen uptake by cyanobacteria Equilibrium between N assimilation, N2 fixation and denitrification (“normal marine production”) High denitrification rates and complete nitrate assimilation

−2‰ to 1‰ +2‰ to +6‰

> +6‰

4. Methodology We collected about 250 rock samples from the Xiaotan section, north-eastern Yunnan, representing different lithologies (carbonates, phosphorites, cherts as well as black shales and organic-rich rocks at higher sampling resolution). N and Corg isotopic analyses were performed at the laboratories of the “Museum für Naturkunde” in Berlin. The device used was a Thermo Finnigan Elemental Analyser Mass Spectrometer - continuous flow. After rock cutting, and selecting unaltered sections of the sample (discarding recrystallized veins and air-exposed sides), samples were powdered, weighed and wrapped in tin foil before measurement. For bulk rock samples, the amount of powder required for a detectable amount of nitrogen ranged between 10 mg and 200 mg depending on clay and/or organic content. The natural drift observed for the reference has been used to correct the samples’ values. All the results characterized by a TN below 0.020 mg (accuracy limit for our AE) have been measured at least twice to demonstrate their fidelity. Nevertheless, in many cases the extremely low nitrogen content did not allow us to obtain a reasonable ␦15 N, obliging us to omit these samples from the present investigation. The final analytical errors for ␦15 N analyses amount to ±0.25‰. Concerning the Corg isotopic values, we analysed our samples after decalcification with 2 N HCL, in order to remove all the inorganic carbon present principally as calcite or dolomite. This procedure was also shown to be suitable for analysis of ␦15 N in samples with low nitrogen content (through increase in nitrogen concentration), despite the possibility of nitrogen loss due to solubilisation of ammonium adsorbed on clay lattices (Meisel and Struck, 2011). Total Organic Carbon (TOC) analyses were performed at the Wolfson Laboratories of University College London. Depending on the likely organic carbon content, an amount ranging between 150 and 350 mg was weighed and then attacked in ceramic crucibles by 3 N HCL until total elimination of carbonate minerals. Afterwards, the residuals were burned in a LECO CS 200 by microwave at 1000 ◦ C in pure oxygen catalysed by iron grains, oxidizing the residue to carbon dioxide. Afterwards, all the gas produced after combustion was measured to assess TOC. For rock samples with nitrogen contents lower than 0.012 mg, isotopic values were obtained after sample decalcification by 2 N HCL (see Table 2). To guarantee the complete removal of carbonates, the samples have been left to react overnight and tested with additional acid before rinsing. Several lines of evidence testify to a different extent of ␦15 N variation before and after this treatment, although mostly regarding recent organic matter (Meisel and Struck, 2011; Scholten, 1991; Bunn et al., 1995; Carabel et al., 2006; Ng et al., 2007), which is dependent on the acid concentration used (Kennedy et al., 2005). For this reason, this treatment should generally be avoided for nitrogen isotopic investigations. In our specific case, we used this technique only for filling gaps in our sequence, whereby good continuity between acid-leached sample values and both overlying and underlying sample values were observed.

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Table 2 Samples with nitrogen concentrations below 0.012 mg/sample have been used only for calculating the C/N ratio, while the ␦15 N from the same samples has not been considered because low accuracy. In some cases the isotopic value after decalcification has been chosen (* samples). Sample

Meters

␦13 Corg

Yuanshan Fm. XTY 61 XTY 60 XTY 59 XTY 58 XTY 57 XTY 56 XTY 55 XTY 54 XTY 53 XTY 52 XTY 51 XTY 50 XTY 49 XTY 48 XTY 47 XTY 46 XTY 45 XTY 44 XTY 43 XTY 42 XTY 41 XTY 40 XTY 39 XTY 38 XTY 37 XTY 36 XTY 35

611.6 591.6 576.6 561.6 550.6 540.6 531.6 526.6 521.6 516.6 511.6 509.6 507.8 505.8 502.8 501.9 500.4 499.4 498.1 495.1 491.1 489.5 487.5 486.5 485.8 485.2 484.2

−25.07 −31.94 −27.75 −26.96 −27.92 −26.43 −26.82 −27.26 −25.95 −28.94 −29.05 −26.24 −27.58 −28.56 −28.99 −28.32 −28.40 −28.54 −28.23 −27.59 −27.95 −27.21 −27.30 −27.44 −27.75 −27.58

2.18 2.28 2.34 2.15 1.86 1.93 1.99 1.86 1.76 1.80

0.065 0.063 0.076 0.065 0.050 0.078 0.075 0.079 0.082 0.088

Shiyantou Fm. XTY 34 XTY 33 XTY 32 XTY 31 XTY 30 XTY 29 XTY 28 XTY 27 XTY 26 XTY 25 XTY 24 XTY 23 XTY 22 XTY 21 XTY 20 XTY 19 XTY 18 XTY 17 XTY 16 XTY 15 XTY 14 XTY 13 XTY 12 XTY 11 XTY 10 XTY 09 XTY 08 XTY 07 XTY 06 XTY 05 XTY 04 XTY 03 XTY 02 XTY 01 CX-XTMS 14 CX-XTMS 13 CX-XTMS 12 CX-XTMS 11 CX-XTMS 10 CX-XTMS 9 CX-XTMS 8 CX-XTMS 7 CX-XTMS 6 CX-XTMS 5 CX-XTMS 4

483.2 473.2 463.2 453.2 443.2 441.7 440.7 439.7 438.7 437.7 436.7 435.7 434.7 433.7 432.7 431.7 430.7 429.7 428.7 427.2 426.2 425.2 424.2 423.2 422.2 421.2 416.7 416.2 415.2 414.2 408.2 404.7 403.4 402.4 389.4 379.4 369.4 359.4 349.4 339.4 329.4 319.4 309.4 299.4 289.4

−28.84 −29.53 −30.44 −30.48 −29.65 −30.01 −29.94 −30.38 −29.78 −29.44 −29.58 −29.76 −30.10 −30.18 −30.28 −30.14 −30.02 −29.68 −30.45 −30.18 −30.07 −30.68 −31.08 −31.37 −31.62 −31.65 −30.52 −31.11 −31.20 −31.99 −31.28 −29.33 −28.83 −30.03 −28.76 −26.96 −28.00 −29.90 −28.83 −30.60 −31.14 −29.55 −29.49 −29.69 −28.96

1.23 1.48 1.36 1.08 1.51 1.69 0.58 0.90 0.85 0.63 0.96 1.01 1.21 0.50 1.13 1.14 0.67 0.28 0.49 0.72 0.70 0.72 0.46 0.66 0.54 1.26 0.87 0.79 0.84 0.81 0.89 0.81 0.30 0.27 1.07 1.40 0.44 1.25 1.21 0.23 1.49 1.73 1.78 2.27 2.30

0.020 0.025 0.022 0.021 0.028 0.032 0.040 0.038 0.046 0.042 0.037 0.044 0.044 0.044 0.043 0.042 0.038 0.036 0.041 0.038 0.039 0.044 0.046 0.043 0.037 0.048 0.033 0.026 0.035 0.038 0.037 0.031 0.035 0.044 0.020 0.017 0.023 0.022 0.024 0.023 0.023 0.026 0.027 0.035 0.028

␦15 N 2.43 2.65 2.54 2.70 2.70 2.62 2.71 2.71 2.53 2.75 2.76 2.45 2.66 1.77 1.96

13 C

TN (%)

TOC (%)

C/N (mol)

0.042 0.060 0.050 0.062 0.063 0.050 0.062 0.068 0.029 0.064 0.064 0.019 0.070 0.057 0.062

0.125 0.168 0.146 0.209 0.197 0.247 0.214 0.284 0.335 0.321 0.299 0.135 0.307 0.473

2.586 2.385 2.512 2.905 2.677 4.238 2.970 3.563 9.883 4.302 4.038 6.159 3.767 7.174

0.159 0.722 0.747 0.713 1.001 0.843 0.522 1.864 3.157 3.324 4.002 4.744

9.942 9.718 11.344 11.166 9.005 20.670 35.994 36.075 42.224 46.608

0.129 0.280 0.196 0.365 0.469 2.255 1.537 2.247 2.031 2.158 1.917 1.860 2.157 2.060

4.404 11.140 7.891 11.101 12.792 47.994 34.705 42.089 41.924 49.734 37.706 36.732 42.185 41.237

1.874 2.174 1.866 1.787 1.882 2.059 1.985 1.285 1.553 1.476 1.005 1.572 1.415 1.647 0.742 0.748 2.304 0.176 0.211

44.278 45.433 41.806 39.675 36.814 38.329 39.544 30.192 28.060 38.768 32.819 38.544 32.414 38.194 20.522 18.543 44.655 7.421 10.481

0.029 0.133 0.302 0.083

1.102 4.780 11.361 3.134

0.131 0.252 0.175

4.161 6.122 5.308

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Table 2 (Continued) ␦13 Corg

Sample

Meters

CX-XTMS 3 CX-XTMS 2 CX-XTMS 1 XT 29 XT 28 XT 27 XT 26 XT 25 XT 24 XT 23 XT 22 XT 21 XT 20 XT 19 XT 18 XT 17 XT 16 XT 15 XT14 XT13 XT12 XT11 XT10 XT9 XT 8 XT 7 XT 6 XT5 XT 4 XT 3 XT2 XT1

279.4 269.4 259.4 254.4 251.9 249.9 248.9 247.9 245.9 244.2 243.2 242.2 240.7 239.2 238.2 236.4 234.4 233.4 232.4 231.4 229.9 228.5 227.5 226.0 225.0 224.0 223.0 221.5 220.2 219.2 218.2 217.9

−29.98 −29.57 −29.93 −29.74 −29.25

Dahai Mb. XT 101 XT103 XT105 XT107 XT 109 XT 111 XT 113 XT 115 XT 117 XT 119 XT 121 XT 123 XT 125 XT 127 XT 129 XT 131 XT 133 XT 135 XT 137 XT 139 XT 145 XT 147 XT 149 XT 151 XT 153 XT 155 XT 156

217.6 216.6 215.7 214.1 211.5 209.0 207.2 203.8 200.4 197.9 195.7 193.5 191.2 188.1 185.2 182.6 180.0 176.8 174.4 172.3 163.2 160.2 157.7 155.2 153.8 152.2 150.8

−27.56 −29.53 −30.26 −29.78 −24.31 −21.70 −21.46 −24.02 −23.69 −23.74 −22.90 −22.33 −23.69 −22.15 −22.26

Zhongyicun Mb. XT 157 XT 159 XT 161 XT 162 XT 163 XT 164 XT 165 XT 166 XT 168 XT 170 XT 172 XT 174 XT 175 XT 176

147.7 146.6 144.3 143.9 142.6 141.6 140.9 139.9 137.9 136.6 134.6 132.9 131.9 130.9

−25.39 −26.19 −24.42 −26.57 −30.26 −29.95 −26.41 −29.26 −28.23

−35.77 −31.72 −31.66 −31.49 −31.33 −34.29 −34.52 −34.39 −34.21 −33.93 −33.79 −33.92 −33.43 −33.89 −33.87 −33.87 −33.67 −32.87 −28.68 −31.84 −29.95 −27.87 −27.78 −28.52 −28.61 −26.38

−21.53 −21.53 −22.15 −22.55 −23.06 −22.51 −22.81 −24.35 −25.49 −24.41 −25.00

−30.37 −30.54 −28.31 −32.03

␦15 N 2.16 2.17 2.12 2.35 2.54 2.34 2.20 2.18 2.02 2.32 1.71 2.25 1.86 2.11 1.88 1.82 2.41 2.68 3.14 2.18 2.23 2.78 2.61 2.06 2.19 2.50 2.42 2.24 2.44 2.46 2.70 2.61

1.30 −0.03 0.01 0.67 0.94 0.94 1.17 2.32 2.27 2.83 2.95 2.66 3.00 1.85 2.05 2.03 4.86 6.25 7.67 6.08

4.93 4.52 2.74 −0.91

1.19 3.17 5.01 5.28 4.47 4.00 6.07 5.49 6.50 8.80 5.62 3.33

13 C

24.18 28.19 27.03 23.44 28.59 30.15 29.54 28.02 27.23 27.64 29.70 30.91 30.70 29.59 28.59 30.35 28.83 29.36 28.00 28.67 29.08 28.93 29.58 28.81 27.43 26.49 26.54 25.77 25.92 26.44 25.60 24.44 20.53

19.97 25.34

28.65 28.48

TN (%)

TOC (%)

C/N (mol)

0.030 0.020 0.050 0.047 0.045 0.056 0.060 0.054 0.057 0.055 0.051 0.059 0.060 0.052 0.058 0.049 0.069 0.073 0.049 0.049 0.057 0.065 0.059 0.062 0.050 0.055 0.033 0.043 0.046 0.047 0.064 0.011

0.303 0.207 0.732 0.932 1.020

8.574 8.710 12.579 17.063 19.413

2.552 2.875 3.360 3.411 2.857 3.816 4.007 2.554 2.952 2.312 3.666 3.328 2.186 2.099 3.172 4.150 3.616 4.256 3.794 2.677 1.943 1.982 2.025 2.187 5.023 0.129

36.875 45.776 50.310 53.189 47.753 55.815 57.770 42.177 43.409 40.397 45.863 39.366 38.630 36.566 47.504 54.497 52.527 58.641 65.197 41.519 51.299 39.838 37.551 39.980 67.795 9.653

0.004 0.014 0.011 0.012 0.101* 0.341* 0.04–0.220* 0.079* 0.127* 0.003–0.149* 0.120* 0.091* 0.005–0.117* 0.121* 0.131* 0.164* 0.004–0.138* 0.122* 0.010 0.006 0.005

0.112 1.760 1.424

143.554 101.111

0.167

0.653

0.051

0.294

0.096

0.706

0.169

1.053

0.074

6.644

0.092

15.642

0.003 0.006 0.006 0.033 0.008

0.098

0.199

5.239

0.083 0.031

0.170

1.757

0.011 0.011 0.025 0.074 0.016 0.009 0.011 0.012 0.017 0.059

0.693

53.147

0.094

1.095

0.310

25.058

2.558

37.543

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157

Table 2 (Continued) Sample

Meters

XT 177 XT 180 XT 181 XT 182 XT 183 XT 184 XT 186 XT 188 XT 189 XT 190 XT 192 XT 193 XT 196 XT 198 XT 199 XT 201 XT 204 XT 206 XT 207 XT 208 XT 210 XT 212 XT 214 XT 216 XT 218 XT 219 XT 221 XT 223 XT 226 XT 228 XT 229 XT 232 XT 234 XT 236 XT 237 XT 238

130.1 127.1 126.3 124.8 122.8

␦13 Corg

␦15 N

120.8 117.8 116.8 115.8 113.8 112.8 109.8 107.3 106.3 104.3 103.7 102.1 101.1 100.9 99.5 97.5 95.5 93.7 90.2 89.6 88.9 87.9 84.3 84.3 83.7 75.2 73.9 72.6 67.6 41.6

−31.60 −30.90 −31.69 −32.87 −31.94 −33.20 −30.93 −33.00 −33.97 −33.76 −33.88 −33.86 −33.51 −34.20 −33.42 −33.74 −33.22 −30.52 −32.31 −27.31 −32.34 −32.12 −32.68 −32.06 −29.24 −33.55 −33.11 −32.63 −30.51 −31.73 −30.48 −32.04 −28.37 −32.26 −33.25 −33.48

Daibu Mb. XT 240 XT 242 XT 246 XT 247 XT 248 XT 250 XT 252 XT 253 XT 255 XT 258 XT 260 XT 262 XT 263

39.1 37.8 34.1 33.1 31.5 29.0 25.5 23.5 20.5 16.3 14.3 12.9 12.4

−31.10 −32.26 −33.17 −32.76 −32.14 −34.76 −35.23 −33.65 −31.58 −27.42 −34.86 −30.37 −32.84

1.81 2.37

Dengying Fm. XT 264 XT 266 XT 268 XT 270 XT 273 XT 276 XT 278

12.3 11.0 8.8 6.8 5.0 2.9 0.0

−29.54 −23.32 −25.37 −23.76 −26.11 −24.88 −21.10

4.80 4.42 4.46 2.67

*

2.26 1.66 4.09

2.37 1.13 2.93 3.00 2.14 2.29 5.07 4.63 4.82 4.64 5.13 3.86 4.81 3.77 4.31 3.34 3.88 3.78 3.38 2.90 4.24 4.43

13 C

TN (%)

TOC (%)

C/N (mol)

27.72

0.008*

0.399

28.01

0.013 0.031

4.152

113.429

1.044

27.784

7.567

52.740

0.031

0.610

0.064

0.161

0.987

29.100

1.064

39.477

1.179

32.012

0.027

0.699

0.016

0.060

3.303

0.014 0.007*

0.003

0.190

24.39 25.41 25.63 31.94 28.48 32.48 28.80 29.12 30.62 29.36 28.90

19.58 27.28 28.27

0.036

0.025 0.032 0.058 0.056 0.123 0.182 0.180 0.044 0.154 0.030 0.342 0.021 0.055 0.029 0.039 0.027 0.023 0.029 0.032 0.027 0.046 0.033

27.38 23.10 4.86 27.98

5.56 4.81

26.86 27.08 25.90 26.47 24.91 27.25 27.11 22.91 19.47 27.02

0.059

0.016 0.011* 0.006*

0.329

25.722

0.012

0.479

27.03 26.92 21.77 24.57 23.95 27.20 25.53 21.78

0.013* 0.019* 0.017* 0.021*

0.416

Value obtained from decalcified samples.

5. Results 5.1. The ı13 Corg profile The organic carbon isotopic profile is shown in Fig. 4a. Its trend may be subdivided into two parts: the first one, from the top Dengying up to the basal Shiyantou exhibits four large fluctuations: C1 and C4 (negative shifts) and C2 and C5 (positive shifts). The first three excursions show considerable variation of up to 10‰, while the last one (C5) is of more minor extent. By contrast, we observe two areas of relative stability in the Daibu and Zhongyicun members (excluding the silty layers) and through almost all of the Dahai

mb. (C3). The upper part, comprising the upper 380 meters shows instead a smooth and generally rising trend (C6-C7) from −30‰ to −26‰. The overall isotopic range through this section is between −21‰ and −36‰. Isotopic investigations on samples from the last 800 Ma show a range of ␦13 Corg around 28‰ to 32‰ lower than contemporaneous DIC (Hayes et al., 1999). Large differences from these values might be due to enhanced burial of 13 C-depleted chemoautotrophic biomass (Hayes et al., 1999) or lower isotopic fractionation during carbon fixation due to lower atmospheric pCO2 , as perhaps was likely during glacial periods (Des Marais et al., 1992; Hayes et al., 1999). Original ␦13 Corg values can become altered because

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Fig. 4. Isotopic profiles: (a) ␦13 Corg profile; (b) ␦15 N profile. Hatched lines represent data from decalcified samples. Colors show the prevalent biochemical processes probably defining the nitrogen isotopic signature trough the section; (c) ␦13 Ccarb profile (from Zhou et al., 1997).

of sample maturation. Although investigation into the H/C ratio of our samples was not performed, the high degree of similarity between the Ccarb and Corg profiles suggests, as a first approximation, the absence of important secondary processes on organic matter isotopic composition. In the same figure (c) the Ccarb isotopic profile from Zhou et al. (1997) is also shown for comparison. The similarity in trend and nick points of both isotopic curves is remarkable. The low resolution of the carbonate isotopic profile is probably due to difficulties in finding suitable samples in all parts of the section (especially in silty and sandy domains), resulting in lower resolution, especially in the upper 300 m. 5.2. The ı15 N profile ␦15 N results are plotted in Fig. 4b. The ␦15 N data from Xiaotan section show fluctuations between −1‰ and +9‰. This range mirrors values that are currently measured in diverse ecosystems on the Earth (Altabet and Francois, 1994; Sigman et al., 2000). Comparison with nitrogen isotope data from ancient rocks also reveals a significant accordance (Thomazo et al., 2009; Godfrey and Falkowski, 2009; Beaumont and Robert, 1999).This profile shows a different cyclicity from the carbon isotope curve, registering several abrupt isotopic excursions as well as scattered values in low nitrogen content samples (see N1-N3). As for the Corg profile, here again the total curve can be divided into two parts: the first one, from the lowermost part of the section ending at the base of the Shiyantou Fm. (from excursion N1 to N7), characterized by high data instability and several meaningful shifts, and the upper part characterized by a flat and relatively smooth increase (2–3‰) around the Shiyantou-Yuanshan transition (N8-N9). The overall range of nitrogen isotope values is 10‰, with its low point at the base of the Dahai Mb. (−1‰) and the high point in the Zhongyicun Mb (+9‰). Hatched segments in the nitrogen curve represent values obtained from decalcified samples aimed at concentrating the nitrogen-bound fraction (organic or clay). This was carried out in

some cases due to the very low nitrogen contents of some samples (principally pure carbonates and cherts) that otherwise would have been impossible to analyse (TN < 0.012 mg). 5.3. TN/TOC vs. ı15 N Absence of systematic correlation of isotopic nitrogen signature with TOC and TN in all the formations and members denotes a mutual independence (Fig. 5a and b), ruling out preferential loss of heavy/light nitrogen after deposition or any other case of decoupling with biochemical processes. Observing the modern ocean, high ␦15 N values in the PON characterize areas where denitrification is the major biochemical process occurring in seawater, responsible for loss of nitrogen in the marine system. A natural outcome would be decreased primary productivity resulting in minor organic content (and consequently carbon and nitrogen) in the sediments. This might explain a generally higher ␦15 N for low nitrogen-content samples. Isotope values for samples with TN < 0.01% and TOC < 0.5% show greater variability, which could be assigned to a decrease in data accuracy, but this possibility may reasonably be ruled out in our study where data show good continuity of trend. 5.4. TOC vs. TN High and low carbon content samples from different geological units have been plotted against total nitrogen and total organic carbon (Fig. 5c), showing in some cases a linear relationship (Zhongyicun Mb.) and always an intercept on the N axis. This is caused by the mixing of two different nitrogen sources: the organic remains, whereby the C/N ratio is broadly constant and inorganic clay-bound nitrogen devoid of carbon. Moreover, the lower is the curve regression coefficient, the higher is the variability in dilution between the two phases due to carbonates or silicate debris (Calvert, 2004). Almost identical co-variations of nitrogen content with TOC are shown for all the intercepts, testifying to a similar

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Fig. 5. Cross-plot showing: (a) ␦15 N vs. TN; (b) ␦15 N vs. TOC; (c) TN vs. TOC; (d). C/N vs. TOC Samples with values not falling within the diagram’s area have been neglected.

organic matter structure in terms of proportion of nitrogen in the biomolecules (Dahai Mb., Basal Shiyantou, upper Shiyantou and Yuanshan Fm.). These trends might be ascribed to a similar effect from early diagenetic microbial digestion on the POM reaching the sea bottom, and/or to a similar source for the biological remains. Data from the Zhongyicun Mb show no meaningful patterns, probably due to the several different lithologies of which it is composed. Despite having a similar chemical composition, the nitrogen content contained in clay minerals is relatively high in these samples (within a range of 50% to 90% of the bulk N) with the highest proportion in the Yuanshan Mb. This feature might be due to different clay contents or enhanced thermal maturity achieved during diagenesis (although seemingly unlikely to vary within a single section), and a more complete transfer of organic-bound nitrogen to the clay layered structures. The Dahai samples are characterized by low TOC and low silty content, and so exhibit relatively low nitrogen contents. 5.5. TOC vs. C/N In Fig. 5d we report TOC against the ratio Corg/Ntot. As already shown in previous studies (Bristow et al., 2009; Calvert, 2004 and references therein), there is a conspicuous increase of this ratio with organic content. This trend has been claimed to relate to the thermal maturity grade of the sediment. In our opinion, this general trend has mostly to do with early diagenesis. By contrast to the situation shown here, an increased maturation temperature would transfer an increased amount of N from the organic fraction to the silicates, but the bulk nitrogen would not generally be affected as the system is commonly closed. One possible explanation might be varied dilution of the POM with other minerals, whereby samples with higher TOC could reflect simply a more minor contribution from other particles settling through the water column, allowing the bio-nitrogen to be more affected by microbial degradation. Differences in slope intercepts could be due to the different nature of biological contributions or redox conditions (Lehmann et al., 2002). For the intercept of the Dahai samples, such a large increase in the C/N ratio might be ascribed to its low clay content and high

carbonate content, which would have limited the incorporation of nitrogen into silicate mineral lattices that might otherwise have been digested by microbes. 6. Discussion The radiation of metazoans and related biological innovations throughout the entire living realm led to numerous changes in biogeochemical cycling (Logan et al., 1995; Dornbos, 2006; Fike et al., 2006; Shen et al., 2008). As a consequence, increased instability of isotopic signatures is often registered in sediments formed during the Paleozoic and Phanerozoic, although no clear explanations have been accepted as yet. Isotopic fluctuations in the ␦13 Corg trend demonstrate communication between DIC and DOC pools at the Precambrian/Cambrian transition in South China (Fig. 6a and c). Similarity between Ccarb and Corg isotopic profiles is ascribed to direct interactions between the organic and inorganic carbon pools, demonstrating that these data truly reflect primary marine organic matter, seawater chemistry and close coupling, unlike some previous geological intervals probably characterized by a very large DOC pool (Rothman et al., 2003; Fike et al., 2006; Swanson-Hysell et al., 2010). A completely different trend is shown by the nitrogen isotopic curve (Fig. 6b) which also registers more frequent fluctuations through the investigated section when compared to carbon isotopes. This arises from the fact that the two proxies respond differently to the biogeochemical settings, being sensitive to diverse physical and chemical processes occurring in the ocean. Differences in frequency are possibly due to dimensions of these isotope pools, intensities of chemical reactions and their associated fractionation factor. Observation of isotopic Corg, N, Ccarb, and TOC profiles (respectively a, b, c, and e in Fig. 6) permit us to constrain the biogeochemical setting characterizing the basin. 6.1. The Dengying-Daibu transition The negative carbon excursion at the end of the Proterozoic (C1 in Fig. 4) has been recorded globally although interpreted in various

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Fig. 6. Collection of all profiles. In the diagram (d) the %N axis is logarithmic in scale.

ways, some of which are related to the advent and/or expansion of new forms of metazoan life. The introduction of faecal pellets may have played an important role here (Logan et al., 1995; Ishikawa et al., 2007), but they did not become abundant until the beginning of the Cambrian Period (Babcock et al., 2001), equivalent to the SSFA 1-2 (Steiner et al., 2007). The bundling of organic matter in faecal packages may be able to explain the weak increase in ␦15 N at this level as faecal pellets would normally be enriched in 15 N with respect to the food source (Montoya, 1994). Other important marine processes invoked involve anaerobic recycling of sedimentary OM (Guo et al., 2007; Goldberg et al., 2007) and oxidation of a large anoxic DOC mass (Wille et al., 2008; Rothman et al., 2003; Ishikawa et al., this volume). Another hypothesis sees the release of very low ␦13 C into the water, due to methane clathrates release or hydrothermal venting (Chen et al., 2009; Kirschvink and Raub, 2003). The slight nitrogen isotope increase observed in this study could reflect enhanced denitrification after a rapid increase in the reduced carbon content of the water column (larger OMZ; Ganeshram et al., 2000; LaPorte et al., 2009), leading to a rise of the chemocline, reduction of the nitrate pool turning then more sensitive to isotopic changes, or simply to the passage of the oxic-anoxic boundary from the sediments to the water column. Unfortunately, given the very low OM concentrations for many samples belonging to the Dengying Fm. and Daibu Mb., their N content was generally too low for isotopic analysis. Therefore, from the few values observed in Fig. 4 ranging between +2‰ and +6‰ we may infer conditions of relatively normal marine productivity (N1, Fig. 4). 6.2. The Zhongycun and Dahai members With the beginning of the Zhongyicun Mb. and the FAD of SSF’s, both carbon isotope profiles show an increase in the second part of the unit (C2), although several positive spikes register some instability already in the first part. Phosphorites define the onset of this member and possibly testify to the beginning of a new stage with enhanced primary productivity and starved sediment accumulation (Karl et al., 1997; Algeo et al., 2007). The positive shift in the ␦13 Ccarb curve can most parsimoniously be interpreted in terms of prolonged deposition of 12 C-enriched organic matter in marine sediments causing a natural increase in the marine carbonate reservoir ␦13 C. The nitrogen profile presents an important positive shift at the base of the unit (N2-N2a), one negative shift in the middle (N3) and one extreme excursion up to +9‰ at the very

top (N4), occurring in phosphorite-rich lithologies. Phosphorus is mostly contained in sedimentary rocks as authigenic carbonate fluoroapatite (Trappe, 1998; Föllmi, 1996) formed after its delivery to the sediment bound to organic matter. While most P is released from the sediment to the seawater during degradation of the OM (Ingall et al., 2005; Goldhammer et al., 2010), in the case of oxic bottom waters a greater proportion of remineralized P can be adsorbed and complexed on Fe-oxyhydroxides (Slomp et al., 1996) or sequestered by biological activity (Davelaar, 1993; Sannigrahi and Ingall, 2005). The best conditions for phosphogenesis arise in cases of high P delivery (benthic oxygen deficiency), suboxic settings in bottom waters aiding its retention in the sediment (within the NO3 − reducing zone), and frequent variations in water redox conditions (Jarvis et al., 1994). Accordingly, the correlation between nitrogen isotopic fluctuations and phosphorite deposition (N2-N4) can be attributed to the establishment of intermittent denitrifying conditions through the base and top Zhongyicun after oxygen depletion due to degradation of increased sinking POM, leading to higher denitrification rates and probably significant depletion of the bioavailable N during the thickest phosphate deposition in its upper part (Kuypers et al., 2005, 2003; Lam et al., 2007). Widespread deposition of phosphorites or carbonate fluoroapatite in these layers testifies to sustained high concentrations in seawater. Development of primary productivity with nitrogen-limited conditions at the base of the Cambrian would then reasonably support episodes of nitrate depletion and N2 -fixation, and consequently isotopic response to transfers of nitrogen among pools (principally nitrate, ammonia and organic-bound). Increase in ␦15 N at the base of the Cambrian might be also driven by the introduction of faecal pellets (Logan et al., 1995), which are characterized by higher ␦15 N (Altabet and Small, 1990; Montoya, 1994; Checkley and Miller, 1989). This important breakthrough in export productivity could have destabilized the nitrogen cycle causing abrupt fluctuations before achievement of a new equilibrium, or shift in the average sediment nitrogen isotope signature toward higher values in the Phanerozoic. The small but distinct increase in the nitrogen isotopic signature labeled N2a in Fig. 6b reveals interesting features. This layer records high silt content and total nitrogen concentrations but low organic carbon content. The most representative sample is XT208 (see table II for details), with a nitrogen concentration of 0.342%, while the TOC is as low as 0.064% showing that the nitrogen must be almost entirely clay-bound. A similar example is XT 240 in the Daibu Mb. The associated isotopic trend may reflect an increase in denitrification rates in the case of authigenic clays or may simply

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be a result of exotic nitrogen supply retaining their original values and thus not representative of marine processes. In the last 20 m at the top of Zhongyicun Mb., nitrogen isotopic values drop 10‰ to −1‰ right at the boundary with the Dahai Mb. (N5), followed by an equally abrupt shift bringing values back toward high denitrification rates in the following 20 m (N6). These two positive peaks (N4-N6), separated by a negative spike (N5), can find a possible interpretation when coupled with the carbonate and organic carbon profiles. In this nearly 50 m, ␦13 Ccarb and ␦13 Corg values both increase significantly (C2), presumably in response to a global scavenging of light carbon from the hydrosphere and consequently from the atmosphere. This intense biological pump could have caused oxygen depletion in the bottom waters, slowly encroaching on wider sections of the water column until the nitrate concentration became too low to sustain a normal N-assimilation. In this situation denitrification and nitrogen assimilation would be weaker, conferring the ecological advantage to N2 -fixation in the upper water column. (Anbar and Knoll, 2002; Tyrrell, 1999; Karl et al., 1997; Montoya, 1994). Nitrogen fixation by prokaryotes requires high activation energy, explaining why it occurs predominantly only in circumstances where other forms of nitrogen assimilation are disadvantageous (Karl et al., 2002; Mulholland et al., 1999). Based on the particularly low nitrogen signature reached in excursion labeled N5, we can also infer activity of anoxic photoautrotophs (green and purple sulfur bacteria; Meyers et al., 2009; Johnston et al., 2009; Ohkouchi et al., 2005) characterizing events of complete oxygen depletion in the water column and presence of sulfides in the photic zone. With the beginning of limestone deposition in the Dahai Mb., the Corg stops its rise and nitrogen isotopes invert their trend indicating renewed oxygenated conditions supporting a normal marine productivity in a denitrification-dominated water column (N6). This scenario could be explained as the attainment of a new equilibrium state between the carbon assimilation and remineralization that would have renewed suboxic/oxic conditions in the seawater, thus favoring a “normal marine production”. The massive enrichment in 13 C observed in the top ZhongycunDahai Mb. (C2-C3-C4) occurred during a period of primary productivity characterized by N-limitation. This is assumed on the one hand by repeated phosphogenic events (Papineau et al., 2009), and on the other hand by the nitrogen isotope trend. Several isotopic peaks at that time in fact testify to episodes of high denitrification rates lowering the N:P ratio until N2 -fixation became probably dominant and able to restore the Redfield ratio (Saltzman, 2005). Before this happened, N-limitation conditions were set, limiting primary productivity and avoiding organic-rich layers all through the Dahai Mb. If positive carbon excursions are characterized by P-limited oceans as proposed by Saltzman (Saltzman, 2005; Falkowski, 1997), based on the carbon profile from the Xiaotan section the time elapsed before complete nitrogen utilization still permits enhanced organic production and rise of the carbon isotopic proxy. The Saltzman hypothesis remains valid here to the extent that when the nitrogen pool appears to have been exhausted, ␦13 C values cease their rise, showing a perfect correspondence in time between the two isotopic cycles (beginning of C3 stage and N5 in Fig. 6). The ability of N-limited waters to curb carbon isotopic excursions through negative feedbacks requires further testing against other geochemical conditions (e.g. dimension of N pools, intensity of nitrogen loss, concomitant sources of renewed nitrogen, etc.).

(volcanic carbon dioxide) or a protracted and widespread remineralization of the OM pool could help to explain these negative shifts and might have caused warming due to the resultant enhanced greenhouse effect from released carbon dioxide. Accordingly, the following 150 m of silty-sandy lithologies testify to changes in the sediment supply to the basin, which may have been driven by increased weathering, induced possibly by climate- and/or tectonic factors. Increased temperatures enable enhanced water stratification which can lead to domination of the N-isotope cycle by N2 -fixation (Capone et al., 1997; Struck et al., 2001; Karl et al., 2002). This is in accordance with the low nitrogen values that characterize the Shiyantou and the Yuanshan Fm. Additionally, an extended continental nutrient influx marking the beginning of the Shiyantou Fm. can be envisaged. At this stage, an increase in the POM rain exhausting the oxygen pool might be responsible for the deposition of the black-shale lithology in anoxic deep waters, again favoring strong N2 -fixation and diminished water mixing (cf. Struck et al., 2009), with possibility of GSB and PSB activity (N7 in Fig. 4b). About 30 meters above, protracted climate-driven ventilation accompanied by changes in oceanic currents would be responsible for the drastic contraction of the OMZ pool and the end of black shale deposition. In Summary, the increase of ␦15 N at the base of the Shiyantou Fm. could be the result of a new equilibrium state, whereby although nitrogen fixation remains dominant, other processes like nitrate assimilation and denitrification play an important role, which is indicative of weaker stratification and a larger nitrate pool. At a general glance, isotopic trends for both nitrogen and carbon are smoother in this portion of the section, and are affected only by relatively minor perturbations (C5-C6-C7-N8-N9). We observe a small increase in both isotopic parameters, starting in the upper Shiyantou black shales and lasting until the end of the siltier units in the basal Yuanshan Formation (C7-N9). The nitrogen signal could instead record a slow recovery from ocean stratification and dominant N2 -fixation to denitrifying conditions, also testified to by repeated carbonate beds intercalated in the Yuanshan Fm. possibly defining an interval of decreased productivity. According to Steiner et al. (2007) and Wang and Mo (1985), North Yunnan paleobasinal settings represent peculiarities among Yangtze platform successions. Apparently, this area is placed in between two shallower platforms, respectively North and South representing sedimentary highs in the area. This geographic shape accompanied by constant subsidence would have created a basin deeper than its surroundings and for this reason characterized by high supply of detritus. Conditions favorable for high sedimentation rates are already testified to by the thicknesses of the Dengying Fm. in this geographical area (Steiner et al., 2007). The Xiaotan section represents conditions of relatively high sedimentation rates, but its shallow water setting and the existence of several erosional surfaces and phosphoritic hard ground (field observations) shed doubt on its continuity. As, nitrogen isotopic fluctuations respond to very short-term oceanographic changes, our record may be affected by loss of information due to insufficient sampling resolution. The important and well defined fluctuations recorded from this section support, however, our sampling strategy and serve to confirm a relatively high sediment supply to the sea floor. Additionally, our data depict a vigorous nitrogen cycle especially during deposition of the Zhongyicun Mb. and Dahai Fm., featured by dominance of different biochemical processes causing imbalances to “normal marine production” settings.

6.3. The Shiyantou-Yuanshan formations

6.4. The importance of enzyme co-factors and bioturbation

Carbon isotope trends in the upper Dahai Member and at the base of the Zhongyicun Member show a significant, steep negative shift of 15‰ (C4). Additional light carbon from geological sources

After complete oxygenation of bottom waters at the end of the Neoproterozoic, the sulphidic conditions responsible for Mo and Fe scavenging seem likely to have diminished (Anbar and Knoll,

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2002), except for in South China where there is some evidence for sulphidic conditions in the basin during the Cambrian (Canfield et al., 2008; Och et al., this volume). The concentrations of these two essential nitrogenase enzyme co-factors for the N2 -fixation operated by prokaryotes would have increased in seawater, affecting profoundly both primary productivity and eukaryotic evolution, as well as increasing OM transfer to bottom waters through faecal pellet delivery and benthic suspension feeding, crucial developments for the oxidative state of the entire oceanic pool. In our opinion, these biogeochemical breakthroughs might be strongly connected to the metazoan diversification observed in the Cambrian since N2 -fixer autotrophic cyanobacterial activity represents the primary biological pump transferring elements from the inorganic to the organic pool, and producing bioavailable compounds at the base of the food chain. Significant nitrogen isotopic variations at the beginning of the Cambrian Period outlined here are correlative with major variations in oxygen concentration through the water column, confirming the early Cambrian as a geological period affected by fundamental steps toward an oxygenated Earth (Canfield et al., 2008; Scott et al., 2008). If bottom water conditions become oxygenated, aerobic early diagenesis of organic matter will occur, which is enhanced by the activities of metazoan bioturbators. Another crucial ecological innovation across the Pc-C boundary is a deeper and more intense bioturbation and evolution of bilaterian lifestyles as testified to by the evolution of body and trace fossils (Weber et al., 2007; Dornbos, 2006; Lin et al., 2010). A continuous enhanced remobilization of sediments probably caused renewed oxidative conditions (redox fluctuations) during early diagenesis in oxic bottom water domains favoring increasing habitability of sediments by O2 diffusion (McIlroy and Logan, 1999), together with loss of short-term isotopic signals due to vertical particle homogenization. Moreover, episodes of nitrogen loss by microbial digestion, denitrification (Grundmanis and Murray, 1977; Gilbert et al., 1995, 1997, 1998) and recycling to the seawater might decrease the organic fraction retained in the sediment, but at the same time represent an important nitrogen source for the marine ecosystem balancing the influx of new nitrogen from the atmosphere. Although these assumptions are plausible, recent studies on the effects of biogenic burrows in western coasts of North America showed important N2 -fixation activity in bioturbated sediment (Bertics et al., 2010), despite enhanced nitrification/denitrification. The organic matter introduced in burrows, in fact, would sustain sulfate reducing bacterial (SRB) activity fixing dinitrogen produced by the biochemical steps ammonia → nitrate → dinitrogen. Consequences for the nitrogen isotopic signature of seawater and sediments are difficult to predict considering the low number of comparative investigations carried out on bioturbated and unbioturbated ecosystems. In the Xiaotan section, bioturbation is apparent from the appearance of SSF’s but becomes increasingly complex and penetrative within the Shiyantou Formation (Tommotian-equivalent) as recorded from time-equivalent successions around the world (Droser, 1987; Dornbos, 2006; Shields-Zhou and Zhu, this volume). This corresponds to a considerable dampening of fluctuations in both isotopic parameters, which may be the result of increasing complexity of early diagenesis due to bioturbation.

7. Conclusions Carbon and nitrogen isotope trends through the Xiaotan section (Yunnan province in South China) were used to reconstruct geological conditions and biological activity on the Yangtze Platform Basin during the early Cambrian. Variations in nitrogen and organic carbon isotope composition can be correlated with changes in the water oxygen concentration, nutrient availability

and physical conditions at regional and global scales. By comparing the ␦15 N, ␦13 Corg and the ␦13 Ccarb records, we tried to determine palaeoenvironmental conditions, and specifically movement of the chemocline between the sedimentary pile and the water column causing widening and contractions of the OMZ. Both isotope systems show significant fluctuation in the Zhujiaqing Fm., testifying to a period of biological and geochemical instability in seawater at the beginning of the Cambrian explosion. The upper part of the section reveals the establishment of relative stability, evidenced by the lack of major isotopic shifts. This study confirms that nitrogen isotope analysis of bulk samples (with some exceptions) as well as decalcified samples can provide important paleoenvironmental information. Nevertheless, the sensitivity and complexity of the nitrogen isotope proxy necessitate study of correlative geological successions around the world before our conclusions can be confirmed. References Algeo, T.J., Ellwood, B., Nguyen, T.K.T., Rowe, H., Maynard, J.B., 2007. The Permian–Triassic boundary at Nhi Tao, Vietnam: evidence for recurrent influx of sulphidic water masses to a shallow-marine carbonate platform. Palaeogeogr. Palaeoclimat. Palaeoecol. 252, 304–327. Altabet, M.A., 2007. Constraints on oceanic N balance/imbalance from sedimentary N-15 records. Biogeosciences 4, 75–86. Altabet, M.A., Higginson, M.J., Murray, D.W., 2002. The effect of millennial-scale changes in Arabian Sea denitrification on atmospheric CO2 . Nature 415, 159–162. Altabet, M.A., Pilskaln, C., Thunell, R., Pride, C., Sigman, D., Chavez, F., Francois, R., 1999. The nitrogen isotope biogeochemistry of sinking particles from the margin of the Eastern North Pacific. Deep-Sea Res. I 46, 655–679. Altabet, M.A., 1996. Nitrogen and carbon isotopic tracers of the source and transformation of particles in the deep-sea. In: Ittekkot, V., Schaffer, P., Honjo, S., Depetris, P.J. (Eds.), Particle Flux in the Ocean. Wiley, London, pp. 155–184. Altabet, M.A., Francois, R., 1994. Sedimentary nitrogen isotopic ratio as a recorder for surface ocean nitrate utilization. Global Biogeochem. Cycl. 8, 103–116. Altabet, M.A., Small, L.F., 1990. Nitrogen isotopic ratios in fecal pellets produced by marine zooplankton. Geochim. Cosmochim. Acta 54, 155–163. Anbar, A.D., Knoll, A.H., 2002. Proterozoic ocean chemistry and evolution: a bioinorganic bridge? Science 297, 1137–1142. Babcock, L.E., Zhang, W., Leslie, S.A., 2001. The Chengjiang Biota: record of the Early Cambrian diversification of life and clues to exceptional preservation of fossils. GSA Today 11, 4–9. Beaumont, V., Robert, F., 1999. Nitrogen isotope ratios of kerogens in Precambrian cherts: a record of the evolution of atmosphere chemistry? Precambrian Res. 96, 63–82. Bertics, V.J., Sohm, J.A., Treude, T., Chow, C.-E.T., Capone, D.G., Fuhrman, J.A., Ziebis, W., 2010. Burrowing deeper into benthic nitrogen cycling: the impact of bioturbation on nitrogen fixation coupled to sulfute reduction. Marine Ecol. Prog. Ser. 409, 1–15. Billings, E., 1871. Proposed new genus of Pteropoda [Hyolithellus]. Canadian Natural. Geol. 6, 240. Bristow, T.F., Kennedy, M.J., Derkowskia, A., Drosera, M.L., Jiang, G., Creaserc, R.A., 2009. Mineralogical constraints on the paleoenvironments of the Ediacaran Doushantuo Formation. PNAS 106 (32), 12190–13195. Bunn, S.E., Loneragan, N.R., Kempster, M.A., 1995. Effects of acid washing on stable isotope ratios of C and N in penaeid shrimp and seagrass: Implications for foodweb studies using multiple stable isotope. Limnol. Oceanogr. 40 (3), 622–625. Canfield, D.E., Glazer, A.N., Falkowski, P.G., 2010. The evolution and future of Earth’s Nitrogen Cycle. Science 330, 192–196, doi:10.1126/science.1186120. Canfield, E., Poulton, S.W., Narbonne, G.M., 2007. Late-Neoproterozoic deep-ocean oxygenation and the rise of animal. Life. Sci. 315, 92–95, doi:10.1126/science.1135013. Calvert, S.E., 2004. Beware intercepts: interpreting compositional ratios in multicomponent sediments and sedimentary rocks. Org. Geochem. 35, 981–987. Capone, D.G., Zehr, J.P., Paerl, H.W., Bergman, B., Carpenter, E.J., 1997. Science 276, 1221. Carabel, S., Godínez-Domínguez, E., Verísimo, P., Fernández, L., Freire, J., 2006. An assessment of sample processing methods for stable isotope analyses of marine food webs. J. Exp. Marine Biol. Ecol. 336, 254–261. Checkley, D.M., Miller, C.A., 1989. Deep Sea Res. 36, 1449. Chen, D., Wang, J., Qing, H., Yan, D., Li, R., 2009. Hydrothermal venting activities in the Early Cambrian, South China: Petrological,geochronological and stable isotopic constraints. Chem. Geol. 258, 168–181. Chen, M., Chen, Y., Qian, Y., 1981. Some tubular fossils from Sinian–Cambrian boundary sequence, Yangtze Gorges. Bulletin of the Tianjin Institute of Geology and Mineral Resources. Chin. Acad. Geol. Sci. 3, 117–124. Cohen, P.A., Knoll, A.H., Kodner, R.B., 2009. Large spinose microfossils in Ediacaran rocks as resting stages of early animals. PNAS 106 (16), 6519–6524, doi:10.1073/pnas.0902322106.

L. Cremonese et al. / Precambrian Research 225 (2013) 148–165 Coveney Jr., R.M., Watney, W.L., Christopher, C.G., 1991. Contrasting depositional models for Pennsylvanian black shale discerned from molybdenum abundances. Geology 19, 147–150, doi:10.1130/0091-7613. Dalsgaard, T., Canfield, D.E., Petersen, J., Thamdrup, B., Acuna-Gonzales, J., 2003. N2 production by the anammox reaction in the anoxic water column of Golfo Dulce, Costa Rica. Nature 422, 606–608. Dalziel, I.W.D., 1997. OVERVIEW: Neoproterozoic-Paleozoic geography and tectonics: Review, hypothesis, environmental speculation. Geol. Soc. Am. Bull. 109, 16–42, doi:10.1130/0016-7606. Davelaar, D., 1993. Ecological significance of bacterial polyphosphate metabolism in sediments. Hydrobiologia 253, 179–192. Des Marais, D.J., Strauss, H., Summons, R.E., Hayes, J.M., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature 359, 605–609, doi:10.1038/359605a0. Ding, L., Zhang, L., Li, Y., Dong, J., 1992. The Study of the Late Sinian–Early Cambrian Biota from the Northern Margin of Yangtze Platform. Scientific and Technical Documents Publishing House, Beijing, pp. 1–156. Dornbos, S.Q., 2006. Evolutionary palaeoecology of early epifaunal echinoderms: Response to increasing bioturbation levels during the Cambrian radiation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 237, 225–239. Droser, M.L., 1987. Trends in extent and depth of bioturbation in Great Basin Precambrian–Ordovician strata, California, Nevada, and Utah. Ph.D. Thesis. University of Southern California, Los Angeles. Fairchild, I.J., Kennedy, J., 2007. Neoproterozoic glaciation in the Earth System. J. Geol. Soc. 164, 895–921. Falkowski, P.G., Fenchel, T., Delong, E.F., 2008. The microbial engines that drive earth’s biogeochemical cycles. Science 320, 1034–1038, doi:10.1126/science.1153213. Falkowski, P.G., 1997. Evolution of the nitrogen cycle and its influence on the biological sequestration of CO2 in the ocean. Nature 387, 272–275. Fanning, C.M., Link, P.K., 2004. U–Pb SHRIMP ages of Neoproterozoic (Sturtian) glaciogenic Pocatello Formation, southeastern Idaho. Geol. Boulder 32 (10), 881–884. Fike, D.A., Grotzinger, J.P., Pratt, L.M., Summons, R.E., 2006. Oxidation of the Ediacaran ocean. Nature 444, 744–747, doi:10.1038/nature05345. Fogel, M.L., Cifuentes, L.A., 1993. Isotope fractionationduring primary production. In: Engel, M.H., Macko, S.A. (Eds.), Organic Geochemistry. Plenum Press, New York, pp. 73–98. Föllmi, K.B., 1996. The phosphorus cycle, phosphogenesis and marine phosphaterich deposits. Earth-Sci. Rev. 40, 55–124. Francois, R., Altabet, M.A., Burckle, L.H., 1992. Glacial to interglacial changes in surface nitrate utilization in the Indian sector of the Southern Ocean as recorded by sediment N15 N. Paleoceanography 7, 589–606. Freudenthal, T., Wagner, T., Wenzhofer, F., Zabel, M., Wefer, G., 2001. Early diagenesis of organic matter from sediments of the eastern subtropical Atlantic: Evidence from stable nitrogen and carbon isotopes. Geochim. Cosmochim. Acta 65 (11), 1795–1808. Fuchsman, C.A., Murray, J.W., Konovalov, S.K., 2008. Concentration and natural stable isotope profiles of nitrogen species in the Black Sea. Marine Chem. 111, 90–105. Galbraith, E.D., Sigman, S.M., Robinson, R.S., Pedersen, F.T., 2008. Past changes in the marine nitrogen cycle. In: Capone, D., Bronk, D., Mulholland, M., Carpenter, E. (Eds.), Nitrogen in the Marine Environment. , second ed. Elsevier. Ganeshram, R.S., Pedersen, T.F., Calvert, S.E., McNeill, G.W., Fontugne, M.R., 2000. Glacial-interglacial variability in denitrification in the world’s oceans: Causes and consequences. Paleoceanography 15 (4), 361–376. Garvin, J., Buick, R., Anbar, A.D., Arnold, G.L., Kaufman, A.J., 2009. Isotopic evidence for an Aerobic Nitrogen Cycle in the Latest Archean. Science 323, 1045–1048. Gaye-Haake, B.T., Lahajnar, N., Emeis, K.Ch., Unger, D., Rixen, T., Suthhof, A., Ramaswamy, V., Schulz, H., Paropkari, A.L., Guptha, M.V.S., Ittekkot, V., 2005. Stable nitrogen isotopic ratios of sinking particles and sediments from the northern Indian Ocean. Marine Chem. 96, 243–255. Gilbert, F., Stora, G., Bonin, P., 1998. Influence of bioturbation on denitrification activity in Mediterranean coastal sediments: an in situ experimental approach. Mar. Ecol. Prog. Ser. 163, 99–107. Gilbert, F., Stora, G., Bonin, P., LeDréau, Y., Mille, G., Bertrand, J.C., 1997. Hydrocarbon influence on denitrification in bioturbated Mediterranean coastal sediments. Hydrobiologia 345, 67–77. Gilbert, F., Bonin, P., Stora, G., 1995. Effect of bioturbation on denitrification in a marine sediment from the West Mediterranean littoral. Hydrobiologia 304, 49–58. Godfrey, L.V., Glass, J.B., 2011. The geochemical record of the ancient nitrogen cycle, nitrogen isotopes and metalc co-factors. Methods Enzymol. 486, 483–506, doi:10.1016/S0076-6879(11)86022-2. Godfrey, L.V., Falkowski, P.G., 2009. The cycling and redox state of nitrogen in the Archean ocean. Nat. Geosci., doi:10.1038/NGE0633. Goldberg, T., Strauss, H., Guo, Q., Liu, C., 2007. Reconstructing marine redox conditions for the Early Cambrian Yangtze Platform: evidence from biogenic sulphur and organic carbon isotopes. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 172–190, doi:10.1016/j.palaeo.2007.03.015. Goldhammer, T., Brüchert, V., Ferdelman, T.G., Zabel, M., 2010. Microbial sequestration of phosphorus in anoxic upwelling sediments. Nat. Geosci. 3, 557–561, doi:10.1038/ngeo913. Grotzinger, J.P., Bowring, S.A., Beverly, B.Z., Kaufman, A.J., 1995. Biostratigraphic and geochronologic constraints on early animal evolution. Science 270, 598–604. Grundmanis, V., Murray, J.W., 1977. Nitrification and denitrification in marine sediments from Puget Sound. Limnol. Oceanogr. 22, 804–813.

163

Guo, Q., Strauss, H., Liu, C., Goldberg, T., Zhu, M., Pi, D., Heubeck, C., Vernhet, E., Yang, X., Fu, P., 2007. Carbon isotopic evolution of the terminal Neoproterozoic and early Cambrian: Evidence from the Yangtze Platform, South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 140–157. Hamersley, M.R., Lavik, G., Woebken, D., Rattray, J.E., Lam, P., Hopmans, E.C., Damste, J.S.S., Kruger, S., Graco, M., Gutierrez, D., Kuypers, M.M.M., 2007. Anaerobic ammonium oxidation in the Peruvian oxygen minimum zone. Limnol. Oceanogr. 52, 923–933. Hayes, J.M., Strauss, H., Kaufman, A.J., 1999. The abundance of C-13 in marine organic matter and isotopic fractionation in the global biogeochemical cycle of carbon during the past 800 Ma. Chem. Geol. 161 (1–3), 103–125. Hoering, T.C., Ford, H.T., 1960. The isotope effect in the fixation of nitrogen by Azotobacter. J. Am. Chem. Soc. 82, 376–378. Hoffmann, K.H., Condon, D.J., Bowring, S.A., Crowley, J.L., 2004. U–Pb zircon date from the Neoproterozoic Ghaub Formation, Namibia; constraints on Marinoan glaciation. Geol. Boulder 32 (9), 817–820. Hoffman, P.F., Kaufman, P.F., Halverson, A.J., Schrag, D.P., 1998. A neoproterozoic snowball earth. Science 281, doi:10.1126/science.281.5381.1342. Holmes, E., Lavik, G., Fischer, G., Segl, M., Ruhland, G.W., 2002. Seasonal variability of •15 N in sinking particles in the Benguela upwelling region. Deep-Sea Res. I 49, 377–394. Hough, M.L., Shields, G.A., Evins, L.Z., Strauss, H., Henderson, R.A., Mackenzie, S., 2006. A major sulphur isotope event at c. 510 Ma: a possibleanoxia–extinction–volcanism connection during the Early–Middle Cambrian transition? Terra Nova 18, 257–263. Hu, S., Steiner, M., Zhu, M., Erdtmann, B.-D., Luo, H., Chen, L., Weber, B., 2007. Diverse pelagic predators from the Chengjiang Lagerstätte and the establishment of modern-style pelagic ecosystems in the early Cambrian. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 307–316. Ingall, E., Kolowith, L., Lyons, T., Hurtgen, M., 2005. Sediment carbon, nitrogen and phosphorus cycling in an anoxic fjord, Effingham Inlet, British Columbia. Am. J. Sci. 305, 240–258. Ishikawa, T., Ueno, Y., Komiya, T., Sawaki, Y., Han, J., Shu, D., Li, Y., Maruyama, S., Yoshida, N., 2007. Carbon isotope chemostratigraphy of a Precambrian/Cambrian boundary section in the Three Gorge area, South China: Prominent global-scale isotope excursions just before the Cambrian Explosion. Gondwana Res. 14, 193–208. Jarvis, I., Burnett, W.C., et al., 1994. Phosphorite geochemistry: state-of-the-art and environmental concerns. Eclogae Geol. Helv. 87, 643–700. Johnston, D.T., Wolfe-Simon, F., Pearson, A., Knoll, A.H., 2009. Anoxygenic photosynthesis modulated Proterozoic oxygen and sustained Earth’s middle age. PNAS 106 (40), 16925–16929. Karl, D., Michaels, A., Bergman, B., Capone, D., Carpenter, E., Letelier, R., Lipschultz, F., Paerl, H., Sigman, D., Stal, L., 2002. Dinitrogen fixation in the world’s ocean. Biogeochemistry 57–58, 47–98. Kennedy, P., Kennedy, H., Papadimitriou, S., 2005. The effect of acidification on the determination of organic carbon, total nitrogen and their stable isotopic composition in algae and marine sediment. Rapid Commun. Mass Spectrom. 19, 1063–1068. Karl, D., Letelier, R., Tupas, L., Dore, J., Christian, J., Hebel, D., 1997. The role of nitrogen fixation in biogeochemical cycling in the subtropical North Pacific Ocean. Nature 388, 533–538. Kirschvink, J.L., Raub, T.D., 2003. A methane fuse for the Cambrian explosion: carbon cycles and true polar wander. C. R. Geosci. 335, 65–78. Krooss, B.M., Jurish, A., Plessen, B., 2006. Investigation of the fate of nitrogen in Palaeozoic shales of the Central European Basin. J. Geochem. Explor. 89, 191–194. Kuypers, M.M.M., Lavik, G., Woebken, D., Schmid, M., Fuchs, B.M., Amann, R., Jorgensen, B.B., Jetten, M.S.M., 2005. Proc. Natl. Acad. Sci. U.S.A. 102, 6478. Kuypers, M.M.M., Sliekers, A.O., Lavik, G., Schmid, M., Joergensen, B.B., Kuenen, J.G., Damste, J.S.S., Strous, M., Jetten, M.S.M., 2003. Anaerobic ammonium oxidation by anammox bacteria in the Black Sea. Nature 422 (6932), 608–611. Lam, P., Lavik, G., Jensena, M., de Vossenberg, V., Schmid, M., Woebken, D., Gutierrez, D., Amann, R., Jetten, M.S.M., Kuypers, M.M., 2009. Revising the nitrogen cycle in the Peruvian oxygenminimum zone. PNAS 106 (12), 4752–4757. Lam, P., Jensen, M.M., Lavik, G., McGinnis, D.F., Müller, B., Schubert, C.J., Amann, R., Thamdrup, B., Kuypers, M.M., 2007. Linking crenarchaeal and bacterial nitrification to anammox in the Black Sea. Proc. Natl. Acad. Sci. U.S.A. 104, 7104–7109. LaPorte, D.F., Holmden, C., Patterson, W.P., Loxton, J.D., Melchin, M.J., Mitchell, C.E., Finney, S.C., Sheets, H.D., 2009. Local and global perpectives on carbon and nitrogen cycling during the Hirnantian glaciation. Palaeogeogr. Palaeoclimatol. Palaeoecol. 276, 182–195. Lehmann, R.F., Bernasconi, S.M., Barbieri, A., Mckenzie, J.A., 2002. Preservation of organic matter and alteration of its carbon and nitrogen isotope composition during simulated and in situ early sedimentary diagenesis. Geochim. Cosmochim. Acta 66, 3573–3584. Li, C., Love, G.D., Timothy, T.W., Fike, D.A., Sessions, A.L., Chu, X., 2010. A Stratified Redox Model for the Ediacaran Ocean. Science 328, 80–83, doi:10.1126/science.1182369. Li, G., Steiner, M., Zhu, X., Yang, A., Wang, H., Erdtmann, B.D., 2007. Early Cambrian metazoan fossil record of South China: Generic diversity and radiation patterns. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 229–249. Li, G., Xiao, S., 2004. Tannuolina and Micrina (Tannuolinidae) from the Lower Cambrian of Eastern Yunnan, South China, and their scleritome reconstruction. J. Paleontol. 78 (5), 900–913. Li, W.X, Li, X.H., Li, Z.X., 2005. Neoproterozoic bimodal magmatism in the Cathaysia Block of South China and its tectonic significance. Precambrian Res. 136, 51–66.

164

L. Cremonese et al. / Precambrian Research 225 (2013) 148–165

Li, X.H., Li, Z.X., Ge, W., Zhou, H., Li, W., Liu, Y., Wingate, M.T.D., 2003a. Neoproterozoic granitoids in South China: crustal melting above a mantle plume at ca. 825 Ma? Precambrian Res. 45–83, 122, House of Yunnan, Kunming, Yunnan, pp. 154. Li, X.H., 1999. U–Pb zircon ages of granites from the southern margin of the Yangtze block: timing of the Neoproterozoic Jinning Orogeny in SE China and implications for Rodinia assembly. Precambrian Res. 97, 43–57. Li, Z.X., Li, X.H., Kinny, P.D., Wang, J., Zhang, S., Zhou, H., 2003b. Geochronology of Neoproterozoic syn-rift magmatism in the Yangtze Craton, South China and correlations with other continents: evidence for a mantle superplume that broke up Rodinia. Precambrian Res. 122, 85–109. Li, Z.X., Li, X.H., Zhou, H., Kinny, P.D., 2002. Grenvillian continental collision in South China: new SHRIMP U–Pb zircon results and implications for the configuration of Rodinia. Geology 30, 163–166. Li, Z.X., Li, X.-H., Kinny, P.D., Wang, J., 1999. The breakup of Rodinia: did it start with a mantle plume beneath South China? Earth Planet. Sci. Lett. 173 (3), 171–181. Li, Z.X, Zhang, L., Powell, C.W., 1995. South China in Rodinia: part of the missing link between Australia–East Antarctica and Laurentia? Geology 23, 407–410. Lin, J.P., Zhao, Y.L., Rahman, L.I., Xiao, S., Wang, Y., 2010. Bioturbation in Burgess Shale-type Lagerstätten — Case study of trace fossil–body fossil association from the Kaili Biota (Cambrian Series 3), Guizhou, China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 292, 245–256. Ling, W., Gao, S., Zhang, B., Li, H., Liu, Y., Cheng, J., 2003. Neoproterozoic tectonic evolution of the northwestern Yangtze craton, South China: implications for amalgamation and break-up of the Rodinia Supercontinent. Precambrian Res. 122, 111–140. Liu, K.K., Kaplan, I.R., 1988. Variation of nitrogen isotope fractionation during denitrification and nitrogen isotope balance in the ocean. EOS 69, 1098. Logan, G.A., Hayes, J.M., Hieshima, G.B., Summons, R.E., 1995. Terminal Proterozoic reorganization of biogeochemical cycles. Nature 736, 53–56. Lourey, M.J., Thomas, W., Trull Daniel, M., Sigman, 2003. Sensitivity of D15N of nitrate, surface suspended and deep sinkingparticulate nitrogen to seasonal nitrate depletion in the Southern. Global Biogeochem. Cylces 17 (3), 1081. Luo, H., Jiang, Z., Wu, X., Song, X., Ouyang, L., Xing, Y., Liu, G., Zhang, S., Tao, Y., 1984. Sinian–Cambrian Boundary Stratotype Section at Meishucun, Jinning, Yunnan, China. People’s Publishing. Macko, S.A., Marilyn, L.F., Hare, P.E., Hoering, T.C., 1987. Isotopic fractionation of nitrogen and carbon in the synthesis of amino acids by mocroorganisms. Chem. Geol. 65, 79–82. Macouin, M., Besse, J., Ader, M., Gilder, S., Yang, Z., Sun, Z., Agrinier, P., 2004. Combined paleomagnetic and isotopic data from the Doushantuo carbonates, South China: implications for the snowball Earth hypothesis. Earth Planet. Sci. Lett. 224, 387–398. Maloof, A.C., Porter, S.M., Moore, J.L., Dudás, F.Ö., Bowring, S.A., Higgins, J.A., Fike, D.A., Eddy, M.P., 2010. The earliest Cambrian record of animals and ocean geochemical change. Geolog. Soc. Am. Bull. 122, 1731–1774, doi:10.1130/B30346.1. Mariotti, A., Mariotti, F., Amarger, N., Pizelle, G., Ngambi, J.M., Champigny, M.L., Moyse, A., 1980. Fractionnements isotopiques de l’azote lors des processus d’absorption des nitrates et de fixation de l’azote atmospherique par les plantes. Physiol. Ve’g. 18, 163–181. McFadden, K.A., Huang, J., Chu, X., Jiang, G., Kaufman, A.J.J., Zhou, C., Yuan, X., Xiao, S., 2008. Pulsed oxidation and biological evolution in the Ediacaran Doushantuo Formation. PNAS 105 (9), 3197–3202, doi:10.1073/pnas.0708336105. McIlroy, D., Logan, G.A., 1999. The impact of bioturbation on infaunal ecology and evolution during the proterozoic-cambrian transition. Palaios 14 (1), 58–72. Meisel, S., Struck, U., 2011. The potential distortion of sedimentary d15N and Corg/N ratios by NH4+ and the effects of pre-analyses sample treatment. Fossil Record. 14, 141–152. Meyers, P.A., Bernasconi, S.M., Yum, J.-G., 2009. 20My of nitrogen fixation during deposition of mid-Cretaceous black shales on the Demerara rise, equatorial Atlantic Ocean. Org. Geochem. 40, 158–166. Minagawa, M., Wada, E., 1986. Nitrogen isotopes ratios of red tide organisms in the east China sea: a characterisation of biological nitrogen fixation. Marine Chem. 19, 245–259. Missarzhevsky, V.V., 1969. Description of hyolithids, gastropods, hyolithelminths, camenides and forms of an obscure taxonomic position, pp. 103–175 (pp. 127–205 in English translation by Amerind Publishing, New Delhi, 1981). In: M. E. Raaben (Ed.), The Tommotian Stage and the Cambrian Lower Boundary Problem. Akademiya Nauka SSSR, Moscow. Missarzhevsky, V.V., 1973. Konodontoobraznye organizmy iz pogranichnykh sloev kembriya i dokembriya Sibirskoj platformy i Kazakhstana. [Conodont-shaped organisms from Precambrian–Cambrian boundary strata of the Siberian Platform and Kazakhstan]. Trudy Instituta. Geologii i Geofiziki SO AN SSSR, vol. 49; pp. 53–57. Montoya, J.P., Carpenter, E.J., Capone, D.G., 2002. Limnol. Oceanogr. 47 (6), 1617–1628. Montoya, J.P., Wiebe, P.H., McCarthy, J.J., 1992. Natural abundance of I5N in particulate nitrogen and zooplankton in the Gulf Stream region and Warm-Core Ring 86A. Deep-Sea Res. 3 (Suppl. 1), S363–S392. Montoya, J.P., 1994. Nitrogen isotope fractionation in the modern ocean: implications for the sedimentary record. In: Zahn, R., Kaminski, M.A., Labeyrie, L., Pederson, T.F. (Eds.), Carbon Cycling in the Glacial Ocean: Constraints on the Ocean’s Role in Global Change. Springer-Verlag, Berlin. Ng, J.S.S., Wai, T.-C., Williams, G.A., 2007. The effects of acidification on the stable isotope signatures of marine algae and molluscs. Marine Chem. 103, 97–102.

Ohkouchi, N., Nakajima, Y., Okada, H., Ogawa, H., Suga, H., Oguri, K., Kitazato, H., 2005. Biogeochemical processes in the saline meromictic Lake Kaiike, Japan: implications from molecularisotopic evidences of photosynthetic pigments. Environ. Microbiol. 7 (7), 1009–1116. Owens, N.J. P., 1987. Natural variations in nitrogen-15 in the marine environment. In: Blaxter, J.H.S., Southward, A.J. (Eds.). Advance in marine biology, vol.24 XII+473 P. Academic Press, INC.: San Diego, California, USA; London, England, UK. Illus. Maps, 389-451. Papineau, D., Puroith, R., Goldberg, T., Pi, D.-H., Shields, G.A., Bhuf, H., Steele, A., Fogel, L., 2009. High primary productivity and nitrogen cycling after the Paleoproterozoic phosphogenic event in the Aravalli Supergroup, India. Precambrian Res. 171, 37–56. Passier, H.F., Bosch, H., Lourens, L.J., Böttcher, M.E., Leenders, A., Damste, J.S.S., de Lange, G.J., de Leeuw, J.W., 1999. Sulfidic Mediterranean surface waters during Pliocene sapropel formation. Nature 397, 146–149. Payne, J.L., Boyers, A.B., Brown, J.H., Finnegan, S., Kowalewski, M., Kraus, R.A., Kathleen Lyonse Jr., D.S., McClain, C.R., McShea, D.W., Novack-Gottshall, P.M., Smith, F.A., Stempien, J.A., Wang, S.C., 2009. Two-phase increase in the maximum size of life over3.5 billion years reflects biological innovation and environmental opportunity. PNAS 106 (1), 24–27, doi:10.1073/pnas.0806314106. Pennock, J.R., Velinsky, D.V., Ludlam, J.M., Sharp, J.H., Fogel, M.L., 1996. Isotopic fractionation of ammonium and nitrate during uptake by Skeletonema costatum: Implications for •15 N dynamics under bloom conditions. Limn. Oceanogr. 41, 451–459. Peterson, B.J., Fry, B., 1987. Stable isotopes in ecosystem studies. Ann. Rev. Ecol. System. 18, 293–320. Qian, Y., 1977. Hyolitha and some problematica from the Lower Cambrian Meishucun Stage in central and southwestern China. Acta Palaeontologica Sinica 16, 255–275. Quan, T.M., van de Schootbrugge, B., Field, P.M., Rosenthal, Y., Falkowsky, P.G., 2008. Nitrogen isotope and trace metal alalyses from the Mingolsheim core (Germany): Evidence for redox variations across the Triassic-Jurassic boundary. Global Biogeochemical cycles 22, GB2014, doi:10.1029/2007GB002981. Rothman, D.H., Hayes, J.M., Summons, R., 2003. Dynamics of the Neoproterozoic carbon cycle. Proc. Natl. Acad. Sci. U.S.A. 100, 8124–8129. Rui, Z.Q., Piper, J.D.A., 1997. Palaeomagnetic study of Neoproterozoic glacial rocks of the Yangzi Block: palaeolatitude andconfiguration of South China in the late Proterozoic supercontinent. Precambrian Res. 85, 173–199. Saltzman, M.R., 2005. Phosphorus, nitrogen, and the redox evolution of the Paleozoic oceans. Geology 33, 573–576, doi:10.1130/G21535.1. Scholten, S.O., 1991. The distribution of nitrogen isotopes in sediments. Geologica Ultraiectina 81, 101. Sannigrahi, P., Ingall, E., 2005. Polyphosphates as a source of enhanced P fluxes in marine sediments overlain by anoxic waters: evidence from 31P NMR. Geochem. Trans. 6 (3), 52–59. Shen, Y., Zhang, T., Hoffman, P.F., 2008. On the coevolution of Ediacaran oceans and animals. PNAS 105 (21), 7376–7738, doi:10.1073/pnas.0802168105. Sigman, D.M., Altabet, M.A., McCorkle, D.C., Francois, R., Fischer, G., 2000. The d15N of nitrate in the Southern Ocean: nitrogen cycling and circulation in the ocean interior. Journal of Geophysical Research 105 (C8), 19599–19614. Sigman, D.M., Karsh, K.L., Casciotti, K.L., 2009. Nitrogen isotopes in the ocean. In: Steele, J.H., Thorpe, S.A., Turekian, K.K. (Eds.). Encyclopedia of Ocean Sciences. Sigman, D.M., Altabet, D.C., McCorkle, D.C., Francois, R., Fisher, G., 1999. The •15N of nitrate in the Southern Ocean: Consumption of nitrate in surface waters. Global Biogeochem. Cycles 13 (4), 1149–1166. Sinninghe Damstè, J.S., Koster, J., Buss, M., 1998. A sedimentary tatrahydrophenanthrene derivative of tetrahymanol. Tatrahedron Lett. 40, 3949–3952. Slomp, C.P., Van der Gaast, S.J., Van Raaphorst, W., 1996. Phosphorus binding by poorly crystalline iron oxides in North sea sediments. Mar. Chem. 52, 55–73. Smith, S.L., Henrischs, S.M., Tho, T., 2002. Stable C and N isotopic composition of sinking particles and zooplankton over the southeastern Bering Sea shelf. DeepSea Res. 49, 6031–6050. Steiner, M., Li, G., Qian, Y., Zhu, M., Erdtmann, B.-D., 2007. Neoproterozoic to early Cambrian small shelly fossil assemblages and a revised biostratigraphic correlation of the Yangtze Platform (China). Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 67–99. Steiner, M., Zhu, M., Zhao, Y., Erdtmann, B.-D., 2005. Lower Cambrian Burgess Shaletype fossil associations of South China. Palaeogeogr. Palaeoclimatol. Palaeoecol. 220, 129–152. Steiner, M.R.J., 2001. Evidence of organic structures in Ediacara-type fossils and associated microbial mats. Geology 29 (12), 1119–1122. Struck, U., Falk, M., Altenbach, A.V., Pollehne, F., Schneider, M., 2009. Nitrogen and carbon isotope ratios in suspended matter and dissolved inorganic carbon in a meromictic lake of the northern Alps (Bavaria, Germany). In: ASLO Aquatic Sciences Meeting 2009, vol. 25-30 January 2009 in Nice, France. Struck, U., Emeis, K.-C., Voss, M., Krom, M.D., Rau, G.H., 2001. Biological productivity during sapropel S5 formation in the Eastern Mediterranean Sea: Evidence from stable isotopes of nitrogen and carbon. Geochim. Cosmochim. Acta 65 (19), 3249–3266. Swanson-Hysell, N.L., Rose, C.V., Calmet, C.C., Halverson, G.P., Hurtgen, M.T., Maloof, A.C., 2010. Cryogenian glaciation and the onset of carbon-isotope decoupling. Science 328, 608–611, doi:10.1126/science.1184508. Thamdrup, B., Dalsfaard, T., Jensen, M., Ulloa, O., Farias, L., Escribano, R., 2006. Anaerobic ammonium oxidation in the oxygen-deficient waters off northern Chile. Limnol. Oceanogr. 51, 2145–2156.

L. Cremonese et al. / Precambrian Research 225 (2013) 148–165 Thamdrup, B., Dalsgaard, T., 2002. Production of N2 through anaerobic ammonium oxidation coupled to nitrate reduction in marine sediments. Appl. Environ. Microbiol. 68, 1312–1318. Thomazo, C., Pinti, D.L., Busigny, B., Ader, M., Hashizume, K., Philippot, P., 2009. General palaeontology (Palaeobiochemistry). Biological activity and the Earth’s surface evolution: Insights from carbon, sulfur, nitrogen and iron stable isotopes in the rock record. C. R. Palevol 8, 665–678. Thomazo, C., Ader, A., Philippot, P., 2011. Extreme 15 N-enrichments in 2.72 Gyr old sediments. Evidemtns for a turning point in the nitrogen cycle. Geociology 9, 107–120. Trappe, J., 1998. Phanerozoic phosphorite depositional sistems. A dynamic model for a sedimentary resource system. Lecture Notes Earth Sci. 76, 316. Tyrrell, T., 1999. The relative influences of nitrogen and phosphorus on oceanic primary production. Nature 400, 525–531. Vannier, J., Steiner, M., Renvoise, E., Hu, S.-X., Casanova, J.-P., 2007. Early Cambrian origin of modern food webs: evidence from predator arrow worms. Proc. R. Soc. B 274, 627–633. Voss, M., Nausch, G., Montoya, J.P., 1997. Nitrogen stable isotope dynamics in the central Baltic Sea: influence of deep-water renewal on the N-cycle changes. Marine Ecol. Prog. Ser. 158, 11–21. Wada, E., Hattori, A., 1991. Nitrogen in the Sea: Forms, Abundances & Rate Processes. CRC Press, Florida, USA. Wada, E., Hattori, A., 1976. Natural abundance of 15 N in particulate organic matter in the North Pacific Ocean. Geochimica and Cosmochimica Acta 40, 249–251. Wada, E., Kadonaga, T., Matsuo, S., 1975. 15 N abundance in nitrogen of naturally occurring substances and global assessment of denitrification from isotopic viewpoint. Geochem. J. 9, 139–148. Wang, J., Li, Z.X., 2003. History of the Neoproterozoic rift basins in South China: implications for Rodinia break-up. Precambrian Res. 122, 141–158. Wang, J., 2000. Neoproterozoic Rifting History of South China: Significance to Rodinia Breakup. Geological Publishing House, Beijing, pp. 1–145. Wang, H., Mo, X., 1985. An outline of the tectonic evolution of China. Episodes 18 (1,2), 6–16.

165

Ward, B.B., Devol, A.H., Rich, J.J., Chang, B.X., Bulow, S.E., Naik, H., Pratihary, A., Jayakumar, A., 2009. Denitrification as the dominant nitrogen loss process in the Arabian Sea. Nature 461, 78–82. Weber, B., Steiner, M., Zhu, M.-Y., 2007. 349 Precambrian–Cambrian trace fossils from the Yangtze Platform (South China) and the early evolution of bilaterian lifestyles. Palaeogeogr. Palaeoclimatol. Palaeoecol. 254, 328–349. Wille, M., Nägler, T.F., Lehmann, B., Schröder, S., Kramers, J.D., 5 June 2008. Hydrogen sulphide release to surface waters at the Precambrian/Cambrian boundary. Nature 453, doi:10.1038/nature07072. Zerkle, A.L., Junium, C.K., Canfield, D.E., House, C.H., 2008. Production of 15 N-depleted biomass during cyanobacterial N2 -fixation at high Fe concentrations. J. Geophys. Res. 113, G03014. Zhao, J.-H., Zhou, M.-F., Yan, D.-P., Zheng, J.-P., Li, J.-W., 2011. Reappraisal of the ages of Neoproterozoic strata in South China: no connection with the Grenvillian orogeny. Geology 39, 299–302. Zhou, C., Zhang, J., Li, G., YU, Z., 1997. Carbon and oxygen isotopic record of the Early Cambrian from the Xiaotan Section, Yunnan, South China. Scientia Geologica Sinica 32, 201–211. Zhou, J.C., Wang, X.L., Qiu, J.S., Gao, J.F., 2004. Geochemistry of Mesoand Neoproterozoic mafic–ultramafic rocks from northern Guangxi, China: arc or plume magmatism? Geochem. J. 38, 139–152. Zhou, J.C., Wang, X.L., Qiu, J.S., Gao, J.F., 2003. Lithogeochemistry of Meso- and Neoproterozoic mafic–ultramafic rocks from northern Guangxi. Acta Petrolei Sinica 19, 9–18. Zhou, M.F., Kennedy, A.K., Sun, M., Malpas, J., Lesher, C.M., 2002a. Neoproterozoic arcrelated mafic intrusions along the Northern Margin of South China: implications for the accretion of Rodinia. J. Geol. 110, 611–618. Zhou, M.F., Yan, D., Kennedy, A.K., Li, Y., Ding, J., 2002b. SHRIMP U–Pb zircon geochronological and geochemical evidence for Neoproterozoic arc-magmatism along the western margin of the Yangtze Block, South China. Earth Planet. Sci. Lett. 196, 51–67. Zhu, M., Zhang, J., Steiner, M., Yang, A., Li, G., Erdtmann, B.-D., 2003. Sinian and early Cambrian stratigraphic frameworks from shallow to deep-water facies of the Yangtze Platform: an integrated approach. Prog. Natural Sci. 13 (12), 951–960.