Organic matter–apatite–pyrite relationships in the Botneheia Formation (Middle Triassic) of eastern Svalbard: Relevance to the formation of petroleum source rocks in the NW Barents Sea shelf

Marine and Petroleum Geology 45 (2013) 69e105

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Organic mattereapatiteepyrite relationships in the Botneheia Formation (Middle Triassic) of eastern Svalbard: Relevance to the formation of petroleum source rocks in the NW Barents Sea shelf Krzysztof P. Krajewski* Institute of Geological Sciences, Polish Academy of Sciences, Research Centre in Warsaw, Twarda 51/55, 00-818 Warszawa, Poland

a r t i c l e i n f o

a b s t r a c t

Article history: Received 2 May 2012 Received in revised form 2 April 2013 Accepted 18 April 2013 Available online 3 May 2013

The Middle Triassic Botneheia Formation of eastern Svalbard (Edgeøya and Barentsøya) comprises an organic carbon-rich, fine-grained clastic succession (w100 m thick) that makes the best petroleum source rock horizon in the NW Barents Sea shelf. The succession records a transgressiveeregressive interplay between the prodelta depositional system sourced in the southern Barents Sea shelf (black shale facies of the lower and middle parts of the Muen Member) and the open shelf phosphogenic system related to upwelling and nutrient supply from the Panthalassic Ocean (phosphogenic black shale facies of the upper part of the Muen Member and the Blanknuten Member). The relationships between organic matter, authigenic apatite, and pyrite in these facies allow to characterize the relative roles of redox conditions and oceanic productivity in the organic carbon preservation. The accumulation of terrestrial and autochthonous marine organic matter in the black shale facies occurred under dominating oxic conditions and increasing-upward productivity related to early transgressive phase and retrogradation of the prodelta system. The phosphogenic black shale facies deposited in an oxygen-minimum zone (OMZ) of the open shelf environment during the late transgressive to regressive phases under conditions of high biological productivity, suppressed sedimentation rates, and changing bottom redox. The phosphatic black shales occurring in the lower and upper parts of the phosphogenic succession reveal depositional conditions indicative of the shallower part of OMZ, including high input of autochthonous organic matter into sediment, oxic-to-dysoxic (episodically suboxic and/or anoxic) conditions, intense phosphogenesis, and recurrent reworking of the seabed. The massive phosphatic mudstone occurring in the middle of the phosphogenic succession reflects the development of euxinia in the deeper part of OMZ during high-stand of the sea. High input of autochthonous organic matter in this environment was coupled with mineral starvation and intermittent phosphogenesis. In mature sections in eastern Svalbard, the petroleum potential of the Botneheia Formation rises from moderate to good in the black shale facies, and from good to very good in the phosphogenic black shale facies, attaining maximum in the massive phosphatic mudstone. Ó 2013 Elsevier Ltd. All rights reserved.

Keywords: Svalbard Triassic Petroleum source rocks Organic matter Apatite Pyrite Redox Productivity

1. Introduction The Mesozoic basins of the circum-Arctic margins contain Triassic successions of fine-grained, organic carbon (OC)-rich deposits that are considered to be important petroleum source rocks in the region (Bird and Houseknecht, 2011; Bjorøy et al., 2010; Brekhuntsov et al., 2011; Dewing and Obermajer, 2011; Embry, 2011). These deposits are widespread in the Alaska North Slope (Magoon and Bird, 1987; Peters et al., 2006; Robinson et al., 1996),

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the Sverdrup Basin of Arctic Canada (Chen et al., 2000; Embry, 1988, 1991), the Barents Sea shelf (Mørk and Elvebakk, 1999; P celina, 1996; Worsley, 2008), and the basin of New Siberian Islands (Egorov et al., 1987; Korcinskaja, 1977; Preobra zenskaja et al., 1975). Paleogeographic reconstructions show their repositories forming a patchy belt of embayments and shallow basins along the northwestern margin of Pangea (Golonka, 2011; Golonka et al., 2003). The search for a major cause of the development of these deposits involves interpretations that highlight the role of stagnation, water stratification, and shelf anoxia in enhanced preservation of organic matter (Gentzis et al., 1996; Høy and Lundschien, 2011; Leith et al., 1992; Steel and Worsley, 1984). However, at many locations these deposits are remarkably enriched in sedimentary phosphate,

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showing facies associations with phosphorite and/or glauconiterich strata (Detterman, 1989; Egorov and Baturin, 1987; Embry et al., 1988; Krajewski, 2000a). This suggests the existence of zones of high biological productivity at least at parts of the Pangean margin, which were probably reinforced by upwelling from the Panthalassic Ocean (Krajewski, 2011; Parrish et al., 2001; Pcelina and Kor cinskaja, 2008). More data are necessary to reconcile these views and improve the understanding of the OC-rich depositional systems of the Arctic Triassic. This paper presents the Middle Triassic Botneheia Formation of Svalbard, NW Barents Sea shelf, which is one of key examples of the OC-rich facies of the Arctic Triassic. The relationships between organic matter and two major authigenic mineral phases, apatite and pyrite, have been investigated along the eastern Svalbard outcrop belt (Edgeøya and Barentsøya) in an attempt to characterize the relative roles of redox conditions and oceanic productivity in the deposition of the OC-rich facies. The results of an integrated study using petrographic methods, Rock-Eval pyrolysis, and PeFeeS geochemistry allow to propose a model that links the migration and expansion of oxygen-minimum zone related to upwelling and high biological productivity with phases of the Middle Triassic transgressiveeregressive cycle. Major depositional regimes in this cycle are defined in terms of processes leading to preservation of organic matter and petroleum potential of the OC-rich sedimentary facies. 2. Geological setting Svalbard (Fig. 1) is the emergent northwestern corner of the Barents Sea shelf (Dallmann et al., 2002; Faleide et al., 2008;

Harland, 1997), which, during the Triassic times, formed a broad clastic shelf basin facing Panthalassic Ocean to the north (Cocks and Torsvik, 2007). The depositional area was likely a big embayment narrowing southward (Worsley, 2008). It was bordered on the west by northern Greenland, and on the east and southeast by extended deltaic systems of the southern Barents Sea shelf (Høy and Lundschien, 2011; Riis et al., 2008). This basin was subject to alternating sea-level rises and progradations of the deltaic systems that resulted in its stepwise infill with clastic sediments (Gløstard-Clark et al., 2010, 2011). At least four transgressiveeregressive cycles can be discerned in the Triassic succession of Svalbard (Egorov and Mørk, 2000; Mørk and Smelror, 2001; Mørk and Worsley, 2006; Mørk et al., 1989). In the local lithostratigraphic scheme, they are classified in the Sassendalen (three cycles during the EarlyeMiddle Triassic) and Kapp Toscana (at least one cycle during the Late Triassic) groups (Mørk et al., 1999). The Middle Triassic cycle was unique in that it resulted in the deposition of black shale facies showing prominent concentration of both oil-prone organic carbon and authigenic mineral phosphorus (Abdullah, 1999; Krajewski, 2000a; Mørk and Bjorøy, 1984; Pcelina, 1985; Schou et al., 1984). It is recognized as the Bravaisberget and Botneheia formations in the upper part of the Sassendalen Group (Krajewski, 2008; Krajewski et al., 2007; Mørk et al., 1999). The formations provide a section of the NW part of the Barents Sea shelf, from deltaic to shallow marine settings in southern Spitsbergen through inner shelf setting in western Spitsbergen (Bravaisberget Formation) to outer shelf setting in central and eastern Spitsbergen and eastern Svalbard (Botneheia Formation). The depositional area in southern and western

Figure 1. (a) Map of Svalbard showing the Triassic outcrop belt and crucial locations of the Middle Triassic Botneheia and Bravaisberget formations. Black circles indicate sections of the Botneheia Formation in eastern Svalbard analyzed in this paper. The boundary between the areas of occurrence of the Botneheia and Bravaisberget formations is marked by a dashed line. The Triassic outcrop belt after Dallmann et al. (2002). (b) Reconstruction of the regional setting of the Barents Sea shelf and major depositional areas during maximum flooding of the Middle Triassic transgressiveeregressive cycle (early Ladinian). Based on Cocks and Torsvik (2007), Gløstard-Clark et al. (2010), Krajewski (2000a), Mørk et al. (1982), Riis et al. (2008), Steel and Worsley (1984), and Worsley (2008). The part of the NW Barents Sea shelf south and east of Svalbard is named the Svalbard Platform after Faleide et al. (1984). Indicated are structural elements of the shelf that are referred to in the text. For detailed information on the tectonic structure of the shelf see Gabrielsen et al. (1990). Br. and Bt. are depositional areas of the Bravaisberget and Botneheia formations, respectively.

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Spitsbergen was sourced from the southwest, most probably from a land area located in northern Greenland (Mørk et al., 1982; Steel and Worsley, 1984). The depositional area in eastern Svalbard was supplied by sediments from the southeast (Gløstard-Clark et al., 2010; Høy and Lundschien, 2011). It contains the outermost facies setting recognized in the shelf basin (Fig. 1). This setting comprises OC-rich, fine-grained clastic facies arranged in a rather simple transgressiveeregressive succession (Fig. 2). The succession is best seen along western Edgeøya and on both sides of Freemansundet in northern Edgeøya and southern Barentsøya (Fig. 3). It contains the non-phosphogenic black shale facies in its lower part that represents an early transgressive phase. This facies is overlain by a phosphogenic black shale facies deposited during a late transgressive phase through a regressive phase, with a well-defined sediment-starved period during high-stand of the sea (Krajewski, 2008). These phases become less pronounced northeastward in Barentsøya, where a submarine swell provided disturbance to the transgressiveeregressive facies arrangement. It was associated with thickness reduction of the Botneheia Formation, and rapid lithological changes and common traces of intraformational reworking in the phosphogenic succession. The Middle Triassic Botneheia Formation in eastern Svalbard (up to 105 m thick) rests on the Vikinghøgda Formation (?InduaneOlenekian) and is overlain by the Tschermakfjellet Formation (Carnian)

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(Mørk et al., 1999). It is subdivided into the lower Muen Member and the upper Blanknuten Member, which are further divided into nine informal lithostratigraphic units (Krajewski, 2008). The boundary between the two members is defined at base of the upper, cliff-forming part of the succession (Mørk et al., 1982). It is not coincident with the boundary between the non-phosphogenic and phosphogenic shale facies (Fig. 3). The latter boundary is discrete in most sections, marked by an appearance of small phosphate nodules in shaly sediment of the upper part of the Muen Member. The Botneheia Formation is of the AnisianeLadinian age, though the biostratigraphic relations in eastern Svalbard are poorly constrained (Buchan et al., 1965; Flood et al., 1971; Klubov, 1965a,b; Kor cinskaja, 1972; Lock et al., 1978; Tozer, 1973; Tozer and Parker, 1968). Detailed magneto-biostratigraphic scheme of the formation is only known for central Spitsbergen (Dagys and Weitschat, 1993; Hounslow et al., 2007, 2008; Weitschat and Dagys, 1989; Weitschat and Lehmann, 1983; Xu et al., 2009). On its basis, it may be suggested that the non-phosphogenic and the phosphogenic facies in eastern Svalbard are of the early to middle Anisian age and the middleelate Anisian through Ladinian age, respectively. However, the boundary between the two facies seems to be diachronous over Svalbard and more complex in Spitsbergen than in Edgeøya and Barentsøya. The Botneheia Formation in central Spitsbergen contains a basal phosphate-bearing unit (Hounslow et al., 2008),

Figure 2. The Botneheia Formation in eastern Svalbard showing its lithostratigraphic subdivision, major facies, and stages of the Middle Triassic transgressiveeregressive cycle. The studied sections and location of samples analyzed in this paper are indicated.

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Figure 3. Facies of the Botneheia Formation in eastern Svalbard. (a) Position of the Botneheia Formation in the Triassic succession; here at Muen in western Edgeøya (stratotype for the Muen Member). The formation records the Middle Triassic transgressiveeregressive cycle in open shelf setting of the NW Barents Sea. Thickness of the formation in the Muen outcrop belt is approximately 100 m. (b) Section of the upper part of the Botneheia Formation at Blanknuten, western Edgeøya (stratotype for the Blanknuten Member) showing succession of the phosphogenic black shale facies over non-phosphogenic black shale facies. Numbers indicate informal lithostratigraphic units of the formation. Note the tripartite nature of the cliff-forming succession of the Blanknuten Member, with the middle part made by massive phosphatic mudstone (Unit 7). This mudstone is the best petroleum source rock in Svalbard, reflecting an euxinic event under high biological productivity conditions during high-stand phase of the Middle Triassic cycle. The maximum flooding surface is suggested to be located in the lower part of Unit 7. Paler-weathering beds in Units 5 and 6 are dolomitic cementstones. (c) Uppermost part of the Botneheia Formation at Palibinranten, northwestern Edgeøya shows the phosphogenic facies succession reflecting regressive phase of the Middle Triassic cycle. Black phosphatic shales and siltstones of Unit 8 are overlain with erosional boundary by a phosphorite conglomerate horizon (Unit 9A) that marks reworking and condensation episode during termination of the regressive phase. This horizon is at places overlain by a thin phosphatic black shale (9B) with accumulations of Tasmanites. Unit 9 is disconformably overlain by sideritic shales of the Tschermakfjellet Formation. (d). Section of the lower part of the Botneheia Formation at Blanknuten showing black shale succession of the Muen Member (Units 2e4). The succession comprises coarsening-upward shale/mudstone packages that terminate in silty horizons variably cemented by microcrystalline dolomite (arrows). They reflect pulses of fine-grained sediment supply to a distal prodelta environment during early transgressive phase of the Middle Triassic cycle. Units 1 (not visible on the photo) and 3 contain more homogeneous shale intervals reflecting periods of increased supply of fine-grained clastics. The appearance of common phosphate nodules at base of Unit 5 marks the onset of high biological productivity conditions during the late transgressive phase.

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which is missing in eastern Svalbard. This unit is most probably of the earliest Anisian age, and seems to represent remnant of an older depositional cycle preserved along the hiatal surface at top of the Vikinghøgda Formation. The high-stand phase in the middle of the Blanknuten Member is dated to be early Ladinian in Spitsbergen (Weitschat and Dagys, 1989; Weitschat and Lehmann, 1983) and in eastern Svalbard (W. Weitschat, personal communication). This is consistent with the suggested age of the Middle Triassic maximum flooding surface in deltaic systems of the southern Barents Sea shelf (Gløstard-Clark et al., 2010, 2011; Riis et al., 2008). It is likely that parts of late Ladinian sediments at the top of the Blanknuten Member are missing due to reworking and condensation during the regressive phase and erosional truncation before the Late Triassic sedimentary cycle (Dagys et al., 1993; Hounslow et al., 2007, 2008). 3. Materials and methods The Botneheia Formation was studied at eight locations in eastern Svalbard (Fig. 1), six in Edgeøya (Vogelberget, Reddikeidet, Muen, Blanknuten, Palibinranten, and Skrukkefjellet), and two in Barentsøya (Skarpryttaren and Isormen). In an attempt to reveal lateral facies changes, supplementary observations were collected between the studied sections. More than one hundred samples of all the discerned facies and lithological units of the Botneheia Formation were selected for this study (Fig. 2). Nine samples from the topmost part of the Vikinghøgda Formation at Blanknuten and seven samples from the bottom part of the Tschermakfjellet Formation at Blanknuten and Skarpryttaren were also analyzed. Samples for Rock-Eval and geochemical analyses were taken from fresh rock in shallow artificial pits (0.1e0.5 m deep) in order to diminish effects of weathering and avoid surficial contaminations. They were transported and stored in hermetic plastic bags. From all samples polished thin sections oriented perpendicular to bedding were prepared. Polished thin sections oriented parallel to bedding were prepared from selected samples of OC-rich lithologies. Thin sections were analyzed using transmitted (TLM), reflected (RLM), and fluorescence light microscopy (FLM). Two NIKON ECLIPSE microscopes, LV100 POL and E600 POL, were used for TLM & RLM and FLM, respectively. For fluorescence microscopy a filter BV2A (EX 400-440; DM 455; BA 470) was used. Photomicrographs were taken using NIKON Digital Sight cameras and a NIS Elements 2.30 software. Classification of detrital rocks after Folk (1974) and phosphate rocks after Föllmi et al. (1991) were followed for the description of discerned lithologies. Qualitative maceral evaluations were performed following the ICCP nomenclature described in Taylor et al. (1998), with later extensions and modifications by Boucsein and Stein (2009) and Ercegovac and Kostic (2006). Semiquantitative assessment of the maceral abundance was done by comparing RLM and FLM images of representative thin sections along the Blanknuten and Skarpryttaren sections. The counts of pyrite particles and measurements of pyrite framboid sizes were done under a combined transmitted and reflected light (TLM & RLM) at 500 magnification. This method was found to be much faster than the counts/measurements using backscattered electron images under scanning electron microscope (Wilkin et al., 1996). Comparative analyses using a JEOL JSM-840A scanning microscope and a NIKON ECLIPSE LV100 POL light microscope on three selected thin sections showed no noticeable differences in the obtained pyrite type and framboid size distributions. In each analyzed thin section, the measurements were done on a surface of approximately 1 square centimeter. Wherever possible, pyrite in the OC-rich rock matrix and in the phosphate fraction (nodules, peloids) was analyzed separately. The sizes of pyrite framboids were measured directly from the microscope to the nearest 1 mm. Such analysis tends to underestimate the true

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diameter of framboids, but it was demonstrated that the deviation from true diameter unlikely exceeds 10% (Wilkin et al., 1996; Wignall and Newton, 1998). Ninety samples representing four sections of the Triassic succession (combined MueneReddikeidet, Blanknuten, Skarpryttaren, and Isormen) were selected for x-ray diffraction analysis (XRD), Rock-Eval pyrolysis, and PeFeeS geochemistry (Fig. 2). In sediments with macroscopically discernible phosphate components (nodules, phosphoclasts, peloidal seams, etc.), the macroscopic phosphate fraction (MPF) and the rock matrix (RM) were analyzed separately. This was made possible for all samples showing macroscopic phosphate fraction but one from the Blanknuten section (E4-41A), in which diagenetic impregnation by carbonate minerals prevented successful separation. The samples were crushed, washed three times with distilled water, and dried. Debris of phosphate fraction were hand-picked under binocular microscope in an attempt to avoid marginal zones of nodules, which usually contain an admixture of black shale material. For XRD analysis the samples were ground to <63 mm fraction. Diffraction patterns were recorded on disordered powder samples using a SIGMA 2070 diffractometer and a Bruker AXS D8 Advance diffractometer in the range 2e120 2Q with CoKa radiation. Diffractionel software v. 03/93 was used to process the obtained data. 100 mg portions of crushed samples were analyzed using a RockEval 6 Standard Analyzer. The analytical procedure followed techniques outlined by Behar et al. (2001). In OC-rich lithologies containing phosphate fraction, only the rock matrix was analyzed. The parameters reported embrace: S1 (the amount of free hydrocarbons released during pyrolysis; in mg HC/g rock); S2 (the amount of pyrolysable hydrocarbons generated during pyrolysis; in mg HC/ g rock); Tmax (temperature of maximum generation of pyrolysable hydrocarbons; in  C); S3 (the amount of carbon dioxide generated during pyrolysis; in mg CO2/g rock), TOC (total organic carbon content; in wt.%); S2/S3; PI (Production Index  S1/(S1 þ S2)), HI (Hydrogen Index; in mg HC/g TOC), and OI (Oxygen Index; in mg CO2/g TOC). The Tmax values were used as a maturity indicator. Approximately 10 g of samples were split into three parts for determination of the oxides of major elements, iron speciation, and isotopic analysis of sulfur. The oxides of major elements were determined using an inductively coupled plasma emission spectrometry (ICP-ES) at Acme Analytical Laboratories Ltd. In this paper, the phosphorus content (as wt.% P2O5) and total iron content (as wt.% Fe TOT RM; recalculated from wt.% Fe2O3) are reported. The detection limits for P2O5 and Fe2O3 were 0.01% and 0.04%, respectively. For sedimentary intervals showing macroscopic phosphate fraction, the average phosphorus content was calculated using P2O5 data of the rock matrix and the phosphate fraction. The volumetric content of the macroscopic phosphate fraction was determined on the basis of field measurements on 0.5e1 square meter of vertical rock section (depending on the degree of sediment inhomogeneity), approximated to tens of percent. The following equation was used: P2O5AVERAGE ¼ [P2O5MPF  vol.% MPF þ P2O5RM  (100  vol.% MPF)]2. It should be noted that this calculation gives approximate values. It is based on low precision estimation of the content of the phosphate fraction, and does not account for differences in volumetric weight between the black shale and the phosphate. However, it provides far better estimate of bulk content of phosphorus in different phosphatic and OC-rich lithologies than only the rock matrix and the phosphate fraction values. The P2O5AVERAGE values of rocks containing macroscopic phosphate fraction were used during the construction of geochemical diagrams. The relationship between ferrous and ferric iron in black shale samples was characterized by comparing the amounts of iron bound in pyrite with the amounts of iron liberated from a standardized acid digestion (Canfield et al., 1992). The pyritic sulfur

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(SPYR RM in wt.%) was liberated from 1 g samples by chromium reduction method under nitrogen atmosphere (Canfield et al., 1996). The content of pyritic sulfur was determined by weighing the precipitated Ag2S using a RADWAG WAA 210/C/1 scale. The measurement precision was 0.1 mg, and reproducibility of the results better than 0.1%. The pyritic iron (FePYR RM in wt.%) was calculated from SPYR RM content assuming FeS2 stoichiometric relationship. The acid extractable iron (FeHCl RM in wt.%) was obtained from samples after 1 min boiling in 12 N HCl (Raiswell et al., 1988). Its content was determined by the atomic absorption spectrometry (AAS) using a PHILIPS PU 9100X spectrometer with airacetylene flame. The detection limit for Fe was 0.1 ppm, and reproducibility of the results better than 1.0%. The degree of pyritization of sediment (DOP; Berner, 1970; Berner and Raiswell, 1983; Raiswell et al., 1988) was calculated using the equation: DOPRM ¼ FePYR RM/(FePYR RM þ FeHCl RM). The analytical procedure for determination of isotopic composition of sulfur in pyrite and in SO4 lattice substitution in carbonate fluorapatite (CFA) followed a general scheme described by Lesniak et al. (2003). In samples containing macroscopic phosphate fraction, the pyritic sulfur in the rock matrix (d34SPYR RM) and in the phosphate fraction (d34SPYR MPF) was analyzed separately. The pyritic sulfur was liberated from samples by chromium reduction method, and the evolved H2S was precipitated by reaction with AgNO3 (Canfield et al., 1986; Zaback and Pratt, 1993). The Ag2S precipitate was oxidized to SO2 by heating in the presence of CuO under vacuum (Fritz et al., 1974). The obtained SO2 was purified by cryogenic distillation. The apatitic sulfur (d34SCFA) was determined in thirty five samples of the phosphate fraction and the OC-rich rocks without macroscopic phosphate that revealed elevated content of CFA. The samples were verified to be devoid of microscopically detectable traces of neoformation or recrystallization of CFA (Krajewski, 2000c). Samples were treated sequentially with 5% NaOCl, 1 M NaCl, and 1 M acetate buffer (pH ¼ 5) (Compton et al., 1993; Sakai et al., 1980; see also Goldberg et al., 2011). CFA was dissolved in 2 N HCl, and the sulfate was precipitated as BaSO4 by reaction with BaCl2. SO2 was liberated from a BaSO4/V2O5/SiO2 mixture under vacuum in the presence of Cu (Yanagisawa and Sakai, 1983), and purified by cryogenic distillation. Isotopic 34 32 S/ S ratios in SO2 obtained from pyrite and CFA were determined using a THERMO-FINNIGAN MAT 253 spectrometer. Reproducibility of the results was better than 0.2&. The sulfur isotope results are expressed in & relative to Vienna Cañyon Diablo Troilite (V-CDT) using a standard d notation: d34Ssample ¼ {[(34S/32Ssample)/ (34S/32SV-CDT)]  1}  1000. 4. Results 4.1. Facies 4.1.1. Black shale facies The black shale facies of the lower and middle parts of the Muen Member (up to 50 m thick) is represented by non-laminated mudshale, mudstone, silt-shale, and siltstone. In western Edgeøya, the black shale succession has been divided into four units (Fig. 2). Units 1 and 3 are dominated by homogeneous shale packages, while Units 2 and 4 contain coarsening-upward shale/mudstone packages 1e3 m thick that terminate in silty horizons variably cemented by microcrystalline dolomite (Fig. 3d). These horizons show evidence of intermittent non-deposition including common winnowing and biological reworking. This division becomes less clear northward, though a poor exposition of the lower part of the Botneheia Formation prevents detailed observations. Field and microscopic examinations suggest that the shaly sediment is bioturbated as a rule. Taenidium and Rhizocorallium (Mørk and

Bromley, 2008) were observed in recurrent dolomitic cementstones (Fig. 4e). Flattened trace fossils can also be seen in dolomitic muddy and silty lithologies. The shale without dolomite cement usually shows homogeneous and patchy textures, which might result from fine bioturbation of the sediment. The black shale contains detrital quartz, feldspars and micas (Figs. 5a and 6g,h). Clay minerals are dominated by illite and chlorite, with a varying admixture of kaolinite. Noticeable amounts of ferric iron compounds (XRD detected hematite, and microscopically detected oxide/oxyhydroxide particles) and pyrite occur throughout the succession. The content of detrital feldspars, micas, and ferric iron minerals irregularly decreases upward the succession along with the overall decrease of the content of the silt fraction. Biogenic remnants embrace rare siliceous sponge spicules and arenaceous foraminifera. Reptilian bones were noted in rare carbonate concretions. Flattened imprints of ammonoids are preserved in the dolomitic cementstone beds. 4.1.2. Phosphogenic black shale facies The phosphogenic black shale facies of the upper part of the Muen Member and of the Blanknuten Member (40e60 m thick) embraces several sediment types, which are moderately to strongly enriched in sedimentary phosphate. Most of the lithologies are OCrich and distinctly black in color (Figs. 3 and 4aec). They are represented by mud-shale and mudstone with subordinate siltstone (Fig. 5c). Sandy siltstone and silty sandstone locally occur at top of the phosphogenic succession (Fig. 4f). Quartz grains dominate the silt and sand fractions (Fig. 6a,b,e,f). Feldspars are accessory, observed in the lower and upper parts of the phosphogenic succession. Clay minerals are dominated by illite and chlorite. The contents of detrital micas and kaolinite are substantially lower than in the underlying black shale facies. Hematite does not occur in amounts detected by XRD. Pyrite constitutes a subordinate, but constant admixture. Common molds of radiolaria are accompanied by bivalve shell detritus. Vertebrate bones and their detritus are also common. Rare sponge spicules, and arenaceous and calcareous foraminifera were noted in the lower and upper parts of the phosphogenic succession. Recurrent dolomitic cementstone beds with flattened ammonoid imprints and horizons of large carbonate concretions occur in the succession (Figs. 4d and 5b). The succession has been divided into five units (Units 5e9), out of which three are cliff-forming (Units 6e8) (Figs. 2 and 3). Units 5 and 6 show maximum lithological diversity, with complex packages and beds of black shale and dolomitic mud- to siltstone containing a spectrum of phosphate accumulations. Many black shale packages are finely bioturbated. Taenidium and Thalassinoides burrows were noted at recurrent levels (Fig. 4b,c). The bioturbated packages interfinger with decimeter- to meter-thick (at maximum) intervals of black shale showing discrete lamination (Fig. 4a). Unit 7 is described in Section 4.1.3. Unit 8 shows facies development similar to Units 5 and 6, but its lithological diversity is less pronounced. Remarkable for this unit are recurrent coquinoid shale beds with accumulations of bivalve shells (Fig. 5f) as well as reptilian and fish bone beds (Krajewski, 2008). This unit is at places truncated by basal erosional surface of Unit 9. It is missing in southern Barentsøya (Fig. 2). The basal phosphorite conglomerate of Unit 9 (9A) is recognized in western Edgeøya and southern Barentsøya, forming a condensed deposit of reworked phosphate nodules, phosphatic fossil molds, and bones and bone debris. It is accompanied by phosphatic sandy siltstone and silty sandstone beds, often impregnated by diagenetic dolomite (Figs. 3c and 4f). The youngest deposits of the Botneheia Formation, making the upper part of Unit 9 (9B), consist of black phosphatic shale up to a few meters thick that is disconformably overlain by sideritic shale of the Tschermakfjellet Formation.

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Figure 4. Sedimentary features of the Botneheia Formation in eastern Svalbard. (a) Bioturbated (b) and finely laminated (l) packages of black shale (BSh) separated by a phosphatic bed showing winnowing and reworking features (w/r) in the phosphogenic black shale facies. The bed comprises two layers of phosphatic grainstone (PhG) separated by a layer containing allochthonous phosphate nodules and grainstone seams (PhN). Horizontal and subvertical burrows in the lower black shale package show advanced compactional deformation. The burrows are better preserved in the silty intercallation cemented by microcrystalline dolomite (DS). Hammer is 40 cm long. Unit 6, Muen. (b) Bedding plane of a phosphatic black shale package showing flattened Thalassinoides tunnels with dominant Y-branching. The tunnels are filled with densely packed, phosphatized fecal pellets. Unit 8, Muen. (c) Vertical section through phosphatic black shale with flattened Thalassinoides tunnels. The cartridge is 4 cm long. Unit 8, Muen. (d) Horizontal burrow filled with phosphatic sediment in dolomitic cementstone bed. Thin section photograph. Unit 5, Blanknuten. (e) Heavily bioturbated dolomitic siltstone with Taenidium in the nonphosphogenic black shale facies. Section parallel to bedding. Unit 4, Blanknuten. (f) Non-compressed Thalassinoides tunnels with dominant T-branching in sandy siltstone bed at top of the phosphogenic succession. Section parallel to bedding. Penknife is 11 cm long. Unit 9, Palibinranten. (b, e) Hammer head is 22 cm wide.

In northeastern Barentsøya, the phosphogenic succession can only be divided into two parts (Fig. 2). The lower, cliff-forming part (referred to as Units 6e7) bears mixed characteristics of Units 6 and 7 as recognized in southern Barentsøya and Edgeøya. The upper part (referred to as Units 8e9) is similar to phosphatic black shale of

Unit 9 in the Freemansundet area, though the basal phosphorite conglomerate splits into scattered phosphate nodules in the OCrich sediment. The section at Isormen shows elevated content of glauconite grains in the phosphogenic succession. The glauconite grains are of the sand to coarse silt fraction. They occur scattered in

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Figure 5. Microfacies of the Botneheia Formation in eastern Svalbard. (a) Silt-shale in the non-phosphogenic black shale facies contains common quartz grains (white) and streaks of terrestrial organic matter (dark brown). Detrital micas and feldspars are noticeable components of the sediment. Unit 2, Blanknuten. (b) Dolomitic cementstone at boundary between the non-phosphogenic and phosphogenic black shale facies. Interlocking mosaic of euhedral to anhedral dolomite crystals reflects diagenetic precipitation of the carbonate. Bottom of Unit 5, Blanknuten. (c) Mud-shale in the phosphogenic black shale facies containing silt-sized quartz grains (white) in clay-dominated matrix. Amorphous organic matter is finely distributed in the matrix. Unit 5, Blanknuten. (d) Massive mudstone in the phosphogenic facies showing molds of radiolaria (ra) and Tasmanites (ta). The matrix is reach in amorphous organic matter. Unit 7, Skarpryttaren. (e) Massive mudstone in the phosphogenic facies with common juvenile bivalve shells and rare silt-sized quartz grains (white). Unit 7, Skarpryttaren. (f) Coquinoid black shale in the phosphogenic facies showing accumulations of Daonella shells (white). Unit 8, Muen. TLM & RLM photomicrographs, normal light. Scale in (e) for all photos. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

the shaly sediment, and are also incorporated as detrital components in authigenic phosphate bodies. 4.1.3. Massive phosphatic mudstone The cliff-forming succession of the Blanknuten Member in Edgeøya is tripartite (Fig. 3b), with the middle part made by a

massive phosphatic mudstone of Unit 7 (15e20 m thick). In most sections in eastern Svalbard, the mudstone appears to be devoid of macroscopic phosphate deposits. However, it contains common microscopic phosphate peloids (see Section 4.3). The rock lacks lamination and bioturbation, though it commonly weathers into paper shale. It contains abundant molds of radiolaria and layers

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Figure 6. Mineral composition of the Botneheia Formation in eastern Svalbard. (a, c, e) Phosphogenic black shale facies; (b, d, f) phosphate fraction in the phosphogenic black shale facies; (g, h) non-phosphogenic black shale facies. XRD patterns of representative samples from the Blanknuten section.

dominated by minute shells of juvenile bivalves (Fig. 5d,e). The detrital components are dominated by quartz, and the clay-rich matrix by illite (Fig. 6c). Microcrystalline quartz cement occurs scattered in the OC-rich matrix along with diagenetic dolomite crystals. Pyrite occurs throughout the mudstone succession. There is observed an indistinct southward increase of the content of the silt-sized clastic fraction in the mudstone. Unit 7 thins and splits into several intervals of phosphatic and dolomitic shale with macroscopic phosphate fraction on transect from the southern to northeastern Barentsøya (Fig. 2). 4.1.4. Black shales of the Vikinghøgda and Tschermakfjellet formations The analyzed samples of the Vikinghøgda Formation show microfacies and mineral composition similar to the nonphosphogenic black shale facies of the Botneheia Formation. In contrast, the black shale of the lower part of the Tschermakfjellet Formation is pyrite-poor, but remarkably enriched in siderite cement.

4.2. Organic matter The Botneheia Formation contains both the sedimentary and secondary organic matter. Sedimentary organic matter embraces indigenous (marine) and allochthonous (transported and/or redeposited) fractions. Secondary organic matter comprises organic fractions generated from sedimenatry organic matter during burial and products of their epigenetic transformations. The black shale facies usually contains compressed indigenous organic debris, owing to diagenetic compaction. The phosphogenic black shale facies contains a variety of compressed and noncompressed organic debris due to punctuated preservation of uncompact sediment texture in the nodular phosphatic bodies (Figs. 7 and 8). 4.2.1. Maceral composition Sedimentary organic matter in the Botneheia Formation comprises macerals of the vitrinite, liptinite, and inertinite groups. Vitrinite components are represented by telinite, telocolinite, and

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Figure 7. Macerals in OC-rich rock matrix in the phosphogenic black shale facies. (a) Lamalginite showing compressed planktonic unicells (la) and telocolinite (tc) in massive phosphatic mudstone. Section perpendicular to bedding. Unit 7, Skarpyttaren. (b) Lamalginite showing agglomeration of planktonic unicells in massive phosphatic mudstone. Section parallel to bedding. Unit 7, Skarpyttaren; (c) Telalginite (Tasmanites) in section perpendicular to bedding. Unit 6, Blanknuten. (d) Telalginite (Tasmanites) in section parallel to bedding. Unit 6, Skrukkefjellet. (e) Lamalginite (la), herbamorphinite (hb), and liptodetrinite (ld) in matrix bituminite. Section perpendicular to bedding. Unit 6, Blanknuten. (f) Vitrodetrinite (vd), pyritized in part, and liptodetrinite (ld) in matrix bituminite. Section perpendicular to bedding. Unit 6, Skarpryttaren. FLM photomicrographs, normal light.

colinite particles, usually less than 200 mm in size, as well as by vitrodetrinite (Figs. 7a,f and 8e,f). Colinite and telocolinite particles occur in layers showing herbamorphinite (bituminite III) streaks and lenses (Fig. 7e). There are observed intermediate forms or replacements of (telo)colinite by herbamorphinite, suggesting that herbamorphinite originated from sedimentary decomposition of

the land plant material. All these components show no fluorescence and higher reflectance than the liptinite components. The liptinite components are represented by alginite and liptodetrinite, which usually exhibit bright fluorescence. Alginite comprises lamalginite and telalginite. Observations in vertical and horizontal sections of black shale show that lamalginite consists of

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Figure 8. Macerals in phosphate fraction in the phosphogenic black shale facies. (a) Telalginite (non-compressed algal unicells) in phosphatized zooplanktonic fecal pellets in porous phosphate nodule. White (orangeeyellow in color photo) are pyrite inclusions. Unit 6, Blanknuten. (b) Telalginite (non-compressed algal unicells) in compact phosphate nodule. Unit 6, Skrukkefjellet. (c) Telalginite (non-compressed Tasmanites) in phosphate nodule. Unit 9, Palibinranten. (d) Fragment of wall of Tasmanites showing pore canals filled with bitumen. Unit 9, Palibinranten. (e) Telinite (tl) showing framboidal pyrite in cellular voids in porous phosphate nodule. Unit 6, Blanknuten. (f) Vitrodetrinite (vd) and liptodetrinite (ld) in compact phosphate nodule. Unit 6, Skarpryttaren. (a, e) TLM & RLM photomicrographs, normal light; (c) TLM photomicrograph, normal light; (b, d, f) FLM photomicrographs, normal light. All photos: sections of phosphate nodules perpendicular to bedding.

accumulations of compressed planktonic algal unicells (Fig. 7a,b). This is confirmed by observations of sections of phosphatic matrices that contain the same unicells, but without compressional deformation (Fig. 8a,b). Telalginite is dominated by Tasmanites (Figs. 7c,d and 8c,d). This alga shows thick thalli that usually are

between 100 and 500 mm in size. However, they may exceed 1 mm in extreme cases. Other liptinite particles embrace diverse acritarchs and rare spores of land plants. Liptodetrinite is closely associated with small alginite particles, suggesting its origin from sedimentary fragmentation of algal material (Figs. 7e,f and 8f). It

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commonly occurs in fluorescing matrix bituminite, which is likely to represent extremely fragmented algal material intermixed with clay-sized mineral grains. Inertinite is represented by inertodetrinite showing no fluorescence and very high reflectance. Secondary organic matter occurs both in the form of macroscopic bodies and microscopic inclusions. Liquid bitumen and exudatinite infillings have been observed in cracks and fissures of carbonate concretions and in incompletely cemented fossils and voids. Microscopic bitumen inclusions occur in Tasmanites pore canals (Fig. 8d) as well as in other alginite particles. In overmature sections in SW Edgeøya, exudatinite impregnations of black shale are common. Micrinite is noted there, but usually masked by exudatinite. 4.2.2. Compositional trends The representation of organic components and their proportional contribution depend on the type of sedimentary facies and the maturity level. The content of vitrinite particles is subordinate. It irregularly decreases upward the non-phosphogenic black shale succession. Units 1 and 3 show elevated contents of vitrodetrinite, herbamorphinite, and inertodetrinite, suggesting periods of enhanced supply of land-derived and reworked organic matter in the lower part of the Botneheia Formation. The transition between the non-phosphogenic and phosphogenic black shales shows an increase in the content of alginite particles and liptodetrinite and a decrease in the content of vitrinite particles and herbamorphinite, but without abrupt compositional changes. Units 5 and 6 are dominated by alginite and liptodetrinite occurring in matrix bituminite. The alginite components are at various stages of mechanical fragmentation, suggesting their enhanced sedimentary reworking. Unit 7 contains mostly lamalginite and telalginite with a negligible admixture of vitrinite (mostly telinite and vitrodetrinite). The algal material is seldom fragmented, forming elongated agglomerations of compressed cells. Tasmanites is locally common, though small, thin-walled unicells predominate. Advanced fragmentation of algal material is noted in northeastern Barentsøya, along with the overall increase of dynamic reworking of sediment. An increase in the content of vitrinite particles and herbamorphinite is again observed in Units 8 and 9. Surprisingly enough, the youngest black shale deposits of Unit 9 (9B) show sharp increase in the content of alginite, the large Tasmanites algae in particular. Similar enrichment has been noted at top of the Botneheia Formation in submarine drilling off Kong Karls Land (Vigran et al., 2008).

4.3. Apatite (carbonate fluorapatite, CFA) Phosphate accumulations in the Botneheia Formation owe their origin to precipitation of authigenic phosphate mineral during seafloor history of the OC-rich facies. The phosphatic microfabrics show cloths and clusters of micron-sized, globose to rod-shaped apatitic bodies indicative of rapid pulses of phosphate precipitation at numerous nucleation sites in the sediment (Fig. 9). XRD reveals that the phosphate is invariably CFA (Fig. 6aef), though its ultrastructures suggest precipitation via intermediate, unstable calcium phosphate phase (Krajewski, 2000a). CFA forms grains that vary in size from microscopic peloids to nodules up to several centimeters across. The nodules are usually flat to ovoid in shape (Fig. 10d). Complex morphologies reflecting phosphatization of various sediment inhomogeneities (laminations, organic remains, burrows) are also observed (Fig. 4a,def). 4.3.1. Pristine versus allochthonous phosphate accumulations Two major types of phosphate accumulations can be discerned: pristine phosphates and allochthonous phosphates. Pristine phosphates are autochthonous deposits that experienced one cycle of phosphate precipitation resulting in the formation of phosphate bodies and grains preserved in sediment without any sign of reworking or transport (Fig. 10). Allochthonous phosphates embrace the accumulations that after formation experienced one or more dynamic events of reworking and/or winnowing, with or without lateral transport involved (Fig. 11). There are also mixed phosphate accumulations that developed as a result of more than one episode of phosphate precipitation separated by reworking event(s) (Fig. 12). However, they are less common than the pristine and allochthonous deposits. Phosphatic microbialites, which are peculiar phosphate accumulations related to benthic activity of microbial mats (Krajewski, 2011), have not been recognized in the Botneheia Formation of eastern Svalbard. 4.3.2. Compositional trends Pristine phosphates are distributed throughout the phosphogenic succession, coinciding in various arrangements with allochthonous deposits. Unit 5 shows mostly allochthonous phosphates forming lensoidal seams and lenses in black shale, which are cemented by microcrystalline dolomite. The intervals with allochthonous phosphate seams interfinger with black shale packages containing pristine nodular accumulations. Unit 6 contains more pristine

Figure 9. Authigenic apatite (carbonate fluorapatite, CFA). (a) Organic-rich clusters cemented by globose to rod-shaped CFA particles in nodular phosphatic matrix. Dark spheres are algal unicells. (b) Delicate cloths of globose and rod-shaped CFA particles in pore space of nodular phosphatic matrix. (a, b) The remaining pore space is filled by late diagenetic calcite cement. Unit 8, Palibinranten. TLM photomicrographs, normal light. Scale in (a) for both photos.

Figure 10. Pristine phosphate. (a) Massive phosphatic mudstone containing discrete layers of pristine phosphatic packstone. The mudstone weathers into paper shale. The visible part of the cliff is approximately 10 m high. Unit 7, Skrukkefjellet. (b) Phosphatic packstone layer composed of pristine peloids originated as a result of pulse of rapid phosphate precipitation in OC-rich mud. Note the compressed nature of organic matter between peloids. White are silt-sized quartz grains. Unit 7, Skrukkefjellet. (c) Pristine phosphate peloids in black phosphatic shale showing common amorphous organic matter (dark). Unit 6, Blanknuten. (d) Phosphatic black shale showing numerous pristine phosphate nodules of flat to ovoid morphology. Note the recurrent horizons enriched in phosphate nodules. Thin, elongated phosphatic bodies in the lower and middle parts of the photo are sections of flattened Thalassinoides tunnels. Hammer is 40 cm long. Unit 6, Blanknuten. (e) Vertical section of a pristine phosphate nodule from sedimentary interval shown in (d). The nodule shows its central part preserving the original texture of radiolaria-rich sediment and mechanically-oriented marginal parts. Arrow indicates cross section of a small burrow filled with fecal pellets of the burrowing organism. (f) Matrix of the phosphate nodule shown in (e). Note the common molds of radiolaria (white spots) and abundant, phosphatized zooplanktonic fecal pellets. The remaining pore space is filled by late diagenetic calcite cement. (b, c, f) TLM photomicrographs, normal light; (e) thin section photograph.

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Figure 11. Allochthonous phosphate. (a) Beds of allochthonous phosphatic grainstone (white) in massive phosphatic mudstone. Hammer is 40 cm long. Unit 7, Reddikeidet. (b) Allochthonous phosphatic grainstone composed of reworked peloids. Inter-particle pore space is filled with late diagenetic calcite cement. Unit 7, Blanknuten. TLM photomicrograph, normal light. (c) Seams and lenses of allochthonous phosphatic grainstone in dolomitic shale. Some of the lenses show low-angle cross bedding and ripple lamination. Unit 6, Blanknuten. (d) Allochthonous phosphatic grainstone composed of reworked peloids showing “oolitic” zonation. Units 6e7, Isormen. TLM photomicrograph, normal light. (e) Dolomitic sandy siltstone with allochthonous phosphate nodules. Hammer is 45 cm long. Unit 9, Palibinranten. (f) Horizontal section through allochthonous phosphate nodule from the siltstone bed shown in (e). The nodule shows two generations of growth in sediment of different texture separated by erosional surface. Thin section photograph.

nodules in undisturbed black shale packages 0.5e2 m thick (Fig.10d), which are separated by phosphatic grainstone beds up to 0.3 m thick (Fig. 11c). The latter beds developed as a result of several episodes of reworking and transport. Phosphatic layers and lenses in the grainstone show low-angle cross bedding and ripple lamination. Allochthonous peloids exhibit at places “ooidal” structures indicative of recurrent events of phosphate accretion and reworking (Fig. 11d).

Traces of sedimentary reworking diminish upward as the black shale with macroscopic phosphate fraction passes into the massive phosphatic mudstone of Unit 7. This muddy sediment contains discrete layers (1e25 mm thick) of pristine phosphatic packstone composed of densely packed, sand- to silt-sized peloids (Fig. 10aec). The mudstone succession shows the presence of a few allochthonous phosphatic grainstone beds (Fig. 11a). Two such beds have been

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Figure 12. Mixed phosphate. (a) Succession of phosphate-cemented, allochthonous grainstone layers and lenses (white) in black phosphatic shale. Penknife is 11 cm long. Units 6e 7, Isormen. (b) Allochthonous phosphate peloids of at least two generations (pale and dark) and glauconite grains (g) cemented by fringing phosphate cement in grainstone layer shown in (a). The fringing cement thickens at places to form void-filling cement. White is late diagenetic calcite cement. TLM photomicrograph, normal light.

recognized at Skarpryttaren and Blanknuten, which are border points of the area of best development of Unit 7. This number increases to three and seven at Muen and Reddikeidet, respectively. Allochthonous peloids composing the grainstone beds show traces of crushing and abrasion, though no “ooidal” structures were observed (Fig. 11b). In Unit 8 there again appear common macroscopic phosphate bodies in black shale. They are dominated by pristine nodules, either scattered in shaly intervals or concentrated in horizons enriched in bone material. Allochthonous phosphates are seldom observed in this unit. The phosphorite conglomerate of Unit 9 (9A) consists of allochthonous phosphate nodules showing generations of growth separated by erosional surfaces (Fig. 11e,f). In northeastern Barentsøya, the succession contains various mixed phosphate accumulations showing the presence of glauconite grains. They are interbedded at irregular intervals with OC-rich and dolomitic shale packages and layers (Fig. 12).

material usually was replaced by ultramicrocrystalline CFA. Perfect replacement microstructures coated by CFA fringe cements suggest that this process occurred very early in sediment during the authigenic phosphate precipitation. 4.4. Pyrite Pyrite in the Botneheia Formation occurs in several different forms, which reflect the mineral formation during sedimentogenesis, authigenesis, and diagenesis of the OC-rich facies. The pyrite is represented by framboids and polyframboidal aggregates, crystals and grains and their aggregates (microgranular pyrite), organic and mineral replacement structures, and cement inclusions (Figs. 13 and 14). Proportional contribution of these forms varies along the succession. It is in part facies dependent, and usually different in the rock matrix and in the phosphate fraction of the phosphogenic shale facies.

4.3.3. Sediment textures in phosphate bodies Pristine phosphate bodies, larger nodules in particular, provide insight into the original composition and texture of the sediment at the time of phosphate authigenesis. Outside the nodules, advanced compaction of the sediment and dissolution of biogenic mineral phases obliterated the primary textures. Observations in thin sections reveal that most of the nodular bodies contain abundant fecal pellets impregnated and cemented by phosphate (Fig. 10e,f). The vast majority of them are within a narrow size range (50e250 mm) and contain algal unicells at various stages of degradation (Fig. 8a,b). These pellets are interpreted to be of zooplanktonic origin. They are accompanied by common molds of radiolaria. Larger, ovoidal to elongated fecal pellets (0.5 mm up to a few milimeters) can be found in phosphatic infillings of burrows and in some phosphate nodules (Fig. 4b,c). They are likely to represent phosphatized feces of the benthic infauna.

4.4.1. Framboids Two types of framboids were identified in the Botneheia Formation (Fig. 13a,b). They are similar to the types discerned by Wignall and Newton (1998), and referred to as type 1 and type 2 framboids. The type 1 framboids (2e25 mm, rarely larger) show compact spheres composed of densely packed, unidimensional and equimorphic microcrystals. The type 2 framboids are usually larger (6e50 mm), and composed of loose aggregates of microcrystals of a greater range of sizes. These latter framboids are commonly compacted in sediment and/or disintegrated to make more or less irregular concentrations of microcrystals. They are far less common than the type 1 framboids. Most framboids occur scattered in sediment, though they are also concentrated in the form of polyframboidal aggregates (Fig. 13c,d). These aggregates are associated with larger organic remains or constitute infillings of algal unicells.

4.3.4. Biogenic phosphate Biogenic phosphate in the form of reptilian and fish bones and bone debris is common throughout the phosphogenic part of the Botneheia Formation. The reptilian and fish remains are scattered in the shaly sediment. They are also concentrated at recurrent horizons to form bone beds containing fossil skeletons at various stages of mechanical disintegration (Krajewski, 2008). Microscopic and XRD investigations show that the original bone

4.4.2. Crystals and grains (microgranular pyrite) Pyrite crystals vary in size from parts of micrometer to several micrometers (Fig. 13e). They exhibit octahedral, pyritohedral, and rarely cubic habit. Anhedral to subhedral pyrite grains of similar sizes occur associated with the crystals in black shale sediment. At places, the crystals and grains are concentrated to form indistinct streaks, usually associated with seams enriched in organic matter. In rare cases, the microgranular aggregates form infillings of algal unicells

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Figure 13. Pyrite (I). (a) Type 1 framboids (compact spheres) and type 2 framboids (larger loose aggregates) in pore space of dolomitic black shale. Units 6e7, Isormen. (b) Type 1 framboids and disintegrated type 2 framboids due to a growth of dolomite crystals in pore space of dolomitic black shale. Units 6e7, Isormen. (c) Polyframboidal aggregate composed of type 1 framboids in pore space of dolomitic black shale. Units 6e7, Isormen. (d) Polyframboidal aggregate composed of type 1 framboids and one large type 2 framboid filling algal unicell. Unit 9, Palibinranten. (e) Scattered pyrite crystals (octahedral to pyritohedral) in massive phosphatic mudstone. Unit 7, Blanknuten. (f) Aggregate of pyrite crystals (octahedral to pyritohedral) filling algal unicell in black shale. Unit 4, Blanknuten. (aef) TLM & RLM photomicrographs, normal light.

and microfossils (Fig. 13f). Microgranular pyrite in the phosphate fraction shows a wider range of sizes and usually more complex morphology resulting from aggregation of the crystals (Fig. 14a). 4.4.3. Replaced organic debris Pyritized organic remains are represented mostly by vitrinite particles (colinite, telocolinite) and herbamorphinite (Fig. 14b),

while indigenous liptinite particles (tel- and lamalginite) are not pyritized as a rule. Exceptions are pyritic infillings of some algal unicells (Fig. 13d,f), but the cell walls show no replacement structures. Pyritization of organic particles and streaks resulted in the formation of replacement structures composed of (poly)framboidal bodies or aggregates of microgranular pyrite or various combinations of the two.

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Figure 14. Pyrite (II). (a) Minute aggregates of subhedral pyrite crystals scattered in matrix of phosphate nodule. Unit 6, Blanknuten. (b) Aggregates of anhedral to subhedral pyrite crystals replacing colinite particles in black shale. Unit 3, Blanknuten. (c) Pyrite-replaced siliceous sponge spicule in phosphatic grainstone. Unit 6, Blanknuten. (d) Aggregates of euhedral, anhedral and subhedral pyrite crystals replacing shells of juvenile bivalves in a matrix of phosphate nodule. Unit 9, Skarpryttaren. (e) Incipient (left) and advanced (right) stages of replacement of dolomite rhombs by skeletal pyrite crystals in massive phosphatic mudstone. Unit 7, Blanknuten. (f) Aggregate of skeletal pyrite crystals forming cement inclusions in pore space of phosphatic grainstone. Note pyritohedral to cubic crystal faces exposed toward open pore space. This pore space is filled by late diagenetic calcite cement. Units 6e7, Isormen. (aef) TLM & RLM photomicrographs, normal light.

4.4.4. Skeletal to massive pyrite Skeletal to massive pyrite forms replacement structures and cement inclusions. Some debris of siliceous sponge spicules is replaced by pyrite (Fig. 14c). Similar structures in bivalve shells are observed, though they show incipient or incomplete replacements (Fig. 14d). Replacement structures in diagenetic dolomite crystals

consist of skeletal pyrite showing well-defined crystal faces (Fig. 14e). Skeletal crystals and their aggregates are also observed in the pore space of phosphate nodules, in radiolaria molds, and between allochthonous peloids in grain-supported matrices. These aggregates usually contain large crystals (up to 50 mm), which are subhedral at contacts with the direct substratum and show well-

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defined faces exposed toward the pore space (Fig. 14f). The pore space that remained after formation of the pyrite is filled by blocky calcite cement containing bitumen inclusions and exudatinite linings in incompletely cemented voids. 4.4.5. Compositional trends Pyrite framboids of type 1 dominate the pyrite fraction in black shale sediment of the non-phosphogenic facies (Units 1e4) and in phosphatic black shale intervals of the phosphogenic facies (Units 5e6, and 8e9) (Figs. 15e18). They constitute between 40% and 70% of total pyrite counts in the rock matrix. The type 2 framboids are far less common, attaining 10% at some levels of the succession. The size distributions of type 1 framboids in the black shale facies show broad size range with mean diameter between 9 mm and 10 mm

(Fig. 15c,d). This range becomes narrower upward the phosphogenic succession (Units 5e6), with mean diameter decreasing to approximately 7 mm (Figs. 15a, b and 16f, g). Similar framboid size distributions characterize the upper part of the phosphogenic succession (Units 8e9). The massive phosphatic mudstone of Unit 7 shows a prominent drop in the content of framboids to less than 20%, associated with narrowing of their size distributions and mean diameter centered around 5 mm (Figs. 16bee and 17c,d). Except the Isormen section, there is observed a distinct difference in the distributions of pyrite types and their abundances in the rock matrix and in the phosphate fraction (Figs. 15e18). The phosphate fraction shows a clear predominance of microgranular pyrite (usually 40e70% of total pyrite counts) over other forms of pyrite. There are many phosphate nodules containing almost

Figure 15. Pyrite type distribution (right) and framboid size distribution (left) diagrams of selected samples from Units 4 (c, d) and 5 (a, b) along the Blanknuten section. The framboid size distributions reflect scattered type 1 framboids. The boundary between Units 4 and 5 is the boundary between the non-phosphogenic and phosphogenic black shale facies. These units record deterioration of oxic bottom conditions during the early to late transgressive phases of the Middle Triassic cycle, related to increase of productivity on a transition from distal prodelta to open shelf depositional settings.

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Figure 16. Pyrite type distribution (right) and framboid size distribution (left) diagrams of selected samples from Units 6 to 8 (aeg) along the Blanknuten section. The framboid size distributions reflect scattered type 1 framboids. Unit 7 is dominated by massive phosphatic mudstone (b, d, e) deposited under euxinic conditions during high-stand phase of the Middle Triassic cycle. It contains a few phosphatic grainstone beds (c) that mark intermittent high-energy events. The under- (f, g) and overlying (a) phosphatic black shales were deposited under dominating oxic-to-dysoxic conditions during the late transgressive and regressive phases of the cycle, respectively. For explanations see Figure 15.

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Figure 17. Pyrite type distribution (right) and framboid size distribution (left) diagrams of selected samples from Units 7 and 9 (aed) along the Skarpryttaren section. The framboid size distributions reflect scattered type 1 framboids. Unit 7 is dominated by massive phosphatic mudstone (c, d) deposited under euxinic conditions during high-stand phase of the Middle Triassic cycle. It is disconformably overlain by phosphorite conglomerate and phosphatic shale of Unit 9 (a, b), deposited under oxic conditions during termination of the regressive phase. For explanations see Figure 15.

exclusively pyrite crystals and grains (Fig. 19). The massive phosphatic mudstone of Unit 7 shows OC-rich matrix dominated by microgranular pyrite (60e70% of total pyrite counts). In phosphatic black shale packages at Isormen (Units 6e7), the proportional contribution of framboids increases and their size distributions become similar to the ones noted in the phosphatic black shales elsewhere in eastern Svalbard (Fig. 18). In this succession, framboids occurring in the rock matrix and in the mixed-type phosphate deposits show similar size distributions. However, a remarkable difference in number of framboids between the rock matrix and the phosphate fraction reflects mechanical enrichment in the rock matrix due to compaction. Pyrite replacements of organic particles provide subordinate component of total pyrite present in the Botneheia Formation. In rare cases, they constitute up to 10% of total pyrite counts in the non-phosphogenic black shale facies (Fig. 15c,d). The contribution of skeletal to massive pyrite (replacements in biogenic mineral debris and dolomite crystals, and cement inclusions) to the total pyrite is variable, though subordinate. It attains 20% of total pyrite counts in the phosphogenic succession at Skarpryttaren (Fig. 17). It drops to 10% and less in the remainder of eastern Svalbard (Figs. 15, 16, 18), except parts of the overmature sections in SW Edgeøya.

4.5. Rock-Eval data1 The results of Rock-Eval analysis (Fig. 20) confirm earlier observations that the Botneheia Formation in SW Edgeøya (MueneReddikeidet sections) is overmature with respect to oil generation (Bjorøy et al., 2006; Brekke et al., 2013; Mørk and Bjorøy, 1984), which reflects proximity to extensive Cretaceous dolerite intrusions (Nejbert et al., 2011). Elsewhere in eastern Svalbard, the background maturity of the Botneheia Formation falls within the oil window, though with a decreasing trend northeastward in Barentsøya. At Isormen, the Botneheia Formation occurs at the upper margin of the oil window. The non-phosphogenic black shale facies shows pyrolytic characteristics typical of mixed kerogen type III and II. The phosphogenic black shale facies shows domination of kerogen type II. The massive phosphatic mudstone of Unit 7 contains kerogen of mixed type II and I. Organic matter in the upper part of the Vikinghøgda Formation shows characteristics similar to the non-phosphogenic black shale facies of the Muen Member. The lower part of the Tschermakfjellet Formation is dominated by kerogen type III.

1

see Appendix A for the Rock-Eval data.

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Figure 18. Pyrite type distribution (right) and framboid size distribution (left) diagrams of selected samples from Units 6e7 (a, b) along the Isormen section. The framboid size distributions reflect scattered type 1 framboids. This phosphogenic succession was deposited under oxic-to-dysoxic conditions on a submarine swell during the late transgressive and high-stand phases of the Middle Triassic cycle. For explanations see Figure 15.

Vertical hydrogen index (HI) profiles in the Botneheia Formation correlate with the type of sedimentary facies (Fig. 21). The nonphosphogenic black shale facies shows irregular upward increase of HI values from approximately 150 mg HC/g TOC at bottom to approximately 400 mg HC/g TOC at top. Transition toward the phosphogenic black shale facies shows no abrupt change in the hydrogen richness of organic matter. The phosphatic black shale intervals of Units 5 and 6 have HI values usually greater than 400 mg HC/g TOC. HI attains a maximum of more than 600 mg HC/ g TOC in the massive phosphatic mudstone of Unit 7. The greatest HI values characterize sections in Barentsøya (including Units 6e7 at Isormen), reflecting lower levels of organic maturity compared to western Edgeøya. OC-rich lithologies of Units 8 and 9 show a drop of HI values back to 400e500 mg HC/g TOC. Organic matter of the phosphogenic facies is very poor in oxygen (OI values usually smaller than 10 mg CO2/g TOC), except the reworked intervals with allochthonous phosphates. In the non-phosphogenic part of the succession, an increase of the OI values in Units 1 and 3 is correlated with elevated content of terrestrial and redeposited organic fractions. Flat HI profiles with values fluctuating around 200 mg HC/ g TOC in the phosphogenic facies at Muen and around 100 mg HC/ g TOC at Redddikeidet reflect the presence of overmature organic matter (Brekke et al., 2013). 4.6. PeFeeS geochemistry2 4.6.1. Phosphorus versus organic carbon The non-phosphogenic and phosphogenic black shale facies show different ranges of bulk phosphorus content, with a clear-cut lower boundary of the phosphogenic sediments at 2% P2O5 (Figs. 21 and 22). The phosphorus content in the non-phosphogenic facies (<1% P2O5) reflects a subordinate admixture of scattered authigenic and detrital apatite microcrystals and iron-related phosphate, with some debris of biogenic apatite. This content does not correlate to organic carbon, the latter shows irregular increase upward the succession from ca. 1% to 4% TOC. The phosphorus content in the phosphogenic facies is first of all related to authigenic CFA. The

2

see Appendix B for the geochemical data.

phosphogenic facies is in general more organic-rich and contains organic matter with higher hydrocarbon potential than the nonphosphogenic facies. Units 5 and 6, which show alternating packages of phosphatic black shale and beds of allochthonous phosphate, have the greatest phosphorus values (up to 25% P2O5), but varying carbon values (2e10% TOC). A drop in the phosphorus content in Unit 7 (2e8% P2O5 in phosphatic mudstone) is associated with high content of organic carbon (7e10% TOC) showing the greatest HI values. An exception is thin phosphatic grainstone beds that concentrate allochthonous phosphate peloids (1% TOC; 17e21% P2O5). Units 8 and 9 show an increase of the phosphorus content (up to 25% P2O5 in phosphorite conglomerate; Unit 9A) and a drop of the organic carbon content (4% TOC). The youngest phosphatic black shale preserved in the Botneheia Formation (Unit 9B) records an organic-rich event at the end of the Middle Triassic cycle. This shale has up to 11% TOC, though it shows HI values smaller than the ones noted in the massive phosphatic mudstone. 4.6.2. C/S systematics Figure 23 shows a plot of pyritic sulfur versus organic carbon values for eastern Svalbard. The flat pattern of data indicates that there was no correlation between the formation of pyrite and the availability of organic matter in the non-phosphogenic and phosphogenic black shale facies. These facies contain no more than 2% pyritic sulfur over a broad value range of organic carbon. Slightly elevated sulfur values for the non-phosphogenic black shale facies are associated with an elevated content of total iron in the lower part of the Botneheia Formation (Fig. 24a). The upper part of the Vikinghøgda Formation and the lower part of the Tschermakfjellet Formation show values within the range to the non-phosphogenic facies of the Muen Member. However, the Tschermakfjellet Formation is pyrite- and organic-poor. 4.6.3. Degree of pyritization (DOP) The degree of pyritization of iron (DOP) in the Botneheia Formation of eastern Svalbard varies from 0.08 to 0.95. The obtained DOP values reflect total pyrite present in samples, which encompasses a late diagenetic pyrite fraction. Inasmuch as this fraction constitutes a subordinate admixture, the observed DOP profiles correlate well with the sedimentary facies (Fig. 21). Units 1e4, 5e6,

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Figure 19. Contrasted developments of pyrite in the phosphatic black shale of Unit 8 (aec; sample E4-45) and in the underlying massive phosphatic mudstone of Unit 7 (def; sample E4-42) in the Blanknuten section. These units were deposited under dominating oxic-to-dysoxic and euxinic conditions, respectively. (a) Phosphatic black shale (PBSh) hosting pristine phosphate nodule (PhN). (b) Pyrite framboids of types 1 and 2, crystals and grains, and replacement structure in dolomite crystal (lower left) in black shale shown in (a). (c) Numerous minute pyrite crystals and their aggregates in phosphate nodule shown in (a). (d) Massive phosphatic mudstone (PMd) showing numerous calcite-filled molds of radiolaria (white spots). (e) Pyrite crystals in matrix of the mudstone shown in (d). (f) Large skeletal pyrite crystal and a small type 1 framboid in calcite-cemented mold of radiolaria in mudstone shown in (d). (a, d) photographs of thin sections perpendicular to bedding; (b, c, e, f) TLM & RLM photomicrographs, normal light. Scales in (c) and (f) are for (b) and (e), respectively.

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Figure 20. Plots of Rock-Eval hydrogen index (HI) versus oxygen index (OI) (a) and hydrogen index (HI) versus Tmax (b) against mean evolution paths of kerogen types I, II, and III. Note the overmature samples from the MueneReddikeidet sections in SW Edgeøya, including data field from Brekke et al. (2013).

and 7 show the DOP value ranges of 0.18e0.52, 0.38e0.66, and 0.77e0.95, respectively. A sharp increase of DOP is noted at lower boundary of the massive phosphatic mudstone of Unit 7. It is associated with a decrease of total iron content to less than 2% (Fig. 24a). A decrease of DOP in the overlying Unit 8 is noticeable (0.69e0.75). It is followed by a further drop of DOP to less than 0.2 in Unit 9. A rather narrow DOP value range is observed in the phosphogenic succession at Isormen (0.39e0.56). Small DOP values (0.01e0.33) characterize the overlying Tschermakfjellet Formation. The Vikinghøgda Formations shows values within the range of the non-phosphogenic facies of the Muen Member. 4.6.4. Isotopic composition of pyritic sulfur The isotopic composition of pyritic sulfur reflects mixtures of different types of pyrite occurring in sediment (Figs. 15e18). The obtained d34SPYR value range for the Botneheia Formation of eastern Svalbard is wide (40 to þ15&), reflecting changing distribution of the pyrite types and their likely isotopic compositions (Fig. 24b). The black shales of Units 1e4 show a rather narrow d34SPYR value range between 15& and 1&, reflecting domination of pyrite framboids. The upper part of the Vikinghøgda Formation shows a slightly broader value range (15 to þ3&), reflecting wider variations in the distribution of pyrite types. The phosphatic black shales of Units 5 and 6 show a d34SPYR RM value range extended toward lighter isotopic compositions (20 to þ2&). However, two samples from the overmature section at Muen

considerably straddle this range (31& and þ8&). A systematic shift of d34S toward heavier isotopic compositions from the rock matrix (d34SPYR RM) to the phosphate fraction (d34SPYR MPF) reflects a superimposed change in the proportional contribution of framboidal versus microgranular pyrite (Figs. 21 and 24b). The massive phosphatic mudstone of Unit 7 shows d34SPYR RM values between 0& and þ15&. These are the heaviest values noted in OC-rich sediments of the Botneheia Formation, but comparable to the ones found in the phosphate fraction of the under- and overlying phosphatic black shales. They are consistent with a clear domination of microgranular pyrite in the massive mudstone and in the phosphate fraction of the shales (Fig. 19). Exceptional situation exists in a facies equivalent of Unit 7 in northeastern Barentsøya (Units 6e7 at Isormen), where there are noted comparable d34SPYR values of the rock matrix and the mixed-type phosphate deposits (Fig. 24b). These values reflect similar distributions of pyrite types in the host shale and in the phosphate fraction due to recurrent events of phosphogenesis, pyrite precipitation and reworking (Fig. 18). The most negative d34SPYR RM values (40 to 20&) have been obtained for black phosphatic shale of the uppermost part of the Botneheia Formation (Unit 9 and Units 8e9 at Isormen). They are associated with a clear domination of framboids among other forms of pyrite (Fig. 17a). Similarly suppressed d34SPYR values have been noted in the lowermost part of the overlying shale of the Tschermakfjellet Formation (Fig. 21). One exceptionally heavy value obtained from the Blanknuten section (þ9.8&) seems to reflect the presence of late diagenetic pyrite micronodules.

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Figure 21. TOC, P2O5, HI, OI, DOP, d34SRM (rock matrix), d34SMPF (macroscopic phosphate fraction), and d34SCFA (carbonate fluorapatite) profiles in sections of the Botneheia Formation in eastern Svalbard. Selected Rock-Eval data (TOC, HI, OI, Tmax) for parts of the MueneReddikeidet sections from Brekke et al. (2013).

4.6.5. Isotopic composition of apatitic sulfur The obtained d34S values of CFA in the Botneheia Formation are between þ8& and þ20& (Fig. 25). However, 80% of the results fall in a narrow range between þ14& and þ19&. The d34SCFA profiles in the phosphogenic succession are flat, compared to the d34SPYR profiles (Fig. 21). Light isotopic compositions of apatitic sulfur have been revealed in two samples from Units 5 and 8 at Blanknuten

(þ10& and þ8&, respectively) and in one sample from Units 6e7 at Isormen (þ11&). 5. Discussion The integrated petrographic and geochemical study of organic matter, apatite (CFA), and pyrite in the Botneheia Formation

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Figure 22. Plots of total organic carbon (TOC) versus P2O5 content (a) and Rock-Eval hydrogen index (HI) versus P2O5 content (b).

provides useful tools in reconstructing processes that were instrumental in the deposition of the OC-rich facies in Svalbard. These processes can be delineated in terms of redox conditions versus biological productivity during consequent stages of the Middle Triassic transgressiveeregressive cycle.

Figure 23. Plot of pyritic sulfur in rock matrix (S

5.1. Redox conditions The dominant redox conditions in OC-rich facies of the Botneheia Formation are reconstructed using the analysis of pyrite types and their distributions (Wignall and Newton, 1998; Wignall et al.,

PYR RM)

versus total organic carbon (TOC) in rock matrix.

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Figure 24. Plots of the degree of pyritization of rock matrix (DOP) versus total Fe content (FeTOT d34SRM (rock matrix) and d34SMPF (macroscopic phosphate fraction) (b).

2005; Wilkin et al., 1996), C/S (Leventhal, 1987; Lyons and Berner, 1992; Raiswell and Berner, 1985), DOP (Berner, 1970; Canfield et al., 1996; Raiswell and Canfield, 1998), and isotopic composition of pyritic sulfur (Anderson et al., 1987; Beier and Hayes, 1989; Raiswell, 1997). These data are combined with sedimentologic criteria commonly used in discriminating different levels of sea bottom oxygenation (Allison et al., 1995; Schieber, 2003; Wignall and Myers, 1988). The observed pyrite types in the Botneheia Formation indicate the mineral formation over a range of sedimentary, authigenic, and

RM)

in rock matrix (a) and the degree of pyritization of rock matrix (DOP) versus

diagenetic environments, including a late diagenetic phase (Figs. 13, 14 and 19). This late pyrite correlates with stages of the formation and corrosion of burial dolomite cement. The spike formation of the dolomite cement has been assigned to thermal decarboxylation processes and cracking of kerogen (Brekke et al., 2013; Krajewski and Wo zny, 2009), suggesting that the late pyrite formed in a catagenic environment, most probably as a result of thermochemical sulfate reduction. Isotopic composition of sulfur in pyrite originated as a result of thermochemical sulfate reduction usually is much heavier than the composition of sulfur in pyrite developed

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Figure 25. Plot of the degree of pyritization of rock matrix (DOP) versus d34SCFA (carbonate fluorapatite).

due to bacterial sulfate reduction (Machel et al., 1995; Worden et al., 2000, 2003). Inasmuch as the late pyrite constitutes a subordinate component of total pyrite in the analyzed samples (Figs. 15e18), its effect on the obtained isotopic results should be insignificant. All the analyzed samples show d34SPYR values lower than the supposed isotopic composition of coeval marine sulfate (Fig. 24b), which is typical of domination of bacterial sulfate reduction in the pyrite formation (Brunner and Bernasconi, 2005; Chambers and Trudinger, 1979; Habicht and Canfield, 1997). The isotopic composition of the Middle Triassic marine sulfate has been here accepted at ca. þ18& related to V-CDT using the data of Kampschulte and Strauss (2004), i.e. at a level approximately 3& heavier that the one assumed during a pilot isotopic survey in Svalbard (Krajewski, 2000b). A flattened part of the oceanic isotopic sulfur curve was suggested for the Middle Triassic (Berner, 2005), allowing the assumption of uniform isotopic composition of marine sulfate throughout the deposition of the OC-rich facies. In modern environments, pyrite framboids form close to the redox boundary under conditions of supersaturation with respect to iron monosulfides as a result of intense bacterial sulfate and iron reduction (Rickard, 1997; Wilkin and Barnes, 1996, 1997). Depending on redox conditions, this boundary may be located in the water column (euxinic conditions), at the water/sediment interface (anoxic conditions) or below the interface in surficial sediment layer (shallowing from oxic to dysoxic conditions) (Canfield et al., 1996; Wilkin and Arthur, 2001; Wilkin et al., 1997). Since the dominant parts of the non-phosphogenic and phosphogenic black shale facies of the Botneheia Formation show evidence of metazoan life at the sea bottom (from bioturbationally homogenized sediment to shallow-dwelling and tunnel-forming infauna), it seems clear that this boundary was located within surficial sediment during prevailing periods of the depositional history (Fig. 4). Rare microfossils of benthic organisms (siliceous sponge, foraminifera) noted in Units 1e4, 5e6, and 8e9 confirm this observation (Fig. 14c). The black shales of these units show a predominance of pyrite framboids over other forms of pyrite, suggesting authigenic pyrite formation close to the water/sediment interface from supersaturated pore fluids. The framboid size

distributions in the non-phosphogenic and phosphogenic facies (except Unit 7) suggest a domination of oxic and dysoxic bottom conditions, respectively (Figs. 15e18). Isotopic composition of pyritic sulfur in framboid-dominated sediments points to a sulfatereplete (open) system of bacterial sulfate reduction, with the likely sulfur fractionation in a range between 15& and 55& (Figs. 21 and 24b). These values are within the range frequently observed in marine sediments with dominating bacterial sulfate reduction (Böttcher and Lepland, 2000; Goldhaber and Kaplan, 1975; Passier et al., 1997). The scatter of the isotopic results may be explained by changing rates of bacterial sulfate reduction and the associated reactions of intermediate sulfur species related to changing sedimentation rate, supply of organic matter, and redox conditions (Böttcher et al., 2001; Canfield and Thamdrup, 1994; Habicht and Canfield, 2001; Habicht et al., 1998). It is commonly accepted that the DOP values of 0.45 and 0.75 mark rough boundaries between oxic/dysoxic and dysoxic/anoxic (euxinic) regimes, respectively (Anderson and Raiswell, 2004; Canfield et al., 1996; Raiswell et al., 1988). The ranges between these values include most of the results obtained from Units 1e4, 5e6 and 8, and 7, respectively (Fig. 24). It should be noted, however, that the DOP values of Units 5e6 and 8 have been obtained from black shale intervals, which interfinger with reworked horizons enriched in allochthonous phosphates indicating periods of better ventilation of the sea bottom. Hence, the oxic and oxic-to-dysoxic redox regimes seem to have dominated deposition of the major parts of the non-phosphogenic and phosphogenic facies, respectively. This does not preclude a possibility of local development of anoxic conditions in laminated packages of black shale in Units 5e6 and 8. However, the fine lamination observed over decimeter-scale intervals along the analyzed sections in Edgeøya and Barentsøya (Fig. 4a) is not supported by “anoxic” DOP values (Fig. 21), suggesting episodes of suboxic (denitrifying) conditions rather than the sulphidic bottoms. A rapid change of the content and size distribution of pyrite framboids in Unit 7 in Edgeøya and southern Barentsøya is interpreted to reflect development of euxinic conditions (Figs. 16 and 17). These framboids are dominated by syngenetic framboids, i.e. the ones that formed in the water column and fell down to the

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bottom. However, their contribution to total pyrite in the mudstone is subordinate. This stands in contrast to modern sediments of euxinic basins, which show concentration of sedimentary framboids as a result of import of surplus dissolved iron to the euxinic water column (Anderson and Raiswell, 2004; Lyons and Severmann, 2006; Raiswell and Anderson, 2005; Raiswell et al., 2001). In the Botneheia Formation, the shift of a major phase of pyrite formation from framboids in the oxic-to-dysoxic phosphogenic sediments to microgranular pyrite in the euxinic muddy sediment correlates with a drop in the content of lithogenic iron and with a sharp increase of DOP and d34SPYR RM values (Fig. 24). This suggests that the pyrite formation in the euxinic water column was of low efficiency due to a shortage of dissolved ferrous iron, and that the pyrite formation in the sediment below became limited by the availability of reactive iron. The supposition is consistent with rapid sea-level rise associated with mineral starvation and diminished supply of reactive land-derived iron to the open shelf environment. Slightly elevated iron and sulfur values for the non-phosphogenic black shale facies occurring at base of the Botneheia Formation (Units 1e4) can be related to increased supply of lithogenic iron to a distal prodelta environment, resulting in a higher availability of reactive iron for pyrite formation during the early transgressive phase (Figs. 23 and 24). Most of microgranular pyrite in the Botneheia Formation reflects direct pyrite precipitation at lower levels of saturation that extended down the sediment column, though still within the limits of bacterial sulfate reduction. The precipitation continued throughout at least partial reduction of porosity of the OC-rich facies. It was more efficient in uncompact matrices of the phosphate fraction that in the host shale sediment. This can be demonstrated by comparing distributions of pyrite types and their abundances in the rock matrix and in the phosphate fraction of the phosphatic black shales (Figs. 15e18). Microgranular pyrite in the massive mudstone of Unit 7 embraces crystals and grains that usually are larger than the ones in rock matrix of the under- and overlying phosphatic black shales (Fig. 19). This is consistent with the suggested very low sedimentation rates for this unit that allowed a prolonged growth of pyrite in uncompact muddy sediment. It is probable that the precipitation of pyrite down the sediment column was affected by upward migration and anaerobic oxidation of methane. This process is associated with formation of pyrite showing heavier isotopic compositions of sulfur (Borowski et al., 2010; Jørgensen et al., 2004; Neretin et al., 2004). A systematic shift toward heavier isotopic compositions of pyrite from the rock matrix to phosphate fraction in the phosphatic black shales (in the range of 10e30&) and heavy isotopic compositions of pyrite in the massive mudstone support this interpretation (Figs. 21 and 24b). These compositions suggest isotopic fractionation of sulfur during bacterial sulfate reduction in the range between ca. 30& and 3&. They are indicative of progressive pyrite precipitation during authigenesis and early diagenesis of the sediment in a closing geochemical system downward the sediment column (Brüchert and Pratt, 1999; Calvert et al., 1996; McKay and Longstaffe, 2003). 5.2. Biological productivity The degree of water column productivity in OC-rich facies of the Botneheia Formation is deduced using the microscopic assessment of the preserved organic fractions (Boucsein and Stein, 2009; Hermann et al., 2011; Rimmer et al., 2004), TOC and pyrolytic data (Barker et al., 2001; Rimmer et al., 1993; Stein, 2007), and the content of authigenic phosphorus (Algeo et al., 2011; Nilsen et al., 2003; Sageman, et al., 2003). In this study, only a semiquantitative assessment of the maceral abundances was possible.

This owes to the fact that most sedimentary organic matter in the Botneheia Formation occurs in extremely fine fraction, being severely compressed in shaly sediments. The black shale facies of the Muen Member (Units 1e4) shows a mixture of land-derived and autochthonous marine organic matter, with a varying contribution of redeposited fractions. This composition is typical for distal prodelta environment (e.g., Alt-Epping et al., 2007; Bourgeois et al., 2011; Goineau et al., 2011). Distal prodelta has been suggested as a sedimentary context of the Botneheia Formation in NW Barents Sea shelf on the basis of sequence stratigraphy (Høy and Lundschien, 2011). The sedimentologic and petrographic observations in eastern Svalbard point to pulses of supply of fine clastic material of low to moderate mineral maturity from the southern part of the Barents Sea shelf (Figs. 3d, 5a, 6g,h). Terrestrial organic matter deposited in this environment was relatively fresh and labile, as it is indicated by preferential pyritization of herbamorphinite/vitrinite during early diagenesis (Fig. 14b). Liptinite components embrace a spectrum of mostly fragmented algal particles, suggesting biological and mechanical reworking of muddy sediment, which is consistent with the oxic bottom conditions, bioturbated nature of sediment, and recurrent winnowing. Irregular increase of the content of autochthonous matter at the expense of terrestrial fractions upward the succession marks the increase and then domination of marine productivity in the upper part of the Muen Member. This increase was associated with a decrease of the content of the silt fraction and an increasingupward mineral maturity of sediment that suggest diminishing influence of the deltaic systems in the NW Barents Sea shelf following the Anisian transgression. The onset of phosphogenic conditions at base of Unit 5 marks the increase of productivity (and the related deposition of organic and skeletal phosphorus) to a level that triggered authigenic precipitation of CFA in sediment (Figs. 20e22). The high biological productivity lasted throughout the deposition of the phosphogenic succession (Fig. 3). The microscopic survey demonstrates that the primary productivity was planktonic (Figs. 7 and 8). It was dominated by microalgae, though Tasmanites was a noticeable component, and abundant at some levels in the upper part of the Blanknuten Member. The analysis of phosphatic matrices points to the importance of zooplankton (including abundant radiolaria) in vertical transport of organic matter from the surficial productivity zone to the sea bottom (Fig. 10e, f). The rapid transport through the water column seems to have been crucial for the abundance and chemical parameters of the preserved organic matter (Calvert, 1987; Suess et al., 1987; Ten Haven et al., 1992). The organic matter is hydrogen-rich and oxygen-poor throughout the phosphogenic succession, and its content covaries with the hydrogen richness (Figs. 20e22). Common lithology changes during the late transgressive phase (Units 5e6), including alternation of laminated and bioturbated phosphatic shale packages, and traces of mechanical reworking and winnowing in allochthonous phosphate beds, record an interplay between OC-rich sedimentation, phosphogenesis, changes in bottom oxygenation, and actions of dynamic agents (Figs. 4 and 11). The development of euxinic conditions during the deposition of Unit 7 was associated with maximum flooding of the Barents Sea shelf. This unit shows the highest concentration of alginite and radiolaria molds, suggesting continuation of high biological productivity conditions during the euxinic period (Figs. 7 and 19). Recurrent intervals enriched in shells of juvenile bivalves indicate events of unsuccessful settlements of the planktonic larvae over sulphidic bottoms (Fig. 5e). The abundance of scattered quartz cement related to diagenetic mobilization of biogenic silica, coupled with scarcity of detrital micas and feldspars, confirms very low sedimentation rates and mineral starvation (Fig. 6c,d). Low

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bottom dynamics is reflected in negligible contribution of liptodetrinite compared to the under- and overlying phosphatic shales (Fig. 7). The regressive trend observed in Units 8 and 9 in the upper part of the Blanknuten Member was associated with continuation of high biological productivity in Svalbard, though the succession is incompletely preserved due to intraformational erosion and condensation (Figs. 2 and 3). A condensed phosphorite conglomerate at bottom of Unit 9 (9A) is recognized in most sections in eastern Svalbard. In southern Barentsøya, it overlies erosional surface that cuts into the massive phosphatic mudstone of Unit 7. During the termination of the regressive phase, Tasmanites blooms contributed to local enrichment in organic carbon of the topmost black shales of Unit 9 (9B). 5.3. The organic mattereapatiteepyrite association Enhanced preservation of organic matter in fine-grained facies of the shelf and epicontinental basins usually is linked to oxygendepleted bottom waters and/or anoxic (euxinic) conditions (Arthur and Sageman, 1994; Calvert, 1987; Demaison and Moore, 1980; Wignall, 1994). These conditions might result from pycnocline development in the water column (salinity and/or temperature stratification) or expansion and strengthening of oxygenminimum zone (OMZ) or combinations of the two mechanisms (Algeo and Heckel, 2008 for discussion). The basic difference between these mechanisms is the nature of processes responsible for oxygen deficiency, which is related to the limited water renewal (suppressed oxygen supply) in the stratified basins or to the enhanced oxygen demand resulting from excess decomposition of organic matter in OMZs. As a consequence, the OC-rich facies of the stratified basins concentrate organic matter derived from various sources, including important land-derived fraction (Arthur and Dean, 1998; Calvert, 1990; Miltner et al., 2005; Tribovillard et al., 2008), while the OC-rich facies underlying OMZs concentrate marine organic matter deposited from indigenous productivity zones (Alt-Epping et al., 2007; Arthur et al., 1998; Trainer et al., 2010; Ten Haven et al., 1992). The OC-rich sediments of both the stratified basins and OMZs show the presence of pyrite originated from bacterial sulfate reduction (Anderson and Raiswell, 2004; Canfield, 2006; Glenn and Arthur, 1988; Wilkin and Arthur, 2001), but differ strikingly in the content of authigenic apatite (CFA), reflecting the absence or the presence of productivity-induced phosphogeniesis, respectively (Arthur and Sageman, 2004; Filippelli, 2001; Froelich et al., 1988; Suess et al., 1987). Thus, the regional association of marine organic matter, pyrite, and CFA in the OC-rich facies is considered typical of upwelling-related zones of high biological productivity and oxygen depletion (Barber and Smith, 1981; Baturin, 1982; Burnett and Roe, 1983; Glenn et al., 1994a). The association of organic matter, pyrite, and CFA is prominent throughout the phosphogenic succession of the Botneheia Formation. It suggests the development of a zone of high biological productivity in the NW Barents Sea shelf during the Middle Triassic transgressiveeregressive cycle. There is growing evidence that the Middle Triassic Panthalassic Ocean was already well-ventilated (Hermann et al., 2011; Wignall et al., 2010; Sano et al., 2012), which would have been associated with turnover of deep oceanic water and formation of hydrodynamic regimes along the oceanic margins. The northern margin of the Barents Sea shelf corresponds to a broad oceanic upwelling zone predicted for the northwestern Pangean margin on the basis of conceptual climate model (Parrish and Curtis, 1982). It is postulated here that the upwelling-related input of nutrient-rich seawater to the margin of the Barents Sea shelf brough about the development of OMZ owing to enhanced cycling of organic carbon and oxygen consumption. The zone of

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productivity and oxygen depletion straddled the shelf margin and expanded southward to the NW Barents Sea shelf following the Anisian transgression. This was made possible after the increase of the water column to a level that allowed inflow of nutrient-rich seawater over the submarine swell in northern Svalbard. OMZs impinging on bottom topography on modern upwelling shelves show patchy belts of OC-rich facies aligned parallel to the shelf margin, with the most oxygen depleted (or anoxic) facies located in their central parts (Calvert and Price, 1983; Emeis et al., 1991; Reimers and Suess, 1983). If the bottom under OMZ receives enough biogenic phosphorus to maintain phosphogenic conditions, the external parts of the OC-rich belt contain bioturbated sediments enriched in phosphate nodules, which decrease in size and vanish toward the central part where laminated sediments with phosphate peloids dominate (Baker and Burnett, 1988; Burnett and Roe, 1983; Glenn and Arthur, 1988; Glenn et al., 1994a). A transgressiveeregressive superposition of facies deposited below the phosphogenic OMZ would give a vertical facies succession similar to the one observed in the Blanknuten Member (Figs. 2 and 3). The lower (Units 5e6) and the upper (Units 8e9) phosphatic black shales are interpreted to represent depositional conditions of a shallower part of OMZ, developed during the late transgressive and regressive phases, respectively. The middle part, encompassing the massive phosphatic mudstone (Unit 7), represents deposition in a deeper part of OMZ during high-stand of the sea. 5.4. The nature of euxinia Major sedimentary phosphate deposits tend to form under oxicto-dysoxic conditions, because the alkalinity rise associated with bacterial sulfate reduction increases the solubility of CFA leading to inhibition of its precipitation in permanently anoxic sulphidic environments (Froelich et al., 1988; Jarvis et al., 1994; Schuffert et al., 1998). Consequently, examples of phosphogenesis under the suggested euxinic (or permanently anoxic bottom) conditions are rare, confined to subordinate phosphate occurrences (Algeo et al., 2004; Föllmi, 1996; Trappe, 1998; Tribovillard et al., 2010; see Piper and Perkins, 2004; Piper et al., 2007 for discussion on redox conditions in the Phosphoria Formation). The data presented in this paper document a regional development of phosphogenic conditions in oxic-to-dysoxic (episodically suboxic and/or anoxic) bottoms of Svalbard (Units 5e6 and 8e9), though the phosphogenesis extended into Unit 7 that shows geochemical and petrographic features indicative of euxinic conditions. The massive phosphatic mudstone contains recurrent phosphatic packstone layers that reflect pulses of phosphate emplacement in the euxinic facies (Fig. 10a,b). Some of these layers show traces of incipient winnowing of sediment from between the peloids, which point to intermittent action of delicate dynamic agents on the sea bottom during the phosphogenic episodes. These traces become common in some layers, though reworked grainstone beds evidencing events of high bottom energy are rare (Fig. 11a,b). Field observations and microscopic survey suggest that there occur up to twenty discrete packstone layers per one meter of vertical section of the massive mudstone in western Edgeøya. This gives ca. 400 such layers formed in the area of best development of euxinia in eastern Svalbard. The likely explanation of their distribution is that the euxinic part of OMZ was a subject of delicate fluctuations, recurrently loosing its strength that allowed intermittent oxygenation of bottom water. Fluctuations in bottom oxygenation related to changes in productivity and/or hydrodynamic regime are observed on different time scales in OMZs of modern upwelling zones (Muñoz et al., 2012; Niggemann et al., 2007; Sifeddine et al., 2008; Scholz et al., 2011). Under euxinic (or anoxic bottom) conditions, degradation of organic matter leads to

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an efficient liberation of organic phosphorus (Filippelli, 2001; Ingall and Jahnke, 1994; Ingall et al., 1993, 2005), which increases the level of porewater phosphate and/or is a subject of biological sequestration (Glenn et al., 1994b; Jarvis et al., 1994; Sannigrahi and Ingall, 2005). This phosphate can rapidly precipitate to form the

apatite phase (possibly preceded by an unstable calcium phosphate) as a result of alkalinity drop and the interface-related geochemical effects caused by oxygenation of bottom environment (Froelich et al., 1988; Krajewski et al., 1994; Schuffert et al., 1994, 1998).

Figure 26. Model of deposition of the Botneheia Formation in eastern Svalbard showing expansion of upwelling-related oxygen-minimum zone (OMZ) on the NW Barents Sea shelf during the Middle Triassic transgressiveeregressive cycle. (a)e(d) illustrate the early transgressive, the late transgressive, the high-stand, and the regressive phases of the cycle, respectively.

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The sulfur isotopic composition of CFA-bound sulfate in the massive phosphatic mudstone (Unit 7) is similar to the compositions in the under- and overlying phosphatic black shales (Units 5e6 and 8e9) (Fig. 21). The obtained d34S values are between the likely composition of coeval marine sulfate and the values lighter by a few per mill (Fig. 25). They indicate that authigenic phosphates of the oxic-to-dysoxic and the euxinic phosphogenic facies originated in pore evnironments containing sulfate with marine isotopic composition, though with changing contribution of sulfate originated due to oxidation of bacterially generated hydrogen sulfide. A slightly narrower d34S range for CFA in Unit 7 suggests short-lasting oxygenation periods that affected phosphogenic pulses in the euxinic facies. The greatest DOP value in Unit 7 (0.95) was noted in a phosphatic grainstone bed that records one of a few well-defined high energy events during the deposition of the euxinic facies (Fig. 21). This value is very close to the values of the under- and overlying mudstone (0.94 and 0.93, respectively), suggesting that temporal bottom ventilations of the euxinic part of OMZ were too short to imprint on the degree of sediment pyritization.

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5.5. The extent of euxinia It has been demonstrated in Section 5.1 that euxinic conditions in the Botneheia Formation can only be documented for the massive phosphatic mudstone of Unit 7. This gives twenty five percent at maximum of the phosphogenic black shale facies deposited in eastern Svalbard (Figs. 2 and 3). The euxinic conditions faded northeastward in Barentsøya as a result of bottom shallowing on a swell area (Fig. 2). Limited outcrops of the formation in southwestern Nordaustlandet (Torellneset) suggest continuation of the swell environment. The swell area showed alternating oxic and dysoxic (possibly also suboxic) bottom conditions, high input of marine organic matter under suppressed mineral sedimentation rates, and noticeable formation of glauconite at the time of deposition of the euxinic facies further south in Svalbard (Figs. 12, 18, 21). It may be suggested that this swell diminished bottom dynamics and enhanced euxinia in Svalbard during highstand of the Middle Triassic cycle. However, exploring this hypothesis requests more geological data from the neighboring shelf area.

Figure 27. Succession of the Botneheia Formation in western Edgeøya showing phases of the Middle Triassic transgressiveeregressive cycle, the suggested sea-level curve, bottom redox conditions, and bullet point characteristics of the OC-rich facies. (a)e(d) illustrate the early transgressive, the late transgressive, the high-stand, and the regressive phases of the cycle, respectively.

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The southernmost outcrops of the massive phosphatic mudstone in Edgeøya (Vogelberget in Tjuvfjorden) suggest continuation of the euxinic depositional regime southward in the Svalbard Platform (Fig. 2). Sejsmic profiles in the Svalbard Platform show fairly uniform characteristics of sequences that can be correlated with the Botneheia Formation (Gløstard-Clark et al., 2010; Høy and Lundschien, 2011). However, the direct data obtained from drillings provide only fragmentary information on their facies development. The drilling east of Kong Karls Land reached topmost part of the Botneheia Formation, which shows Tasmanites-rich, phosphatic mudstone and shale with chert concretions (Riis et al., 2008; Vigran et al., 2008). The depositional redox conditions have not been reported, though the chert concretions are interpreted to have originated as a result of diagenesis of siliceous sponges. This suggests non-euxinic conditions for the drilled part of the phosphogenic succession. In the core obtained from a drilling on the Sentralbanken High, the late Ladinian mudto siltstones with phosphate nodules are interpreted to represent open marine deposition (Riis et al., 2008). The bioturbations, bivalve fossils, and traces of bottom currents (ripple marks) observed in the core also suggest non-euxinic conditions. On Hopen, the Middle Triassic succession is below sea-level. On Bjørnøya, a thin, condensed phosphorite conglomerate (Verdande Bed) marks the southern limit of the phosphogenic facies in the NW Barents Sea shelf (Mørk et al., 1990). It was deposited on a complex horst structure (BjørnøyaeSørkapp and BjørnøyaeStappen highs) located between the southeastern and southwestern deltaic systems of the Barents Sea shelf (Faleide et al., 2008; Gabrielsen et al., 1990; Riis et al., 2008). There are also known OC-rich, phosphatic shale intervals in the Steinkobbe Formation (SpathianeAnisian) at Svalis Dome located in the southern Barents Sea shelf (Mørk and Elvebakk, 1999). However, using the available data it is difficult to constrain correlations between the depositional systems of the Svalis Dome and Svalbard. Westward in Svalbard, the massive phosphatic mudstone in the Blanknuten Member can be recognized in eastern (Weitschat and Dagys, 1989) and central Spitsbergen (Krajewski, 2000b,d; Hounslow et al., 2008). A black shale interval with “anoxic” DOP values and heavy isotopic compositions of pyritic sulfur is known from the upper part of the Passhatten Member (Bravaisberget Formation) in Van Keulenfjorden in western Spitsbergen (Karcz, 2010). It is considered a facies equivalent of the massive phosphatic mudstone of the Botneheia Formation, deposited during maximum transgression of the Middle Triassic cycle. Southwestward in Spitsbergen, the OC-rich phosphogenic facies of the Bravaisberget Formation continues to the northern margin of the SørkappeHornsund High, but without intervals suggesting euxinia or extended bottom anoxia (Krajewski, 2000e). The above data suggest that the euxinic (or sulphidic bottom) conditions in the Middle Triassic transgressiveeregressive cycle were confined to a period of high-stand of the sea in mostly deep shelf setting. These conditions extended from Barentsøya and Edgeøya southward into the Svalbard Platform, and westward to central and western Spitsbergen. It seems plausible that coeval euxinic (anoxic) OC-rich facies developed north of the submarine swell in northern Svalbard, reflecting deeper depositional conditions and diminished bottom dynamics. However, the Triassic succession is not preserved in this area. 6. Conclusions The results of this study allow to propose a model of sedimentation of the Botneheia Formation in eastern Svalbard that links the migration and expansion of the oxygen-minimum zone (OMZ) related to high biological productivity in the NW Barents Sea shelf

with phases of the Middle Triassic transgressiveeregressive cycle (Fig. 26). The model postulates upwelling of nutrient-rich water from the Panthalassic Ocean as a major factor affecting the productivity and burial of organic matter in the shelf environments. Oxygen depletion in these environments was a supplementary factor favoring organic matter preservation. It resulted from the enhanced carbon cycling and oxygen consumption, but also from inflow of oxygen-depleted water originated in deeper parts of OMZ. A combination of these two factors during the transgressivee regressive cycle accounts for the observed types of the OC-rich facies and their mineral and organic associations (Fig. 27). Four stages of the evolution of the OC-rich succession can be discerned. They overlap on the early transgressive, the late transgressive, the high-stand, and the regressive phases of the Middle Triassic cycle. These stages are summarized below in terms of processes leading to the preservation of organic matter and burial petroleum potential of the OC-rich facies. 6.1. Early transgressive phase The black shale facies (Units 1e4) deposited during the early transgressive phase of the Middle Triassic cycle. Interplay between the prodelta and open shelf sedimentation in eastern Svalbard resulted in a mixture of terrestrial and autochthonous marine organic fractions in the silty to muddy sediments of low to moderate mineral maturity (Fig. 26a). The deposition occurred under dominating oxic conditions, though the input of fresh organic matter contributed to oxygen consumption and anoxic sulfidic conditions below the surficial sediment layer (Fig. 27a). No phosphogenic conditions developed in this environment. The TOC and HI of Units 1e4 indicate moderate to good petroleum potential in mature sections in Edgeøya and Barentsøya. There is an irregular increase in organic richness, kerogen quality, and proportion of type II/III organic material toward the overlying phosphogenic black shale facies. It embraces at least two episodes of enhanced input of fine-grained clastics and terrestrial organic matter (Units 1 and 3). These episodes were separated by longer periods during which pulses of supply of fine-grained clastics and terrestrial organic matter overlaped on the organic deposition from marine productivity (Units 2 and 4). Elevated HI values in the upper part of the black shale succession already show the dominance of type II kerogen. 6.2. Late transgressive phase Units 5e6 of the phosphogenic black shale facies deposited during the late transgressive phase of the Middle Triassic cycle. The ongoing marine transgression bought about an inflow of nutrientrich and oxygen-depleted water from OMZ that formed at the margin of the Barents Sea shelf (Fig. 26b). This was made possible after drowning of the submarine swell in northern Svalbard to a depth that no longer provided topographic restriction for the upwelled oceanic water. The drowning seems to have been slow enough to gradually increase productivity in the shelf area. In result, a broad zone of high biological productivity developed in Svalbard and further south in the Svalbard Platform. Enhanced deposition of fresh marine organic matter (and organic and skeletal phosphorus) affected regional development of the OC-rich phosphogenic facies. Intense cycling of organic carbon in the water column combined with actions of dynamic agents led to fluctuating bottom redox conditions from oxic to dysoxic, with episodes of suboxia and or/anoxia (Fig. 27b). The phosphatic black shales embrace packages of bioturbated and laminated sediment with pristine phosphates that interfinger with beds and layers of allochthonous phosphates. They show enhanced mineral maturity

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related to suppressed sedimentation rates in open marine environment. The TOC and HI in Units 5e6 indicate good to very good petroleum potential in mature sections in Edgeøya and Barentsøya. Kerogen is dominated by alginite and liptodetrinite (type II). However, substantial internal source rock quality variations are observed. Notwithstanding this fact, the OC-rich sediment dominates the succession, and the phosphatic black shales of Units 5e6 are widespread in eastern Svalbard. 6.3. High-stand phase The massive phosphatic mudstone (Unit 7) of the phosphogenic black shale facies deposited during the high-stand of the Middle Triassic cycle. Rapid increase of the water column at the beginning of this phase enabled further expansion of OMZ and the productivity zone in the shelf area. The environments in eastern Svalbard were characterized by very low terrigenous sediment input, and high organic matter and biogenic (radiolarian) silica inputs. Diminished dynamics and intense carbon cycling over the deep shelf led to the development of euxinic bottom water south of the submarine swell of northern Svalbard. Intermittent ventilations of the euxinic environment affected the development of phosphogenic conditions that resulted in the deposition of recurrent pristine and allochthonous phosphatic layers. The TOC and HI in Unit 7 indicate very good petroleum potential in mature sections in Edgeøya and Barentsøya. Kerogen is dominated by lamalginite, with varying admixture of telalginite (type II/I). The massive phosphatic mudstone is the best petroleum source rock in the Botneheia Formation. However, it shows narrower stratigraphic and lateral extent than the host phosphatic black shales. 6.4. Regressive phase Units 8e9 of the phosphogenic black shale facies deposited during the regressive phase of the Middle Triassic cycle. The regressive trend terminated the euxinic conditions in Svalbard, though phosphogenesis was widespread in the oxic-to-dysoxic (episodically suboxic and/or anoxic) bottom environments. Recurrent non-deposition events aided the formation of bone and coquinoid beds in Unit 8. The erosion at base of Unit 9 resulted in local truncation and removal of parts of the underlying succession. The youngest OC-rich phosphatic shale (9B) overlying the phosphorite conglomerate (9A) witnesses to a continuation of biological productivity in the closing shelf basin. The TOC and HI in Units 8e9 indicate good to very good petroleum potential in mature sections in Edgeøya and Barentsøya. Similarly to Units 5e6, kerogen is dominated by alginite and liptodetrinite (type II). However, these sediments are thin compared to the underlying OC-rich facies deposited during the transgressive and the high-stand phases of the Middle Triassic cycle. Acknowledgments Fieldwork in eastern Svalbard was done during two expeditions of the Polish Academy of Sciences (PAS) and two expeditions of the Norwegian Petroleum Directorate (NPD) between 2005 and 2009 using logistical support of SSB Horyzont II and M/S Kongsøy. Special thanks are due to Atle Mørk for inviting the author to NPD expeditions in 2007 and 2009. The Governor of Svalbard (Sysselmannen på Svalbard) granted permissions to use facilities of the Kapp Lee Station in Edgeøya and the Kinnvika Station in Nordaustlandet in 2005 and 2007, respectively. Field assistance of Trond Brekke, Przemys1aw Karcz, Marcin Klisz, Stanis1aw Mazur, Marta  Slubowska-Woldengen, and Mateusz Zab1ocki is greatly appreci_ ated. Bozena qa˛ cka and Tatiana Weso1owska processed samples for

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chemical analyses, and Ewa Deput preparated thin sections. Krzysztof Nejbert kindly allowed microscopic facilities at the University of Warsaw. The Rock-Eval analysis was done at the Institute of Oil and Gas in Kraków, the ICP-ES analysis at ACME Analytical Laboratories Ltd in Vancouver, the isotopic sulfur analysis at the Stable Isotope Laboratory PAS in Warszawa, and the XRD and AAS analyses at the Institute of Geological Sciences PAS in Warszawa. The author thanks two anonymous reviewers for their constructive criticisms and useful recommendations of ways to improve the manuscript. This is a contribution to the SvalGeo research project at the Institute of Geological Sciences PAS. Appendices A and B. Supplementary material Supplementary material associated with this article can be found, in the online version, at http://dx.doi.org/10.1016/j.marpetgeo.2013. 04.016. References Abdullah, W.H., 1999. Organic facies variations in the Triassic shallow marine and deep marine shales of central Spitsbergen, Svalbard. Marine and Petroleum Geology 16, 467e481. Algeo, T.J., Heckel, P.H., 2008. The Late Pennsylvanian Midcontinent Sea of North America: a review. Palaeogeography, Palaeoclimatology, Palaeoecology 268, 205e221. Algeo, T.J., Schwark, L., Hower, J.C., 2004. High-resolution geochemistry and sequence stratigraphy of the Hushpuckney Shale (Swope Formation, eastern Kansas): implications for climato-environmental dynamics of the Late Pennsylvanian Midcontinent Seaway. Chemical Geology 206, 259e288. Algeo, T.J., Kuwahara, K., Sano, H., Bates, S., Lyons, T., Elswick, E., Hinnov, L., Ellwood, B., Moser, J., Maynard, J.B., 2011. Spatial variation in sediment fluxes, redox conditions, and productivity in the PermianeTriassic Panthalassic Ocean. Palaeogeography, Palaeoclimatology, Palaeoecology 308, 65e83. Allison, P.A., Wignall, P.B., Brett, C.E., 1995. Palaeo-oxygenation: effects and recognition. In: Bosence, D.W.J., Allison, P.A. (Eds.), Marine Palaeoenvironmental Analysis from Fossils. Geological Society, London, Special Publications, vol. 83, pp. 97e112. Alt-Epping, U., Mil-Homens, M., Hebbeln, D., Abrantes, F., Schneider, R.R., 2007. Provenance of organic matter and nutrient conditions on a river- and upwelling influenced shelf: a case study from the Portuguese margin. Marine Geology 243, 169e179. Anderson, T.F., Raiswell, R., 2004. Sources and mechanisms for the enrichment of highly reactive iron in euxinic Black Sea sediments. American Journal of Science 304, 203e233. Anderson, T.F., Kruger, J., Raiswell, R., 1987. CeSeFe relationships and the isotopic composition of pyrite in the New Albany Shale of the Illinois Basin, U.S.A. Geochimica et Cosmochimica Acta 51, 2795e2805. Arthur, M.A., Dean, W.E., 1998. Organic-matter production and preservation and evolution of anoxia in the Holocene Black Sea. Paleoceanography 13, 395e411. Arthur, M.A., Sageman, B.B., 1994. Marine black shales: depositional mechanisms and environments of ancient deposits. Annual Review of Earth and Planetary Sciences 22, 499e551. Arthur, M.A., Sageman, B.B., 2004. Sea-level control on source-rock development: perspectives from the Holocene Black Sea, the mid-Cretaceous Western Interior Basin of North America, and the Late Devonian Appalachian Basin. In: Harris, N.B. (Ed.), The Deposition of Organic-Carbon-Rich Sediments: Models, Mechanisms, and Consequences. SEPM (Society for Sedimentary Geology) Special Publication, vol. 82, pp. 35e59. Arthur, M.A., Dean, W.E., Laarkamp, K., 1998. Organic carbon accumulation and preservation in surface sediments on the Peru margin. Chemical Geology 152, 273e286. Baker, K.B., Burnett, W.C., 1988. Distribution, texture and composition of modern phosphate pellets in Peru shelf muds. In: Burnett, W.C., Froelich, P.N. (Eds.), The Origin of Marine Phosphorite. The Results of the R. V. Robert D. Conrad Cruise 23-06 to the Peru Shelf Marine Geology, vol. 80, pp. 195e213. Barber, C., Smith, R.L., 1981. Coastal upwelling ecosystems. In: Longhurst, A.R. (Ed.), Analysis of Marine Ecosystems. Academic Press, New York, pp. 31e68. Barker, C.E., Pawlewicz, M., Cobabe, E.A., 2001. Deposition of sedimentary organic matter in black shale facies indicated by the geochemistry and petrography of highresolution samples, Blake Nose, western North Atlantic. In: Kroon, D., Norris, R.D., Klaus, A. (Eds.), Western North Atlantic Palaeogene and Cretaceous Palaeoceanography. Geological Society, London, Special Publications, vol.183, pp. 49e72. Baturin, G.N., 1982. Phosphorites on the sea floor. Origin, composition and distribution. In: Developments in Sedimentology, vol. 33. Elsevier, Amsterdam, 343 pp. Behar, F., Beaumont, V., Penteado, H.L. De B., 2001. Rock-Eval 6 technology: performances and developments. Oil & Gas Science and Technology e Revue d’IFP 56, 111e134.

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