Photic zone euxinia in the central Rhaetian Sea prior the Triassic-Jurassic boundary

Palaeogeography, Palaeoclimatology, Palaeoecology 461 (2016) 55–64

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Photic zone euxinia in the central Rhaetian Sea prior the Triassic-Jurassic boundary Martin Blumenberg a,⁎, Carmen Heunisch b, Andreas Lückge a, Georg Scheeder a, Frank Wiese c,d a

Federal Institute for Geosciences and Natural Resources (BGR), Stilleweg 2, 30655 Hannover, Germany State Authority for Mining, Energy and Geology (LBEG), Stilleweg 2, 30655 Hannover, Germany c Department of Earth & Environmental Sciences, Palaeontology and Geobiology, Ludwig-Maximilians-Universität München, Richard-Wagner-Str. 10, 80333 München, Germany d Department of Geobiology, Geoscience Centre, Georg-August-University Göttingen, Goldschmidtstr. 3, 37077 Göttingen, Germany b

a r t i c l e

i n f o

Article history: Received 14 April 2016 Received in revised form 1 August 2016 Accepted 3 August 2016 Available online 4 August 2016 Keywords: Biomarkers Central Atlantic Magmatic Province (CAMP) Isorenieratane Hebelermeer 2 drillcore

a b s t r a c t Shortly after the deposition of black shales in the Rhaetian Sea (Central European Basin, CEB), the Triassic/Jurassic (Tr/J) boundary witnessed one of the Big Five mass extinction events in Earth's history. Aiming at a better understanding of paleoenvironmental changes in the (geochemically) less well-known middle to late Rhaetian (German subdivision), we studied samples from a cored borehole from the central Rhaetian Sea in NW Germany (Hebelermeer 2). Biomarkers, palynomorphs and bulk geochemistry all support a marine/brackish setting with inputs from terrestrial plants. Dinosteranes were found throughout the core, most likely suggesting the spread of dinoflagellates. It is widely accepted that volcanic exhalations during the development of the Central Atlantic Magmatic Province (CAMP) had a major impact on marine and terrestrial environments in the earliest Jurassic e.g., the development of photic zone euxinia and H2S poisoning of benthic life. There is evidence from the studied core, however, that comparable conditions already thrived in the central Rhaetian Sea during the middle Rhaetian. A first indication of euxinia/anoxia is evident from low to very low total organic carbon versus total sulfur (TOC/TS) ratios (~b1) with a minimum preceding the Tr/J boundary (0.03). The latter very low value hints at a decoupling of S and C cycles and eventually abiogenic pyrite formation. Water-column anoxia during the middle Rhaetian is indicated by the occurrence of the biomarker gammacerane, which records ciliates living at the O2-H2S chemocline. The strongest support for a stratified water column even with photic zone euxinia comes from high abundances of isorenieratane. This biomarker is a pigment of anoxygenic phototrophic green sulfur bacteria, which use H2S for photosynthesis. Our data point to perturbations in the biogeochemical cycles of sulfur and carbon already in the middle Rhaetian, which are possibly linked to early volcanic activities and SO2, H2S, and CO2 eruptions. © 2016 Elsevier B.V. All rights reserved.

1. Introduction The Triassic-Jurassic (Tr/J) boundary marks one of the most important mass extinction events in the Phanerozoic (Hesselbo et al., 2002; Hesselbo et al., 2004; Jaraula et al., 2013; Kasprak et al., 2015; Raup and Sepkoski, 1982; Williford et al., 2014). It is linked to the development of extensive and large igneous provinces (i.e. Central Atlantic Magmatic Province; CAMP; Marzoli et al., 1999) in the context of the breakup of the super-continent Pangaea. A similar scenario can be envisaged for the Permian/Triassic boundary extinction (Siberian Trap volcanism; Renne et al., 1995). Mechanisms such as massive CO2, SO2, and thermogenic methane emissions with subsequent perturbations in the carbon and sulfur cycles were crucial for mass extinctions in the marine and terrestrial realms (for the Tr/J boundary see Beerling and Berner

⁎ Corresponding author. E-mail address: [email protected] (M. Blumenberg).

http://dx.doi.org/10.1016/j.palaeo.2016.08.007 0031-0182/© 2016 Elsevier B.V. All rights reserved.

(2002); Hesselbo et al. (2002); Richoz et al. (2012)). Ocean acidification, eutrophication of the seas due to increased continental weathering and phosphate and iron run-off, resulting in anoxia and photic zone euxinia, are believed to have played a major role during both extinction events (see for review van de Schootbrugge and Wignall, 2015). In case of the Tr/J boundary, the major event is recorded as a globally occurring, negative carbon isotope excursion (initial-CIE) in, e.g., sedimentary organic carbon (Hesselbo et al., 2002), alkanes of terrestrial plants (Williford et al., 2014) and fossil leaf (Ruhl and Kürschner, 2011). The exact timing of the onset of magmatic activities, biogeochemical changes, and mass extinction, however, is still a matter of discussion. Based on 187 Os/188Os data, Kuroda et al. (2010) suggested early Rhaetian volcanism, and also Ruhl and Kürschner (2011) proposed from stable carbon isotope data a Rhaetian age for the initial CAMP activities. This data is in accordance with a marked drop of biodiversity already during the Rhaetian (Kiessling et al., 2007; McElwain et al., 2007). Early Rhaetian magmatic activities are proposed to be expressed by a “precursor-CIE” (Ruhl and Kürschner, 2011).

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One of the key arguments for the linkage of H2S poisoning and the Tr/J marine mass extinction – comparable to a possible Permian/Triassic kill mechanism (Wignall and Hallam, 1992; Grice et al., 2005; Kump et al., 2005; Meyer et al., 2008) – comes from the occurrence of isorenieratane in earliest Jurassic sediments of the Rhaetian Sea (Richoz et al., 2012; Whiteside and Grice, 2016). This biomarker is produced by anoxygenic phototrophic green-sulfur bacteria, which require a hydrogen sulfide (H2S) environment in the photic zone. This environmental situation is termed photic zone euxinia. Jaraula et al. (2013) observed isorenieratane in a sedimentary succession from the southern United Kingdom (St. Audrie's Bay), and as isorenieratane was also found in samples deposited slightly before the Tr/J boundary, it was postulated that episodic photic zone euxinia could have occurred temporarily before the onset of the Jurassic. However, the general conclusion of the study was that photic zone euxinia is directly linked to CAMP, the initial CIE, and the Tr/J boundary (Jaraula et al., 2013). Following the interpretation above it has to be considered though that: the studied sections neither record the (i) situation in the central Rhaetian Sea (Fig. 1; e.g. proximal position of core Mariental 1; Richoz et al., 2012) nor (ii) did samples of (middle) Rhaetian age receive much attention (e.g. core Rosswinkel FR 204-201; Richoz et al., 2012). This time interval, however, is crucial in understanding late Triassic changes (Kiessling et al., 2007). To bridge this gap, we studied an interval of a core from the central Rhaetian Sea (Fig. 1; Hebelermeer 2 site). This sequence includes black shales with high contents of total organic carbon (TOC) as well as laminated shales/silt stones, which are generally indicative of a low energy system and the deposition below storm wave base. Our work focused on the interval covering the Middle Rhaetian to the Lower Jurassic. We

used bulk carbon and sulfur geochemistry, palynomorphs, and extractable hydrocarbon biomarkers to reconstruct palaeoenvironmental changes in the Rhaetian Sea. 2. Regional geology and paleogeography The working area is located in an intracontinental depositional setting, which traditionally has been referred to as the Germanic Basin of the Triassic period. In recent literature, this well-established term has progressively been abandoned in favor of Southern Permian Basin (e.g. van Wees et al., 2000) or Central European Basin, CEB (e.g. Fischer et al., 2012), a term applicable for the Permian to Jurassic depositional area within central Europe. In this study, we follow the term CEB, and we use the name Rhaetian Sea for the western part of the CEB, which was presumably permanently flooded during the terminal Triassic. Conditions were most likely brackish in its central parts, and clastic input was shed from the Bohemian and Fennoscandian Highs (Fischer et al., 2012; Fig. 1). In the central parts, laminated shales were periodically deposited, while in the eastern part of the sea deltaic successions occurred. Repeated marine ingressions came from the SW through the Burgundy Alemannic Gateway and from the NW (Fischer et al., 2012). During the late Triassic (middle Keuper), the CEB was affected by synsedimentary extensional tectonism resulting in an uplift of salt structures, which locally were truncated by Lower Jurassic strata (Maystrenko et al., 2013, with further references). The Hebelermeer 2 core presented herein was retrieved from the central part of the former Rhaetian Sea (Fig. 1; Fischer et al., 2012). Bio- and lithostratigraphic subdivisions of the core are based on detailed micro- and macrofossil studies as well as on lithologic characters of the

Fig. 1. Hebelermeer 2 drillcore site near Meppen in northern Germany and the paleogeographic situation during the middle Rhaetian. 1 = position of the St. Audrie's sequence in S UK (Hesselbo et al., 2004); 2 = Mariental 1 in northern Germany (van de Schootbrugge et al., 2009; Heunisch et al., 2010); 3 = Rosswinkel in Luxemburg (Richoz et al., 2012). The paleogeographic map including the paleolatitudes was modified after Fischer et al. (2012) and references therein.

M. Blumenberg et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 461 (2016) 55–64

sediments (Schwarzenhölzer, 1955). Although the original core covered an almost complete succession of Cenozoic and Mesozoic sediments down to the Lower Rhaetian, only samples of Upper Triassic to Lower Jurassic (Hettangian) strata were available for this study (1571.5– 1710 m depth).

3. Material and methods 3.1. Samples Seven samples from core Hebelermeer 2 (52.675842° N, 7.208970° E; maximum core depth 1714 m), drilled in 1955 by the Deutsche Vacuum Öl, GEW Elwerath, Deutsche Schachtbau, Wintershall, KCA Deutag, were sampled in 2014 at the core depository of BGR and LBEG in the Geocenter Hannover. The investigated samples have geological ages from the late Triassic to the early Jurassic. Detailed data on coring and stratigraphy of the whole core are available in the LBEG archive in Hannover. The detailed biostratigraphy and petrography was described in 1955 by Hubert Steghaus and Klaus Hoffmann, respectively (both formerly in the LBEG precursor-institution Amt für Bodenforschung). Biostratigraphy using microfossils was performed on numerous subsamples throughout the entire core. However, only selected samples within scientifically and economically interesting intervals have been stored and were available for further studies. For stratigraphy of the investigated section see the left panel in Figs. 2 and 3. For the discrete samples used in this study, precise depths are not available, but a depth range (noted on the core boxes) is given, which is sufficiently detailed for this study (see depth ranges in Table 1 and respective grey intervals in Figs. 2 and 3). Clearly, with the relatively low sampling resolution we cannot show the development and cease of the crucial CAMP-induced oceanic perturbations in the sense of a dynamic record. However, we are able to provide a succession of paleobiological and geochemical snap-shots from the late Triassic into the earliest Jurassic, each of which is indicative enough to justify the conclusions in this work.

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3.2. Bulk geochemistry (CS-analyses, δ13CTOC, Rock-Eval, XRF) To exclude contamination, surfaces of the samples used for geochemical work were removed with a precision saw. Aliquots of the dried sediment samples (50 °C for 48 h) were crushed and ground (grain size b 100 μm) using a vibratory disc mill. Bulk sulfur and carbon were determined on sediment aliquots by a LECO CS-230 elemental analyzer (LECO Instrumente, Germany). TOC was determined on decalcified (acidification with 10% hydrochloric acid; HCl at 80 °C) and dried samples. About 180 mg of each sample was burned in a highfrequency induction furnace in an oxygen atmosphere by use of the absorption signal at the infrared detector. The instrument was calibrated using a suite of commercially available standards (LECO). Reproducibility of TS- and C-measurements was ±0.02%. Bulk δ13CTOC analysis was made via elemental analysis-isotope ratio mass spectrometry (EA-IRMS) using a coupled Flash EA 1112 (Thermo Fisher Scientific), ConFlow IV (Thermo Scientific), and a Delta V Advantage IRMS (Thermo Fisher Scientific). Rock-Eval pyrolysis was performed using a Rock-Eval-6 analyzer (Vinci Technologies). This technique has been used as a rapid method to determine the hydrocarbon source rock potential of sedimentary rocks. X-ray fluorescence (XRF) analyses were performed using a PANanalytical Axios WDRF System calibrated with certified reference material (among others rocks, soils, ores, and sediments from USGS, NIST). 1 g of pulverized sample was initially heated up to 1030 °C to determine the loss on ignition, then mixed with 5 g lithium borate to prepare a fusion at 1200 °C prior to the XRF analysis. Differential Thermal Analysis (DTA) mineralogical investigations were performed using a Netzsch 449 F3 Jupiter thermobalance equipped with a sample holder linked to a Netzsch QMS 403 C Aeolus mass spectrometer (MS). 100 mg of powdered material previously equilibrated at 53% relative humidity (RH) was heated from 25 to 1150 °C with a heating rate of 10 K/min. For Scanning Electron Microscopy (SEM) investigation, a FEI Quanta 600 F operated in low-vacuum mode (0.6 mbar) was used. Therefore,

Fig. 2. Bulk geochemistry in the Hebelermeer 2 core. Grey boxes mark intervals available for geochemical and palynological studies. Exact depths of the samples may vary within the individual grey areas. Biostratigraphic zones (in italics) were defined by Lund (1977) and were used for biostratigraphy. PT = Pinuspollenites-Trachysporites zone, RP = RicciisporitesPolypodiisporites zone, RL = Rhaetipollis-Limbosporites zone, CE = Corollina-Enzonalasporites zone.

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Fig. 3. Biomarkers in the Hebelermeer 2 core from the central Rhaetian Sea. Grey boxes mark intervals available for geochemical and palynological studies. Exact depths of the samples may vary within the individual grey areas. Concentrations for isorenieratane are semi-quantitative.

sputtering of the samples with gold or carbon is not necessary. The microscope is equipped with the Energy Dispersive X-Ray (EDX)-system Genesis 4000 of EDAX. 3.3. Palynomorphs Samples for palynological investigations are aliquots of those used for geochemical analyses. The preparation followed the standard preparation procedure at BGR/LBEG: treatment of the crushed sediment sample (about 20 g) with HCl 10%, HF 40%, after each step neutralizing the sample with distilled water, sieving the acid insoluble residue ultrasonically and mesh size 10 μm, and mounting the organic residue on slides with glycerin jelly. 3.4. Extraction and biomarker analyses For analysis of mid- to high-molecular weight hydrocarbons, organic compounds were extracted from the sediment samples with dichloromethane (DCM) using an accelerated solvent extraction device (Dionex ASE-350). 10 g of the sample was diluted with 22 g sea sand (Carl Roth,

Germany, extracted and annealed, 4 h at 400 °C) within extraction cells. Organic extracts were collected automatically, and elemental sulfur was subsequently removed from the extract by treatment with activated (10% HCl at 60 °C for approximately 1 h) copper granula. Samples that yielded N10 mg of extract were subjected to a fractionation procedure. Prior to this, asphaltenes in the extracts were precipitated by adding 2 mL DCM and 60 mL petroleum ether to (at maximum) 100 mg of extract (reaction time 12 h). Subsequently, the extracts were centrifuged at 3000 rpm for 15 min. The supernatant solution containing maltenes and resins was collected and the solvent removed through evaporation in a nitrogen atmosphere at 35 °C. Parallel extraction and asphaltene precipitation of samples of known composition (MGS-1 sediment and/or the Norwegian Geochemical Standard NSO-1 oil) for every sample sequence assured reproducibility control of the method. The residual maltenes and resins (up to 100 mg) were separated into aliphatic and aromatic fractions as well as into heterocompounds (NSO-compounds) on silica gel (activated at 240 °C for 12 h) by mid-pressure liquid chromatography (BESTA-Technik für Chromatographie GmbH, using a sequence of organic solvents of different polarity (isohexane, isohexane/DCM (mix 2:1; v:v), DCM/methanol

Table 1 Samples from the Hebelermeer 2 core used for the study. nd = not determined due to a too low S2-peak after Rock-Eval. Sample code

Depth interval in core (m)

Age

TS (%)

TOC (%)

Ccarb (%)

δ13C TOC (‰)

Tmax (°C)

HI (mg HC/g TOC)

Kerogen type

Heb-1 Heb-2 Heb-3 Heb-4 Heb-5 Heb-6 Heb-7

1571.5–1576 1656.8–1660.8 1670–1676.5 1683–1689.6 1689.6–1695.6 1695.6–1701 1704–1710

Jurassic (Lias α 1) Triassic (Upper Rhaetian) Triassic (M. Rhaetian) Triassic (M. Rhaetian) Triassic (M. Rhaetian) Triassic (Lower–M. Rhaetian) Triassic (Lower–M. Rhaetian)

2.09 8.78 4.24 4.01 4.45 1.55 0.23

2.34 0.25 7.98 3.33 6.02 1.23 0.16

1.74 0.60 0.57 1.30 0.01 0.00 9.76

−27.6 −23.8 −26.6 −27.3 −27.1 −27.7 −27.0

435 nd 432 432 427 424 nd

285 nd 551 375 357 69 nd

II/III – II II II II/III? –

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(mix 2:1; v:v)). For samples with extract amounts below 10 mg, only the aliphatic fraction has been retrieved from the chromatographic column. The distribution of compounds contained in the aliphatic and aromatic fractions was determined with an Agilent 6890 gas chromatograph (GC) equipped with a 50 m Ultra 1 column (Agilent; 0.2 mm inner diameter; 0.11 μm film thickness) and connected to a flame ionization detector (FID). Individual biomarkers were analyzed after gas chromatographic separation (Agilent 6890) with a mass spectrometer system (MS; Agilent 7000). Measurements of aliphatic fractions were carried out as multiple-reaction-monitoring (MRM) using parentdaughter-scans. Relative abundances of pristane and phytane were calculated from FID responses, while individual MRM responses were used to determine peak areas for the other biomarkers (hopanes: M+ → m/z 191; steranes (dinosteranes): M+ → m/z 217 (231)). Aromatic fractions were analyzed in full scan mode. Isorenieratane was identified from coelution experiments with an aliphatic fraction from the Bächental Fm. (Rhaetian, Upper Triassic), Allgäu in Bavaria (S. Germany) (and the presence of spectral characteristics typical for aryl isoprenoids; e.g. m/z 133; Grice et al., 1996; Koopmans et al., 1996a). Semiquantification was achieved through comparison of the FID trace of the isorenieratane peak of the sample with the highest signal with an n-C40 alkane calibration sample set with known concentrations. The samples with lower signal abundance for isorenieratane were quantified relative to their abundance on the m/z 133 trace compared to the previously calibrated sample. Relative abundances of short chain aryl isoprenoids (AIR) were calculated according to Schwark and Frimmel (2004). 4. Results 4.1. Bulk geochemistry of the Upper Triassic (Rhaetian) to Lower Jurassic (Hettangian) in the Hebelermeer 2 core Total organic carbon (TOC) and total sulfur (TS) were generally high (Table 1; Fig. 2). Only in sample Heb-7, TOC and TS were low with 0.23 and 0.16%, respectively. While TOC was also low in Heb-2 (0.25%), the accompanying TS was very high (8.78%). Ccarb was, except for the carbonate-rich oldest studied sample Heb-7 (9.76%), relatively low.

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TOC/TS ratios were for all samples lower than described for normal marine situations (2.8; Berner and Raiswell, 1983) and were extremely low in the sample from the uppermost Rhaetian (Heb-2). Reactive Iron (Fe*) was estimated from Fe-data from XRF assuming that aluminosilicate bound Fe is not available for pyrite formation. Fe* was therefore calculated as Fe* = Fe − (0.25 × Al) (Lückge et al., 1999), and its relative distribution is shown in a ternary diagram with TS and TOC in Fig. 4. Except for Heb-7, all samples plot close to the pyrite line and EDX and DTA analyses of sample Heb-2 with the highest sulfur abundance (not shown) support that sulfur is mainly fixed as pyrite. Samples from the Middle Rhaetian (Heb-3 to Heb-5) and the Lower Jurassic (Heb-1), exhibited a good to excellent hydrocarbon generation potential (hydrogen indices (HI) up to 550; Table 1) and Tmax values typical for “early Oil window” maturities (Peters, 1986; Table 1; Supplementary Fig. S1). Rock Eval data HI versus OI (oxygen indices) also revealed that kerogens were either solely or mainly built of type II (marine) kerogen with minor contributions from type III (terrestrial) organic matter (OM; Supplementary Fig. S3). 4.2. Palynomorphs Seven samples were palynologically prepared and investigated. All samples contain identifiable palynomorphs and are biostratigraphically classified after Lund (1977). The content of palynomorphs is variable both in number of specimens and taxa (see Supplementary Table S2). The state of preservation is moderately good (palynomorphs identifiable) to very poor (fragmentation of palynomorphs, extensive destruction of the exines). Apart from palynomorphs, which supply only a minor part of the organic residue after acidic treatment, phytoclasts, amorphous organic matter and pyrite crystals are present to abundant in variable amounts (Fig. 5). The palynomorphs of sample Heb-1 (1571.5–1576.0 m) are poorly preserved. The fossil content is different from the underlying samples and the diversity is low. The presence of Pinuspollenites minimus (Fig. 5A), Chasmatosporites spp., Cerebropollenites sp. together with several phytoplanktonic taxa of marine affiliation (acritarchs indet, dinocysts indet (?Beaumontella cf. caminuspina; Fig. 5C), Campenia gigas, Leiosphaeridia spp., Micrhystridium spp., Tytthodiscus sp., cf. Mancodinium sp.) supports the early Jurassic, probably Hettangian age

Fig. 4. Ternary diagram of TOC, TS and reactive iron (Fe*). Fe* which was calculated from XRF-data.

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Fig. 5. Palynomorphs in the core Hebelermeer 2. A–C: Heb-1 (1571.5–1576 m, Lower Jurassic), D–G: (“dark brown”) palynomorphs from sample Heb-2 (1656.8–1660.8 m, Upper Rhaetian), H–I: Heb-3 (1670–1676.5 m, Middle Rhaetian), J–L: Heb-4 (1683–1689.6 m, Middle Rhaetian), M–O: Heb-5 (1689.6–1695.6 m, Middle Rhaetian), P–S: Heb-6 (1695.6– 1701 m, Lower–Middle Rhaetian), T–V: Heb-7 (1704–1710 m, Lower–Middle Rhaetian). A: Pinuspollenites minimus, B: Trachysporites sp., C:? Beaumontella cf. caminuspina, D: Deltoidospora spp., E: Polypodiisporites polymicroforatus, F: Zebrasporites laevigatus, G: Densosporites sp., H: Riccisporites tuberculatus, I: Rhaetogonyaulax rhaetica, J: Rhaetipollis germanicus, K: Granuloperculatispollis rudis, L: Botryococcus sp., M: Limbosporites lundbladii, N: Spiritisporites spirabilis (reworked?), O: Dapcodinium priscum, P–S: Reduviasporonites chalastus, T: Acanthotriletes varius, U: Calamospora spp., V: “Spinotriletes” sp.

(Pinuspollenites-Trachysporites zone; Lund, 1977); see Supplementary Table S2). Sample Heb-2 (1656.8–1660.8 m) contains abundant clusters of partly irregular shaped pyrite minerals. In addition, some lignitic clasts and few scattered palynomorphs occur, which are often darker than the specimens of the over- and underlying samples and which are only moderately preserved. The dark color and the occurrence of Polypodiisporites polymicroforatus (Fig. 5E) as well as Densosporites spp. (Fig. 5G) support its biostratigraphic position in the Upper Rhaetian (“dark zone”; van de Schootbrugge et al., 2009; Heunisch et al., 2010) (Upper Rhaetian Ricciisporites-Polypodiisporites zone; Lund, 1977). Samples Heb-3 to Heb-5 (1670.0–1695.6 m) contain a spectrum of palynomorphs, which confirm the biostratigraphic microfaunal assignment from 1955 (see Materials & Methods) as Middle Rhaetian (Rhaetipollis-Limbosporites zone; Lund, 1977). Characteristic are the dinocysts Rhaetogonyaulax rhaetica, Dapcodinium priscum and Beaumontella spp. together with Rhaetipollis germanicus, Ricciisporites tuberculatus and Limbosporites lundbladii (Fig. 5H–L). Interestingly, the central sample Heb-4 contains no marine acritarchs or dinocysts (conflicting with high dinosterane abundances; see discussion below). In Heb-5 (1689.6–1695.6 m) few specimens of Reduviasporonites chalastus are present. This taxon belongs to Zygnemataceae and is an extant freshwater representative of these algae accepting also acidified settings (Kleeberg, 2004). The content of spores and pollen in the samples Heb-6 to Heb-7 (1695.6 m to 1710 m) are indicative of an early (to middle) Rhaetian age (e.g. Geopollis zwolinskae, Acanthotriletes varius, Conbaculatisporites mesozoicus, Enzonalasporites spp.) (Corollina-Enzonalasporites zone; Lund, 1977). In Heb-6 more specimens of Reduviasporonites chalastus than in Heb-5 were found. In conjunction with the occurrence of dinocysts, acritarchs and prasinophycean algae (Cymatiosphaera) as

well as abundant amorphous organic matter, a brackish (to marine) environment is likely. Palynomorphs (Botryococcus sp.; Schizosporis sp.) in Heb-7 suggest brackish to freshwater environment at least for (parts of) the photic zone. 4.3. Biomarkers Hydrocarbon biomarkers were found in all samples and revealed maturity indices typical for immaturity to early oil window thermal maturity (e.g. from carbon preference index, CPI, and C29S/(C29S + R)steranes; Supplementary Fig. S3). 4-Desmethyl sterane distributions were not varying in the profile, demonstrating marine/brackish conditions for all samples (with minor terrestrial input; see Supplement). Isorenieratane abundance was highest in the samples from the Middle Rhaetian (Heb-3 to -5) and was moderately abundant in the Lower Jurassic claystone (Fig. 3). Isorenieratane was absent in the sediments of the Lower and Upper Rhaetian (Heb-7 and Heb-2). The AIR was calculated for samples with abundant isorenieratane (Heb-1 and Heb-3 to -6). In samples from the Middle Rhaetian an AIR of 0.7 to 1.1 was calculated whereas for Heb-1 from the Lower Jurassic an AIR of 0.5 was observed. Dinosteranes (4α,23,24-trimethylcholestanes) were found in all samples. Among total steranes, dinosteranes were particularly abundant in the Middle Rhaetian samples (Heb-3 to Heb-5). Similar comparatively high abundances like for isorenieratane and dinosteranes were found for gammacerane (Fig. 3). The ratio of pristane (Pr) to phytane (Ph), which are both commonly attributed to phototrophic organisms (terrestrial plants, algae and bacteria), records with low specifity the setting/redox potential during deposition (Peters et al., 2004). Pr/Ph ratios were generally around or b1 except for the samples Heb-1 (1.9) and Heb-6 (2.3).

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5. Discussion 5.1. Primary productivity and photic zone euxinia during the Middle Rhaetian (and during the Tr/J-boundary interval) in the central Rhaetian Sea High TOC concentrations in the Middle Rhaetian shales (Table 1) indicate that primary productivity and/or preservation was high during this interval. Steranes support a mixed input of predominantly phytoplanktonic (marine/brackish) algae (including dinoflagellates) and, to a minor extent, terrestrial plants for the studied interval (Supplementary Fig. S3). Limited input of terrestrial organic matter is also suggested by the low pristane/phytane ratios (Peters et al., 2004). However, the quality of this index suffers from influences of the varying source organisms, and, therefore, pristane/phytane ratios are relatively unspecific. Here, however, the ratios fall within a range often reported for oils and black shales deposited under oxygen deficiency (Fig. 3). While C27 to C29 4-desmethyl steranes indicate relatively uniform phytoplanktonic communities in the studied interval (Supplementary Fig. S3), dinoflagellate abundance appears to fluctuate. The abundances of dinosteranes (4α,23,24-trimethylcholestanes), a diagenetic product of dinosterol, were low in the oldest studied sample (Lower to Middle Rhaetian), but it increased significantly in the Middle Rhaetian. This indicates that dinosterol-producing dinoflagellates (Boon et al., 1979) were prominent during the Middle Rhaetian. As one of the first records of dinosteranes was from marine sediments of a Rhaetian age (Thomas et al., 1993), the dinoflagellates in the studied core section represent most likely early forms capable of dinosterol-biosynthesis. The marine dinoflagellate R. rhaetica was suggested as good candidate for dinosteranes in Rhaetian sediments (Thomas et al., 1993). Given the presence of R. rhaetica in samples Heb-3 to Heb-5, our data are in line with this interpretation (Supplementary Table S2). It was further hypothesized that their occurrence is linked to a marine (brackish) incursion (Thomas et al., 1993). No dinocysts were found in sample Heb-4 (Middle Rhaetian). The presence of many Botryococcus clusters in this sample points to brackish or even freshwater conditions in the surface waters. However, bulk geochemical data are clearly supportive for marine/brackish conditions and the sample contains also abundant dinosteranes. The conflicting biomarker and palynological data may on the one hand be related to the still unknown origin of dinosterol (cysts or viable organisms). On the other hand, not all dinosterolproducing dinoflagellates form cysts (only 10% of today's dinoflagellates form cysts; Head, 1996). Furthermore, hydrodynamic sorting may also have affected the record of palynomorphs. Phytoplanktonic (dinoflagellate) primary productivity was likely elevated during the Middle Rhaetian, which is indicated by increased TOC, sulfur concentrations and enhanced dinosterane abundance. From the observation of a “precursor CIE” it was proposed that CAMP occurred in different phases and that pulses of volcanogenic CO2 (and other gases) emission took place already during the middle Rhaetian (in the Westbury Fm.; Ruhl and Kürschner, 2011). This would result in fertilizing effects from volcanic ashes and an increased continental phosphate and Fe run-off due to increased silicate weathering (Meyer et al., 2008), which is in good accordance with our data (high TOC, high abundances of algal palynomorphs and steranes in the Middle Rhaetian samples). The ratio of Ts (18α-22,29,30-trisnorhopane) to Tm (17α-22,29,30trisnorhopane) should increase with increasing maturity. However, as the maturities were rather uniform (Tmax values in Table 1), the relative increase in Ts over Tm during the Middle Rhaetian (Fig. 3) requires a different explanation. The Ts/Tm is higher in anoxic samples compared to samples from suboxic depositional environments leading us to interpret the relatively high Ts/Tm ratios as hint to strong anoxia (Peters et al., 2004 and references therein). A low redox scenario during the middle Rhaetian is supported by the distribution of other, more specific biomarkers. The ratio of 17α,21β-C35-homohopanes versus the sum of

61

17α,21β-C31-C35 homohopanes, for instance, is suggested as a measure for redox conditions with high values representing enhanced preservation and, thus, anoxia (Peters et al., 2004 and references therein). In the Hebelermeer 2 core Hopane indices suggest that anoxia was most intense in the Middle Rhaetian (Fig. 3), where the biomarker gammacerane indicates also water column stratification. Gammacerane is believed to derive from tetrahymanol, which is produced by ciliates at oxic-anoxic transition zones (relative abundances are expressed with the Gammacerane Index; Sinninghe Damsté et al., 1995). The strongest evidence for Middle Rhaetian (and to a lesser extent in the Earliest Jurassic) water column stratification, however, comes from the high abundance of isorenieratane. Isorenieratane is a C40 diaryl isoprenoid and a diagenetic product of the pigment isorenieratene (Hartgers et al., 1994) from phototrophic brown green-sulfur bacteria. Green sulfur bacteria thrive at oxic-anoxic transition zones and rely on light and H2S for energy generation. Consequently, isorenieratane is a marker for photic zone euxinia (Sinninghe Damsté et al., 1993). Recently, however, this relationship was challenged due to the unexpected widespread occurrence of isorenieratane and related pigments in numerous oils and source rock extracts of different age (French et al., 2015). The author concluded that in addition to photic zone euxinia other factors e.g., an origin from microbial mats, an allochthonous input or production of these biomarkers in “no analog” paleoenvironments, should also be taken into account to explain isorenieratane findings. Furthermore isorenieratane may also have formed from the ubiquitous βcarotene pigment under specific diagenetic conditions (Koopmans et al., 1996b). 13C-enrichements in the δ13C-values (due to the specific carbon fixation pathway of green-sulfur bacteria; e.g. Grice et al., 1996; Koopmans et al., 1996a) proof that isorenieratane in geological samples is a product of isorenieratene. For the determination of reliable compound specific isotopic values, sufficient abundances and base line peak separation is necessary. This was, however, impossible for the studied samples, due to co-elutions with other compounds in the aromatic fractions. In concert with other biomarkers in the core and low TOC/TS ratios (see below), which also indicate water column stratification, however, we are confident of a green sulfur-bacterial source for isorenieratane in the Hebelermeer 2 core. Consequently, the central Rhaetian Sea is suggested to have been strongly stratified during the middle Rhaetian with a relatively shallow oxic-anoxic transition zone. Relatively low AIR for samples from the Middle Rhaetian (0.7 to 1.1) suggest that the stratification during this interval was permanent (Schwark and Frimmel, 2004). The AIR, however, has to be used with caution because particularly in samples with low organic carbon this ratio can be considerably affected by weathering (Marynowski et al., 2011) and potentially other secondary effects. Photic zone euxinia was used to explain massive hydrogen sulfide poisoning during the earliest Jurassic as consequence of CAMP (Richoz et al., 2012; Jaraula et al., 2013), and our data suggest that photic zone euxinia in the central Rhaetian Sea were already established during middle Rhaetian times. A lack of isorenieratane in according samples from Mariental 1 could be related to its proximity to the coast during the Rhaetian (Fig. 1; Richoz et al., 2012) and the core may, thus, not be suitable to describe the depositional conditions for the entire basin. The reported absence of isorenieratane in the Middle Rhaetian in UK, the Westbury Fm. at St. Audrie's (Jaraula et al., 2013; Fig. 6), may be due to less restricted conditions compared to the semi-enclosed central Rhaetian Sea basin (Fig. 1). This is also supported by the high input of terrestrial Type III organic matter (Jaraula et al., 2013). Organic matter in our samples consists, according to Rock-Eval data (Fig. S2), mostly of marine Type II kerogen, which is in line with a distal position of the Hebelermeer 2 site in the basin. The development of water column stratification during the middle Rhaetian can have different causes. Movements of Permian salt in the German part of the Central European Basin started already in the Middle Keuper (Mohr et al., 2005). As a result, density separation caused by brines from leaking diapirs, could have played a role for the

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5.2. TOC/TS relationship during the Tr/J-boundary interval (and during the Middle Rhaetian) in the central Rhaetian Sea In the sample from the Upper Rhaetian (Heb-2; from the interval closely preceding the Tr/J boundary), the TOC content was extremely low and sulfur very abundant (Table 1; Fig. 2). Most TOC/TS ratios are below the threshold described for “normal marine sediments” of 2.8 (Berner and Raiswell, 1983; Leventhal, 1995). Generally, TOC/TS ratios are under normal marine conditions positively correlated, and the availability of degradable organic matter controls the S fixation via sulfate reduction (SR; Berner, 1984). If the supply of Fe is not limited, HS− is fixed as pyrite. Most of the samples from the Hebelermeer 2 core plot close to the pyrite line, reflecting that Fe was not limited during deposition and organic matter turnover; including the uppermost Rhaetian sample (Heb-2), where sulfur concentrations were extraordinary high (Table 2 and Fig. 4). Fe may have had a detrital, terrigeneous (Brumsack, 1989; Dean and Arthur, 1989) and/or volcanic ash origin. EDX and DTA analyses support that the vast majority of S is fixed as pyrite (not shown). Assuming now that all pyrite results from bacterial SR, one can estimate with the use of Eqs. (1) and (2) the original amount of organic matter (TOCoriginal) before decay through SR

Fig. 6. Relative abundances of isorenieratane and the potential control, photic zone euxinia (PZE), in sequences/core from the Rhaetian Sea or the surrounding as well as proposed relative sea level changes (relative abundances of isorenieratane are grey color coded). Boxes indicate times for which data were available and reported (white = not detected). The Germanic and British stages were related according to Wiedmann et al. (1979). Relative sea level fluctuations are still controversially discussed (Hallam and Wignall, 2004) but were taken and simplified after Hesselbo et al. (2004). Westb. beds = Westbury beds; 1 = Jaraula et al. (2013).; 2 = Richoz et al. (2012), 3 = this study.

establishment of a stratified water column. However, our biomarker and palynological data exclude that enhanced salinities developed at least in the surface waters in which primary producers thrived. We consider enhanced (marine/brackish) primary productivity (e.g. through volcanic CO2 and related Fe and phosphate input from enhanced continental weathering), successive strong microbial turnover and oxygen depression in deep water masses as the crucial drivers for the permanently stratified water column during the Middle Rhaetian.

TOCoriginal ¼ TOCmeasured þ TOClost through SR

ð1Þ

TOCoriginal ¼ TOCmeasured þ TSmeasured  0:75  1:33

ð2Þ

The TOClost through SR can be estimated from TSmeasured and a factor considering diffusional loss of H2S in the recent sediment (Vető et al., 1994; Lückge et al., 1999). This simple relationship is proposed for recent and immature sediments, but was also used for the interpretation of geological settings (Lallier-Vergès et al., 1997). Applying the equations on samples of core Hebelermeer 2 it indicates high losses of TOC through SR for most samples (~30–50%; Table 2). The data further indicate that TOC losses were particularly high for the sample originating close to the Tr/J boundary (Heb-2). For such calculations of TOCoriginal caution is advised, because they are developed for recent sediments and do not account for further diagenetic changes. For instance, organic matter maturation lowers the original TOC/TS ratios so that for instance an original 2.8 ratio changes to 1.9 at a vitrinite reflectance of 0.5% (Raiswell and Berner, 1987). This vitrinite reflectance mirrors that of the maturity of the studied core (see Tmax values in Table 1). Further, the equations do also not account for different magnitudes in the losses of H2S during diagenesis, presumably varying over time. Nonetheless, we still consider TOC/TS ratios to mainly record biogeochemical changes with low values for the Middle Rhaetian being in line with euxinic conditions. In such a setting, organic matter degradation by sulfate reduction started already in the water column, and anoxia was established for a longer time period (which is in line with the conclusions from biomarkers above). The very low TOC/TS ratios for the uppermost Rhaetian suggest extreme depositional conditions. Palynomorphs from this interval exhibit a dark color, which can be a result of exposure to an acidic environment during deposition (van de Schootbrugge et al., 2009). The specific setting is also supported by the calculated (theoretical) 97% TOC loss through SR of the organic input (Table 2). Indeed, SR in the uppermost sediments may still have proceeded on “old” TOC even if productivity may have collapsed due to CAMP (Schoene et al., 2010). Under such a scenario, TOC would have been strongly lowered. The lack of primary production would further have enhanced relative enrichment of sedimentary pyrite due to the absence of dilution effects. Such a scenario is feasible, but we consider a loss of 97% TOC solely by these processes as unlikely if taking into account that a considerable fraction of OM is inert (1 to 90%; Raiswell and Canfield, 2012) and that ~ 25% are degrade(able) in (normal) marine sediments through SR (Dean and Arthur, 1989). This indicates either a decoupling of the normal carbon and sulfur cycles shortly before the Tr/J-boundary (Williford et al.,

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63

Table 2 Total measured organic carbon (TOCmeasured) and total sulfur (TSmeasured) and theoretical calculated TOCoriginal, which was calculated from a theoretical loss through sulfate reduction (SR).

Heb-1 Heb-2 Heb-3 Heb-4 Heb-5 Heb-6 Heb-7 *

Jurassic (Lias α 1) Upper Rhaetian Middle Rhaetian Middle Rhaetian Middle Rhaetian Lower–M. Rhaetian Lower–M. Rhaetian

TSmeasured (%)

TOCmeasured (%)

TOCoriginal (%)

TOC loss through SR (%)

2.09 8.78 4.24 4.01 4.45 1.55 0.23

2.34 0.25 7.98 3.33 6.02 1.23 0.16

4.42 9.01* 12.21 7.33 10.46 2.78 0.39

47 97* 35 55 42 56 59

= calculated values are considered too high (see text).

2009) or that a high portion of the pyrite in this sample is e.g. of late diagenetic, perhaps volcanic origin. Indeed, SEM analyses revealed that pyrite in sample Heb-2 has irregular and not framboidal shapes, which are commonly formed as a consequence of bacterial sulfate reduction (Wilkin and Barnes, 1997). 5.3. Implications for the biogeochemical development of the central Rhaetian Sea during the middle to late Rhaetian We believe that CAMP and sea-level change played key roles for the changes in the central Rhaetian Sea situation observed in the Hebelermeer 2 core. During the middle Rhaetian, permanently stratified, brackish and shallow marine conditions were established in the central Rhaetian Sea. High primary production was potentially triggered by volcanic CO2 and fostered by fertilization effects of external (also volcanic) Fe (Blain et al., 2007) and phosphate delivered through acidic continental weathering (Meyer et al., 2008). The development of water column stratification could have additionally be supported by sulfate-rich brines from leaking salt structures, which developed in the context of diapirism from the middle Keuper on. The situation abruptly ended during the uppermost Rhaetian, which is indicated by very low TOC values, low numbers and low taxa diversity of palynomorphs, and low abundances of biomarkers (e.g., dinosteranes). This falls into an extensive regressive phase (Hesselbo et al., 2004), probably caused by glaciations as result of (Middle to Upper Rhaetian) volcanic ash and SO2 exhalation and the development of a “volcanic winter” (Schoene et al., 2010). These processes shallowed the Rhaetian Sea above storm-wave base (Schoene et al., 2010), terminating water column stratification (also indicated by the lack of isorenieratane). Our data are in line with such scenarios for the Tr/J boundary interval. Shortly after this collapse, a stratified water column re-established in the lowest Jurassic (Richoz et al., 2012; Jaraula et al., 2013) (samples for the lowest Jurassic are unfortunately not available from the Hebelermeer 2 core; Fig. 6). For the following Lower Jurassic, there is compelling evidence for a progressing sea-level rise (Hallam, 1997) and development and re-establishment of marine conditions (Hallam and Wignall, 1999). As described, our data are supportive of a CAMP-related volcanic influence on the central Rhaetian Sea starting already during the middle Rhaetian (Ruhl and Kürschner, 2011) or even earlier (Kuroda et al., 2010). This could imply that H2S poisoning suggested for the earliest Jurassic (Richoz et al., 2012; Jaraula et al., 2013) may have impacted life already earlier during the Upper Rhaetian. 6. Conclusions A drillcore from the central Rhaetian Sea (NW Germany), covering sediments from the Middle Rhaetian to the Lower Jurassic revealed strong evidence for perturbations of the biogeochemical cycles and primary production during the middle Rhaetian. Biomarkers and bulk geochemistry suggest a predominantly marine/brackish setting with minor terrestrial input (with early dinosterol-producing dinoflagellates). Dinoflagellate cysts and acritarchs were common in Middle Rhaetian

samples, but pollen and spores were also prominent in the palynomorph spectrum. Bulk TOC and TS geochemistry and biomarker occurrences (gammacerane from ciliates and isorenieratane from green-sulfur bacteria) clearly point to water column stratification and photic zone euxina, triggering black shale deposition during the Middle Rhaetian. We infer that photic zone euxinia was a result of CAMPrelated volcanism, causing a complex interplay of CO2 and SO2. Our observations indicate that (pre)-CAMP volcanism started already at the beginning of the Middle Rhaetian and appear to have developed gradually (from moderate activities at the beginning to the major eruptive and catastrophic phase during the Tr/J-boundary). A catastrophic input of sulfur into the depositional system and/or a collapsing primary productivity shortly before the Tr/J boundary is indicated by extraordinary high sulfur concentrations (8.8%). In such a scenario, volcanic CO2 eruptions and perhaps Fe and phosphate input from continental weathering could add to the observed gradual increase of sedimentary TOC-contents (up to 8%) during the Middle Rhaetian due to an enhanced primary productivity before potentially rapid SO2-rich eruptions turned sea-waters into a sulfidic brew. Our data suggest that biogeochemical cycles particularly during the uppermost Rhaetian were imbalanced before the system turned back to less harsh but still euxinic conditions in the Lower Jurassic. Acknowledgments Christian Ostertag-Henning, Jochen Erbacher, Gerd Röhling, Jürgen Messner, and Stephan Kaufhold are thanked for helpful discussion and Monika Weiß, Sabrina Koopmann, Petra Adam, Sylvia Kramer, and Barbara Piesker for laboratory assistance. We thank the GDF Suez (now Engie SA) and ExxonMobil for the permission to publish the results. We also thank Ken Williford and the editor Thomas Algeo for their helpful comments as well as three anonymous reviewers for their comments on a previous version of the manuscript. The work was conducted in the frame of the BGR NiKo project (Shale oil and shale gas potential in Germany). Appendix A. Supplementary data Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.palaeo.2016.08.007. References Beerling, D.J., Berner, R.A., 2002. Biogeochemical constraints on the Triassic-Jurassic boundary carbon cycle event. Glob. Biogeochem. Cycles 16. http://dx.doi.org/10. 1029/2001GB001637. Berner, R.A., 1984. Sedimentary pyrite formation: an update. Geochim. Cosmochim. Acta 48, 605–615. Berner, R.A., Raiswell, R., 1983. Burial of organic carbon and pyrite sulfur in sediments over phanerozoic time: a new theory. Geochim. Cosmochim. Acta 47, 855–862. Blain, S., et al., 2007. Effect of natural iron fertilization on carbon sequestration in the Southern Ocean. Nature 446, 1070–1074. Boon, J.J., Rijpstra, I.C., DeLange, F., De Leeuw, J.W., 1979. Black Sea sterol - a molecular fossil for dinoflagellate blooms. Nature 277, 125–217. Brumsack, H.-J., 1989. Geochemistry of recent TOC-rich sediments from the Gulf of California and the Black Sea. Geol. Rund. 78, 851–882.

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