Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 Shallow-water facies setting around the Kačák Event: a multidisciplinary approach P. KÖNIGSHOF1*, A. C. DA SILVA2, T. J. SUTTNER3, E. KIDO3, J. WATERS4, S. K. CARMICHAEL4, U. JANSEN1, D. PAS2 & S. SPASSOV5 1 Senckenberg Research Institute and Natural History Museum, Senckenberganlage 25, 60325 Frankfurt, Germany 2 Sedimentary Petrology, B 20, University of Liege, Sart-Tilman, B-4000 Liège, Belgium 3 University of Graz, Institute of Earth Sciences, Heinrichstraße 26, A-8010 Graz, Austria 4 Department of Geology, Appalachian State University, Boone, NC 28608, USA 5 Geophysical Centre, Royal Meteorological Institute of Belgium, 1 rue du Centre Physique, B 5670 Dourbes (Viroinval), Belgium *Corresponding author (e-mail: peter.koenigshof@senckenberg.de) Abstract: In the Eifel area (western Rheinisches Schiefergebirge), a shallow- to deep-subtidal sequence of mixed carbonates and siltstones around the Kačák Event Interval close to the Eifelian– Givetian stage boundary was studied. An overall transgressive trend is inferred by the microfacies evolution. The stratigraphic variations of magnetic susceptibility in carbonates and in shale intervals show an overall decreasing evolution towards the top, which fits well with the transgressive trend. In addition, carbon and oxygen isotopes, and major, trace and rare earth element (REE) analysis have been used to get a better understanding of palaeoenvironmental variations in a shallow-water realm in the late Eifelian (kockelianus and ensensis conodont biozones): for example, the d13C excursion and Ce anomaly are interpreted to be the local representation of the beginning of the Kačák Event Interval, which is also consistent with the stratigraphy and microfacies analyses. Black shales at the Eifelian–Givetian boundary sections are very common in Europe and elsewhere (summarized in House 1996), representing a global (eustatic) sea-level rise. The Kačák Event (Budil 1995; House 2002) occurs just below the Eifelian– Givetian Stage boundary at the Jebel Mech Irdane section in Morocco, the Global Stratotype Section and Point (GSSP). The Kačák Event is a typical black shale event, connected with the global rise in sea level at the base of the transgressive– regressive (T– R) Cycle Ie of Brett et al. (2011). The sea-level rise is connected with increasing d13C values of about 2‰ at the Eifelian –Givetian boundary (e.g. Buggisch & Mann 2004; van Geldern et al. 2006), and the positive excursion can be attributed to an increased burial of isotopically light organic carbon or increased riverine weathering flux (Kump et al. 1999; Saltzman 2002). This event is characterized by a biotic turnover that is primarily recorded in pelagic faunas, such as conodonts, cephalopods and dacryoconarids (e.g. Chlupáč & Kukal 1986; Bultynck 1987, 1989; Becker & House 2000). The dacryoconarid Nowakia otomari is considered to be an indicator of the Kačák Event or the otomari Event sensu Walliser (1985). In Germany, the otomari Event was described in pelagic carbonates of, for instance, the Odershausen Formation in the eastern Rheinisches Schiefergebirge (eastern RSG) in the Kellerwald area (e.g. Walliser 1985; Schöne 1997) (Fig. 1). The Eifelian –Givetian boundary interval in the Eifel area (western RSG) is characterized by shallow-water successions. The otomari Event has been placed at the boundary between the Nims and Giesdorf members by Struve et al. (1997), whereas Schöne et al. (1998) suggested placing this event at the boundary between the Junkerberg and Freilingen formations (see Table 1). A detailed stratigraphic correlation of sections in the eastern RSG with those of the Eifel area is difficult owing to the different depositional settings (deep-water facies setting v. shallowwater facies setting) and correlation problems: for example, the lack of the dacryoconarid Nowakia otomari in the shallow-water setting. The effects of anoxic events in Earth’s history on shallow-water realms are still poorly documented. In this paper, we describe a shallow-water sequence around the Kačák Event Interval, with special focus on a comparison of magnetic susceptibility data, facies analysis and geochemical data in order to From: Becker, R. T., Königshof, P. & Brett, C. E. (eds) Devonian Climate, Sea Level and Evolutionary Events. Geological Society, London, Special Publications, 423, http://doi.org/10.1144/SP423.4 # 2015 The Geological Society of London. For permissions: http://www.geolsoc.org.uk/permissions. Publishing disclaimer: www.geolsoc.org.uk/pub_ethics Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 1. Geological map of the Rheinisches Schiefergebirge (RSG) and Ardennes (slightly modified from Wehrmann et al. 2005). The rectangle demarcates the study area (Fig. 2). get a better understanding of mainly climate-driven changes and sea-level fluctuations. Geological setting The RSG belongs to the European Variscides, which have been interpreted by a number of studies as a collage of microplates (e.g. Matte 1986; Franke et al. 1990; Franke & Oncken 1990, 1995; Franke 2000) successively accreted during the Early Devonian–Late Carboniferous and leading to the amalgamation of the supercontinent Pangaea (e.g. Scotese & Barret 1990; Kroner & Hahn 2003; Romer et al. 2003; Kroner et al. 2007; Linnemann et al. 2010; Eckelmann et al. 2014). It is widely accepted that the bulk of the RSG is part of the Table 1. Stratigraphic position of the outcrop west of Blankenheim within the Junkerberg Formation, and the transition to the Freilingen Formation Biostratigraphy after Struve (1996), Ochs & Wolfart (1961) and Struve (1982). Geological events: see Schöne et al. (1998), Struve (1982) and Struve et al. (1997). Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 2. Facies model of the Middle Devonian of the Eifel region (modified from Winter 1977). Facies type (a) facies dominated by clastic input; facies type (b) carbonate platforms and biostromal reefs (including the mid-Eifelian High); facies type (c) facies characterized by reduced clastic input and increasing carbonate sedimentation. The black star marks the investigated section near Blankenheimerdorf (50826′ 29.22′′ N, 6838′ 12.79′′ E). Avalonia terrane, which was separated from Gondwana in the Early Ordovician and drifted northwards (e.g. Oncken et al. 2000; Tait et al. 2000; Torsvik & Cocks 2004; Linnemann et al. 2008, 2010; Romer & Hahne 2010). At about 450 Ma, Avalonia collided with Baltica in the Late Ordovician –early Silurian, which led to the closure of the Tornquist Sea. Owing to the collision of Baltica and Avalonia with Laurentia, the Iapetus Ocean was closed at around 420 Ma and Laurussia was formed (Kroner et al. 2007; Linnemann et al. 2008; Nance et al. 2010). The closure of the Rheic Ocean began in the late Silurian –Early Devonian and continued until the Early Carboniferous by successive closing from west to east. The ‘Amorican Terrane Assemblage’ (ATA) is believed to have separated from Gondwana in the Ordovician and followed Avalonia in a northwards direction. An island arc may have formed between Avalonia and the ATA during closure of the Rheic Ocean and accreted to Avalonia in the Early Devonian (Franke & Oncken 1995; Franke 2000). This arc is documented by magmatic rocks in the ‘Northern Phyllite Zone’ and the ‘Mid-German-Crystalline Zone’ (MGCZ) (e.g. Brinkmann 1948; Dombrowski et al. 1995; Reischmann et al. 2001). During the Early Devonian, the Rhenohercynian Basin developed as a narrow (about 250–300 km) but rather elongate (more than 2000 km) sedimentary trough south of the Old Red Continent (Stets & Schäfer 2009). Towards the south, the trough was confined by the MGCZ during the Lochkovian and Pragian. Detrital supply came from northern (Wierich 1999), as well as from southern, source areas (Hahn 1990; Hahn & Zankl 1991). Lower Devonian sequences are characterized mainly by sandstones and siltstones, whereas carbonates are less frequent. Sedimentation generally changed during the Middle Devonian. Biostromes began to flourish in the northern RSG (Pas et al. 2013 and references therein). In the SE RSG, biostromes are associated with volcaniclastic deposits indicating increased volcanic activity (e.g. Königshof et al. 1991, 2010; Nesbor et al. 1993; Braun et al. 1994; Nesbor 2004). The RSG east of the river Rhine is characterized by several autochthonous and parautochthonous (e.g. Wachendorf 1986; Meischner 1991; Schwan 1991; Bender & Königshof 1994), as well as allochthonous, units (e.g. Engel et al. 1983; Oczlon 1992; Franke 2000; Huckriede et al. 2004; Salamon & Königshof 2010; Eckelmann et al. 2014). West of the river Rhine, the Eifel synclines, interpreted as part of the north –south-trending axial depression of the RSG, are the dominant structure (Fig. 2). In contrast to the entire RSG east of the river Rhine, the Eifel area shows very low to lowgrade thermal alteration (e.g. Teichmüller & Teichmüller 1979; Helsen & Königshof 1994; Königshof & Werner 1994). A remagnetization event has been described in the carbonate rocks of the Devonian of the Rhenohercynian Fold-and-Thrust Belt in Belgium and NW Germany, near Cologne (Zwing et al. 2002, 2005, 2009; Zegers et al. 2003; Da Silva et al. 2012, 2013), but without specific studies on the Eifel area. In the Eifel area, siliciclastics were delivered from the north during the Early Devonian and early Middle Devonian (Eifelian), but diminished during Givetian times when shallow subtropical carbonates were established over much of the region. Struve (1963) established a depositional model of the Eifel area with a north– south trending basin surrounded by landmasses, which he considered as the so-called ‘Eifel Sea Strait’. In contrast to this model, Winter (1977) defined three facies belts Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. (Fig. 2) in the Eifel synclines. Later, Faber (1980) modified this model, based on detailed microfacies studies, which gave evidence of small-scale cyclicity developments of a carbonate platform during the early Eifelian, and of a flat shelf lagoon during the late Eifelian and early Givetian, affecting the eastern part of the Eifel synclines. Paproth & Struve (1982) distinguished between north-, west-, and south-Eifel biofacies based on faunal differences. During the Givetian this facies differentiation broke down to some degree and mainly stromatoporoid/coral biostromes extended over the entire area. The studied section lies within the Blankenheim Syncline (Figs 2 & 3), between the villages of Blankenheim and Blankenheimerdorf, and comprises shallow-shelf mixed carbonate and siliciclastic facies of Middle Devonian age (Eifelian) accumulated on the southern margin of the former Avalonia microcontinent. Methods Microfossil separation and microfacies analysis Microfacies analysis is based on facies changes observed in the field and observations from more than 80 thin sections of different formats (most of them in the format of 7.5 × 11 cm). Thin sections are stored at the Senckenberg Research Institute and Natural History Museum, Frankfurt. Fourteen rock samples, between 1 and 3 kg, were treated with formic acid diluted in water at a ratio of 1:5. Conodonts were extracted from residues by hand picking after heavy liquid separation (sodium polytungstate: density 2.79 g cm23) and are stored at the University of Graz. Magnetic susceptibility and specific magnetic measurements Magnetic susceptibility (x) is considered to be a proxy parameter for detrital input, and is used for correlations of different coeval stratigraphic sections (e.g. Ellwood et al. 1999) and palaeoenvironmental or palaeoclimatic reconstructions (Hladil et al. 2009; Da Silva & Boulvain 2010; Whalen & Day 2010; Koptı́ková 2011; De Vleeschouwer et al. 2012; Da Silva et al. 2013). Initial magnetic susceptibility measurements on 77 samples, cleaned of surface alteration, were performed on a KLY-3S Kappabridge (AGICO, noise level 2 × 1028 SI) at the University of Liège (Belgium). Magnetic susceptibility (MS) is expressed in m3 kg21; each data point is the average of three measurements. The sample mass was weighed with a precision of 0.01 g. Fig. 3. Lithological column of the Blankenheim section. The section has a thickness of about 8 m, and is composed of bioclastic wackestones, floatstones and rare grainstones with intercalations of shales and siltstones. Diagenetic processes can considerably influence the detrital susceptibility signal (primary signal); an aspect that is often underestimated (Schneider et al. 2004; De Vleeschouwer et al. 2010; Riquier et al. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT 2010; Da Silva et al. 2012, 2013). Because magnetic minerals ((titano)magnetite, pyrrhotite and greigite) are very sensitive to post-depositional diagenesis, it is essential to understand the nature and origin of the magnetic minerals carrying the magnetic susceptibility signal. In this respect, the nature and grain size (detrital grains are classically coarser than diagenetic grains) of the magnetic minerals should be scrutinized. Natural materials have three major magnetic behaviours: diamagnetic minerals have extremely weak negative values of MS (e.g. calcite and quartz); paramagnetic minerals have weak positive values (e.g. smectite, illite, biotite, dolomite and pyrite); and ferromagnetic (sensu latu) minerals have high and positive values (e.g. magnetite, pyrrhotite, hematite and goethite). In this paper, the nature and grain size of magnetic minerals is constrained through hysteresis loop measurements, acquisition of isothermal remanent magnetization (IRM), backfield curves and short-term remanence decay, which were measured with a J-coercivity rotational magnetometer at the Geophysical Centre of the Royal Meteorological Institute of Dourbes in Belgium. We selected 20 samples on a regular interval, which were cut into a cuboidal shape (approximately 0.9 × 0.7 × 2.5 cm) and weighed with a precision of 0.001 g. Remanent and induced magnetizations were measured between +500 and 2500 mT, with averaged field increments of 0.5 mT per magnetization step. The magnetization duration at each magnetization step is in the order of tenths of a second. When the backfield curve is finished, the direct current is switched off automatically and the magnetizing field decreases towards a constant residual field of about 0.4 mT within the following 0.4 s. The decay of the remaining remanence (IRM 500 mT, 0.4 s) is monitored over 100 s, and is called the short-term remanence loss. The following parameters were derived from the hysteresis loops: saturation magnetization, Ms (A m2 kg21); remanent saturation magnetization, Mrs (A m2 kg21); high-field magnetic susceptibility, xHF (m3 kg21); coercive force, Bc (mT); and remanent coercive force Bcr (mT). Ms, Mrs and xHF were normalized with respect to sample mass. xHF corresponds to the high-field magnetic susceptibility values and represents the paramagnetic plus diamagnetic contributions. The ferromagnetic contribution corresponds to xFerro, which is the total susceptibility (x) minus the xHF. The hysteresis parameters are compared with the classic ‘Day plot’ of Mrs/Ms v. Bcr/Bc (Day et al. 1977; Dunlop 2002), allowing an assessment of the grain size of the magnetic minerals. The amount of high- and low-coercivity minerals can be roughly estimated through the high-field remanence (%), corresponding to the difference in remanence acquired at 300 and 500 mT. Geochemical proxies Carbon (d13Ccarb) and oxygen (d18O) isotope ratios of whole-rock carbonate total organic carbon (TOC) and sulphur content were analysed for 45 rock samples collected through the section. Rock powders were produced from polished rock surfaces of hand specimens by drilling the micritic matrix under a binocular microscope. In addition, carbon isotopes were analysed on isolated prasinophytes (green algae) in 14 samples in order to compare d13Ccarb and d13Corg trends. All ratios of d13Ccarb and d18O, as well as d13Corg, are expressed in the conventional d-notation as per mil (‰) relative to the Vienna Pee Dee Belemnite (Vienna-PDB) standard. In this study, oxygen isotope values of conodont apatite were also measured for two samples (BL1229c and BL12-22) to reconstruct the palaeo-seasurface temperature. They were analysed by using a monogeneric icriodontid conodont assemblage (exclusively well-preserved platform elements). The colour alteration index (CAI) of conodont elements is very low, ranging between 1.5 and 2. The analysis of d18Oapatite was calibrated using a value of 21.7‰ for standard NBS120c. Carbon and oxygen isotope analysis on wholerock carbonate were performed using a Kiel II preparation line and Finnigan MAT Delta Plus mass spectrometer at the University of Graz (Austria). Carbon isotope analysis of organic carbon and oxygen isotope analysis on conodont apatite were performed at the GeoZentrum Nordbayern of the University of Erlangen-Nürnberg (Germany) using an elemental analyser (Carlo-Erba1110) coupled online to a ThermoFinnigan Delta Plus mass spectrometer, and a high-temperature reduction furnace (TC-EA) connected to a ThermoFinnigan Delta Plus mass spectrometer, respectively. TOC and sulphur content were analysed on a LECO CS-300 (version 1.0, Year 1992) at the University of Graz, and are reported as Corg and Stot. One gram of fine powder of each sample was treated three times with 2N HCl for 24 h to remove the carbonate component. Following acid treatment, the samples were rinsed to neutrality (three times with distilled water). The resulting dried powder was then analysed. Whole-rock geochemical analyses (major, trace and rare earth element (REE)) were performed on 45 samples at Activation Laboratories in Ancaster, Ontario, Canada. Proxies for detrital input include total Al2O3, Rb, K2O, TiO2 and SiO2. The carbonate content was measured by CaO + Loss of Ignition (LOI) and Sr. Proxies for anoxia include the chalcophile elements (S, Cu, Zn and Pb), authigenic U (where Uauthigenic ¼ Utotal 2 Th/3) (Wignall & Myers 1988), Ce anomaly (Wright et al. 1987), Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. V/Cr (Pujol et al. 2006), d13C excursions and TOC (T. C. Bond et al. 2013). Primary productivity proxies include excess Ba (where Baexcess ¼ Ba 2 0.15 Al2O3) and excess P2O5 (where Pexcess ¼ P2O5 2 0.15 Al2O3) (Pujol et al. 2006). Bulk d18O values are not correlated with increases in MnO concentrations, nor decreases in total Sr (ppm), nor increases in Ca/Mg ratio. When using these geochemical data in combination with microfacies observations, we therefore conclude that diagenetic alteration is not significant in any of the samples. Results Stratigraphy and microfacies Microfossil biostratigraphy. The Blankenheim section exposes the Junkerberg and Freilingen formations and reaches the latest Eifelian (Table 1). The conodont assemblage is dominated by ‘shallow-marine forms’ of the genus Icriodus (Weddige 1977; Weddige & Ziegler 1979). Conodonts average approximately 10 –30 elements per kg, with a higher number of specimens occurring between samples BL 12-29c and BL 12-22. Elements include icriodontids (commonly platform and rare coniform elements), one broken specimen of Belodella sp., two broken specimens of Polygnathus linguiformis linguiformis Hinde, 1879 (two different morphotypes), and two indeterminable platform fragments of Polygnathus sp. Icriodontid species (Icriodus arkonensis Stauffer, 1838, I. regularicrescens Bultynck, 1970 and I. struvei Weddige, 1977) are the most diagnostic, and constrain the section as ranging from the Tortodus kockelianus to the Polygnathus ensensis Biozone. Icriodus regularicrescens occurs from near the base to the top of the section. The lower boundary of the ensensis Biozone was tentatively set at the first occurrence of Icriodus arkonensis near the top of the section (sample BL 12-2) because the marker species I. obliquimarginatus Bischoff & Ziegler, 1957 was not found. Illustrations of representative specimens of diagnostic conodont species are provided in Figure 4. Apart from conodonts, pyritized foraminiferid tests belonging to the genus Pseudopalmula, Semitextularia, probably Paratikhinella, and ostracod valves of Polyzygia (sample BL 12-34a) and other forms are obtained from several samples, but they are less diagnostic in terms of stratigraphy. Macrofossil biostratigraphy. The succession from BL 12-25 to BL 12-32 contains a brachiopod fauna of biostratigraphic significance (Fig. 5): † Sample BL 12-S 25: Schizophoria schnuri blankenheimensis Struve, 1965; † Sample BL 12-28D: numerous specimens of Iridistrophia? cf. undifera (Schnur, 1853) and Vandercammenina? latistriata (Frech, 1911); † Sample BL 12-29D: numerous Iridistrophia? cf. undifera; † Sample BL 12-32: particularly rich in brachiopods: Vandercammenina? latistriata, Iridistrophia? cf. undifera and few poorly preserved Athyris (Alvarezites) wolfarti Struve, 1992? We collected the following taxa from loose blocks in the sample BL 12: Iridistrophia? cf. undifera, Schizophoria schnuri blankenheimensis, Carpinaria vel Subcuspidella sp. (one specimen), Helaspis sp. (one specimen, probably H. plexa (Wolfart, 1956)) and Strophomenida gen. et sp. indet. (one specimen). The associations suggest an assignment to the late middle Eifelian, equivalent to the Junkerberg Formation and most possibly to the Blankenheim Member (‘latistriatus-Horizont’ sensu Wolfart in Ochs & Wolfart 1961). The correlation of the latter with the ‘Type Eifelian’ succession is, however, controversial. Based on brachiopods, Struve (in Struve et al. 2008) assigns a ‘Nims age’ to the Blankenheim Member, whereas Basse & Müller (2004), arguing with the trilobite stratigraphy, regard a partial Giesdorf age as possible. The study of additional material of the spinocyrtiids from the section may provide a better-constrained age. Microfacies analysis. The studied section is about 8 m thick and is characterized by bioclastic wackestones, floatstones and rare grainstones, with a local occurrence of laminated limestones. The limestone beds vary in thickness from centimetre- to decimetre-scale and are intercalated with shales. The section is composed of shallow-shelf mixed carbonate and siliciclastic facies. Open-marine planktonic organisms, such as dacryoconarid tentaculitids, cephalopods and/or pelagic ostracodes, are absent. Bioclasts are dominated by brachiopods, crinoids, bryozoans and subordinate numbers of trilobites, corals, algae and/or molluscs, depending on environmental changes. All bioclasts are very wellpreserved due to the low diagenetic overprint (Helsen & Königshof 1994). Carbonates lack indicators of pressure solution or strong diagenetic overprint. Microfacies analysis allowed the discrimination of five microfacies, which are described from shallowest to deepest. Microfacies 1: Bioclastic wackestone, partly strongly burrowed. In the outcrop, the grey carbonates of this facies show wavy-nodular to planar bedding. Wackestones have 20– 25% bioclasts, including crinoids, brachiopod shells, bryozoans, trilobites, rare corals (rare Thamnopora), stromatoporoids, ostracodes (Fig. 6a –c: BL 12-23(2), BL Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 4. (a) Icriodus struvei Weddige, 1977, I element, upper view (sample BL 12-34A). (b) Icriodus werneri Weddige, 1977, I element, upper view (sample BL 12-22). (c) Icriodus regularicrescens Bultynck, 1970, I element, upper view (sample BL 12-22). (d) Belodella sp., coniform element, upper view (sample BL 12-5). (e) Icriodus retrodepressus Bultynck, 1970, I element, upper view (sample BL 12-3). (f) Icriodus arkonensis Stauffer, 1938, I element, upper view (sample BL 12-2). (g) & (h) Polygnathus linguiformis linguiformis Hinde, 1879, Pa element, upper and lateral views (sample BL 12-29C). (i) Polygnathus linguiformis linguiformis Stauffer, 1879, posterior part of broken Pa element, upper view (sample BL 12-12). 12-24A and BL 12-25) and occasional gastropods. The matrix is primarily composed of a pelmicrite with a local occurrence of pelsparite. Burrowing is very common, so the texture is more or less homogenized. Minor phosphorite intraclasts show variable grain sizes (fine silt to small pebbles) and, in addition, pyrite and iron oxide crusts occur, mainly as coverings on bioclasts, such as trilobite fragments or brachiopod shells (Fig. 6d, e: BL 12-33(2) and BL 12-34A). Interpretation: The facies points to an upperramp position with moderate water depth just below the fair-weather wave base. The diverse fauna includes primarily brachiopods, trilobites and bryozoans, and suggests an open-marine environment. Gastropods could be an indication of a more restricted environment in the lower part of the section but they are less common. The sediments have been reworked and were deposited close to the place of origin; the sediments are poorly sorted and bioclasts show no abrasion. The more clayey/ marly sediments indicate a former soft substrate, which is strongly burrowed. More calcareous layers occur in the lower part of the sequence. These layers can be correlated with a sampled section (Ernst et al. 2011), which is located very close (less than 20 m) to the described section herein. These authors described a number of bryozoans with increasing calcifying taxa, such as trepostome bryozoans, encrusting worms and calcimicrobes, Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 5. Brachiopods from the section BL 12 (whitened with MgO before photographing, shown in natural size). (a) & (b) Iridistrophia? cf. undifera (Schnur, 1853), ventral (a) and dorsal (b) view of partly exfoliated articulated shell. (c) Schizophoria schnuri blankenheimensis Struve, 1965, dorsal view of articulated shell (sample BL 12-S-25). (d) Carpinaria vel Subcuspidella sp., ventral valve. (e) Helaspis sp., probably H. plexa (Wolfart, 1956), ventral valve. (f) Vandercammenina? latistriata (Frech, 1911), ventral valve. particularly Girvanella from that horizon (Ernst et al. 2011, fig. 10c, d). Girvanella is an indicator of low sedimentation rates and may occur in a water depth of tens of metres (Flügel 2004), although Riding (1975) suggested that no confident depth limits can be set for Girvanella occurrences. A period with preferred calcite precipitation in a shallow-water setting within a photic zone is likely. The existence of phosphorite intraclasts and iron-oxide crusts around some bioclasts, such as trilobites, brachiopod shells (Fig. 6d, e) and corals, may suggest occasional subaerial exposition. However, Fe oxide crusts may also indicate submarine weathering of pyrite under oxic conditions (e.g. Flügel 2004), which is more likely. Microfacies 2: Brachiopod/crinoid floatstone. The prevailing bioclasts are brachiopods (Fig. 6f: BL 12-32) and crinoids; bryozoans are less abundant. Often they occur together as a brachiopod– crinoid- or brachiopod–bryozoan-floatstone (Fig. 7a: BL 12-29C), but monospecific biota also occur, such as crinoid-, brachiopod- (Fig. 6f) or bryozoan-floatstones (Fig. 7a: BL 12-29C). The large bioclasts are embedded within a bioclastic wackestone matrix or within a lime-mudstone, as shown in Figure 6f. Large components show no abrasion and were transported a relatively short distance from the source area (many biotic components are complete) and show few effects as a result of compaction (e.g. very rare clay steams and/or stylolites). Sometimes, shells have been infested with boring organisms (Fig. 7b: BL 12-33(3)). Interpretation: Based on the sizes of the components, the preservation, particularly the accumulation of brachiopods (which may be more or less autochthonous) and facies/sedimentological criteria, a shallow-water low-energy environment (midramp setting, shallow subtidal) is inferred. The lower part of the section described herein can be correlated with another section sampled 20 m to the north (see Ernst et al. 2011), which also exhibited fenestrate bryozoans that are an indication of quiet water conditions and the presence of soft substrates (Ernst et al. 2011). Borings are common, which also suggests low sedimentation rates. Microfacies 3: Bioclastic wackestone to grainstone. This microfacies is composed of poorly sorted wackestones to grainstones (e.g. Fig. 7c: BL 12(1)), typically with a fine-grained micritic matrix that also contains quartz grains. Burrowing is absent. This lithofacies occurs in the middle part of the sequence and in the uppermost part (sample numbers BL 12-1 –BL 12-3), representing the youngest sediments of the section. Main bioclasts of up Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 6. (a) – (e) Allochthonous bioclastic wackestone, locally strongly burrowed due to an increasing number of endobenthic organisms. Main bioclasts are crinoids, brachiopods, bryozoans and trilobites; corals, stromatoporoids and ostracodes are less abundant. (a) Sample BL 12-23(2); (b) sample BL 12-24A; (c) sample BL 12-25; (d) & (e) occasionally small phosphorite intraclasts, pyrite crystals and iron-oxide crusts occur (d, sample BL 12-33(2); e, sample BL 12-34A). (f ) Brachiopod– crinoid-floatstone embedded in a lime-mudstone (sample BL 12-32). to 15% are composed of crinoids, trilobites, brachiopod shells and rare corals. Bioclasts are less abundant than in Microfacies 1 and they are much smaller. Furthermore, calcareous algae and gastropods are absent. Very small phosphorite clasts and some extraclasts occur (Fig. 7d: BL 12-3A(1)). Some layers show graded bedding. Wackestones to grainstones, which are dominated by crinoids (Fig. 7e, f: BL 12-12(2) and BL 12-14) and common occurrences of tube worms (Fig. 8a: BL 12-20(2)), also occur, but they have less palaeoenvironmental significance as they occur in shallow-, as well as in deeper-marine settings (Flügel 2004). Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 7. (a) Brachiopod– bryozoan-floatstone (sample BL 12-29C); (b) brachiopod-floatstone; brachiopod shells are often infested by boring organisms (sample BL 12-33(3)); (c) & (d) poorly sorted bioclastic wackestone with small phosphorites and intraclasts (c, sample BL 12-1; d, sample BL 12-3A(1)); (e) & (f) crinoid-dominated wackestones (e, sample BL 12-12(2); f, sample BL 12-14). Interpretation: The faunal association, the preservation, the poorly sorted sediment and the fine bioclastic micritic matrix suggest a generally lowenergy environment, in a mid- to deep-ramp (subtidal environment) setting. The occurrence of rare grainstones is interpreted as storm layers. Sedimentation (e.g. poorly sorted, reworking) is most probably associated with a sea-level rise. Microfacies 4: Microbioclastic wackestone. This microfacies exhibits a variable amount of biota, which ranges from below 5%, with large isolated fragments, such as ostracodes (Fig. 8b: BL 12-11(2)), up to 20% of shell hash of trilobites, brachiopods and ostracodes, among other organisms. Those layers can exhibit graded bedding. Rare bryozoans also occur (Fig. 8c: BL 12-9A). The main difference Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 8. (a) Large occurrences of tube worms in a partly sparitic wackestone (sample BL 12-20(2)); (b)–(c) microbioclastic wackestone with large fragments of ostracodes; sediment is strongly burrowed (b, sample BL 12-11(2); c, sample BL 12-9A); (d) matrix composed of quartz-bearing pelmicrite (sample BL 12-15(2)); (e) peloidal calcisiltite with rare bioclasts (sample BL 12-30). to Microfacies 1 is the size of bioclasts (microbioclasts in contrast to large bioclasts in Microfacies 1) and the matrix. The matrix is a quartz-bearing, burrowed pelmicrite (Fig. 8d: BL 12-15(2)) containing occasionally larger crinoid and echinoderm fragments. Bryozoans are very rare. Burrowing is a common feature. A bioturbation index, in which a descriptive grade was assigned to the degree of bioturbation, was established by Taylor & Goldring (1993). The degree of bioturbation in this section can be assigned to grade 4, which is characterized by high abundance and density. This facies occurs mainly in the upper part of the section. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Interpretation: This microfacies suggests generally a low-energy, outer-ramp position. Those layers that show graded bedding are interpreted as material reworked during higher-energy events (e.g. storm deposits). Microfacies 5: Peloidal calcisiltite. The rock is composed of densely packed micrite peloids and silt-sized terrigenous quartz grains. Bioclasts are very rare, and are composed of crinoid ossicles and brachiopod remnants. Small-scale lamination is common. This facies type occurs only in layers BL 12-30 and BL 12-31 in the lower part of the section (Fig. 8e: BL 12-30). Interpretation: The densely packed, quartz-rich calcisiltite could be interpreted as a post-drowning phase of a carbonate platform comparable to an example from Morocco (e.g. Blomeier & Reijmer 1999). Similar finely laminated lime mudstones also occur in very shallow intertidal and supratidal environments (e.g. Gammon & James 2001), but these facies also show microbial mats, fenestral fabrics and/or desiccation features, which do not occur in Microfacies 5. Therefore, we interpreted the latter as representing a deeper, low-energy setting. Magnetic susceptibility data The mean initial magnetic susceptibility for the whole section is 2.25 × 1028m3 kg21, which is a little bit lower but in the range of the MS marine standard ¼ 5.5 × 1028 m3 kg21 – the median value for about 11 000 lithified marine sedimentary rocks, including siltstone, limestone, marl and shale samples (Ellwood et al. 2011). The lowest susceptibility is 0.45 × 1028m3 kg21 (no negative values, indicating that the section is not composed of very pure carbonates) and the highest value is 9.15 × 1028 m3 kg21. The section is composed of an alternation of carbonates and shales, and both lithologies were sampled. It clearly appears that the shale levels have systematically higher susceptibility (mean value of 4.27 × 1028 m3 kg21) than the carbonates (1.27 × 1028 m3 kg21) (Fig. 9). Ms (magnetization at saturation) and xFerro (ferro-magnetic susceptibility) are both proxies for the concentration of ferromagnetic minerals. The contribution of Ms is rather low (5.02 × 1024 and 12.62 × 1024 A m2 kg21) and Ms data are not well correlated with xin (correlation factor r ¼ 0.40: Fig. 9. Lithological column (scale bar in cm) and magnetic susceptibility evolution compared with magnetic hysteresis and backfield curve data on selected samples. xin, low-field magnetic susceptibility; Ms, magnetization at saturation; xhf, high-field magnetic susceptibility; xferro, ferromagnetic susceptibility. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 10. Comparison between magnetic susceptibility values (xin) and hysteresis parameters from selected samples from Blankenheim section: (a) Ms, magnetization at saturation, r ¼ 0.40; (b) xferro, ferromagnetic susceptibility, r ¼ 0.54; (c) xHF, high-field magnetic susceptibility, r ¼ 0.98; (d) Bc, coercivity, r ¼ 0.40. Fig. 10a). xFerro values are also very low (0.01 × 1028 and 0.86 × 1028 m3 kg21) and correlate weakly with xin (r ¼ 0.54: Figs 9 & 10b). These low correlations observed between x and proxies for ferromagnetic minerals are good arguments in favour of a low influence of ferromagnetic minerals. xHF is a proxy for the sum of the paramagnetic and diamagnetic minerals. xHF values are between 0.26 × 1028 and 6.70 × 1028 m3 kg21 (Fig. 9). All values are positive, indicative of a relatively large amount of paramagnetic minerals. There is a very strong correlation between xHF and xin (r ¼ 0.98: Figs 9 & 10c), indicating that, in this case, the magnetic signal is mostly linked to paramagnetic minerals. This fits well with the fact that the mean susceptibility values are systematically higher for clay mineral levels. It is possible to obtain information on the nature of the ferromagnetic minerals through the backfield curve. The high-field remanence saturation represents the proportion of non-saturated (at 300 mT) magnetic minerals (high-coercivity minerals, such as hematite, as opposed to low-coercivity minerals, such as magnetite, that are easily saturated). High-field remanence is between 0 and 20% for 10 samples (small amount of high-coercivity minerals, such as hematite), between 20 and 50% for seven samples (larger amount of high-coercivity minerals), and higher than 50% for three samples (Fig. 9). The shape of the hysteresis loops, Bcr/Bc and Mrs/Ms, provide information on the magnetic grain size, with the help of the Day plot (Day et al. 1977; Dunlop 2002). Main grain-size categories correspond, from the coarser to the finest, to multidomain (MD), pseudo-single domain (PSD), single domain (SD) and super paramagnetic (SP). Coarser grains are often interpreted as related to a detrital origin, while the smallest grains are interpreted as having formed during diagenesis. The samples from Blankenheim fall along the PSD and SD + MD mixing curve (Dunlop 2002), with only one sample (S12) along the SD + SP mixing curve (Fig. 11). The remanence decay is also indicative of the grain size. The relative decay viscosity coefficient (Sd, a dimensionless quantity) was calculated from the remanence decay measured over 100 s. Sd represents the slope in the IRM2500 mT v. Log(time) diagram. The Blankenheim samples range between 6 × 1023 and 30 × 1023, with only one sample (S12) reaching 116 × 1023 (Fig. 9). According to Spassov & Valet (2012) Sd values between 3.9 and 6.5 × 1023 are indicative of SD grains, while values above 100 are indicative of Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 11. Bi-logarithmic Day plot (Day et al. 1977) for selected samples from the Blankenheim section, compared with the mixing curves by Dunlop (2002), with classic remagnetized limestones (Jackson 1990), MD magnetite grains (Parry 1982) and with results from Zwing et al. (2005) on different lithologies from NE Germany (near Cologne). The majority of the Blankenheim samples fall on the left-hand side of the plot, within the PSD and SD + MD mixing curve. SP + SD grains. This fits relatively well with the results obtained with hysteresis data, with a majority of the sample with relatively coarse grains and only one sample with very-fine SP granulometry (sample S12). Interpretation From the magnetic measurements, we can deduce that we have a magnetic signal, which is mostly carried by paramagnetic minerals (probably clay minerals considering the clear link between claydominated intervals and magnetic susceptibility highs: Fig. 12). Ferromagnetic (sensu lato) minerals are present in very low abundance, and they are dominated by magnetite and hematite, but these minerals have a small influence on the total bulk magnetic signal. The Day diagram (cf. Fig. 9) indicates that the ferromagnetic mineral content consists of a mixture of SD and MD, dominated by MD grains (i.e. coarse grain sizes indicating detrital origin). Only one sample shows the influence of smaller grains (BL S12). Hematite is the most abundant in this sample, suggesting that it was subjected to a stronger or different diagenetic pathway. The Rhenohercynian zone has experienced a widespread remagnetization event, identified through detailed magnetic studies in Belgium (Molina Garza & Zijderveld 1996; Zegers et al. 2003; Da Silva et al. 2012, 2013) and in the NW RSG in Germany, near Cologne (Zwing et al. 2005). The latter study compared magnetic susceptibility data from biohermal carbonate rocks, platform carbonate rocks and siliciclastic rocks (Fig. 11). They identified the most important carrier of the late Palaeozoic magnetization component as magnetite. However, results were different, depending on the lithology, with MD (detrital) magnetite dominating in the siliciclastic rocks and SP (diagenetic) magnetite in the biohermal carbonates, with the platform carbonates showing intermediate hysteresis properties. Zwing et al. (2005) interpreted the remagnetization as a widespread event, affecting all lithologies, but that this remagnetization would be disguised by the large amount of MD magnetite in siliciclastic rocks. This remagnetization is interpreted as related to complex processes. In the Late Devonian and Early Carboniferous sedimentary rocks, the remagnetization is coeval with clay diagenesis (through the smectite to illite transition releasing iron); in the Middle Devonian strata, however, clay diagenesis and remagnetization are not coeval and pyrite oxidation processes are observed (Zwing et al. 2009). In the case of the Blankenheim section, either the MD detrital magnetite ‘hides’ a remagnetization event, as in Zwing et al. (2005), or the MD magnetite has a detrital origin and no remagnetization processes occurred in this section. Helsen & Königshof (1994) have shown that the conodont alteration Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 12. Lithological column showing magnetic susceptibility evolution (in carbonates and shales), stable carbon and oxygen isotopes of bulk rocks, palynomorphs and TOC and sulphur data. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. index from Lower and Middle Devonian rocks within the Eifelian area indicate little, if any, effect of heating, with alteration index ranging between 1.5 and 2.0, indicating very low temperatures of approximately 55 8C. From thin-section observations, the alteration appears to be very weak, with no indication of pressure solution and with remarkable preservation of fossils (see above). This is in accordance with the generally low Sd values, which indicate that SP grains are not very common in the samples. However, one would expect an SP presence in remagnetized samples. Low alteration and low compaction are explained by a reduced sedimentation overburden (Helsen & Königshof 1994). Figure 11 displays a Day plot, which includes the data from Jackson (1990), corresponding to classic remagnetized limestone samples (Channel & McCabe 1994), and from Parry (1982), corresponding to coarse-grain magnetite. Figure 11 also includes data from Zwing et al. (2005) on biohermal carbonates, platform carbonates and siliciclastic Devonian rocks, with different magnetic properties observed for the different lithologies. Our Blankenheim data are on the left-hand side of the plot, separated from the other results, and follow the PSD and SD + MD mixing curve. Thus, the data correspond to grain size probably even coarser than that observed by Zwing et al. (2005) on the siliciclastic rocks in the area of Cologne. The carbonate and the shale intervals in Blankenheim are carrying magnetite of the same grain size despite the lithological variations. This is in opposition to Zwing et al. (2005), in which the carbonates have smaller grains. In Zwing et al. (2005), the different grain size related to different lithologies was explained by the remagnetization event. The steady magnetite grain size in Blankenheim is in favour of a detrital origin for the magnetite grains, without any impact of a remagnetization event, although this needs to be cross-validated by further studies. Furthermore, a comparison of magnetic susceptibility results with geochemical data, and specifically with elements that are considered as proxy for detrital inputs, such as Ti, Al, Rb and Zr (Tribovillard et al. 2006; Calvert & Pedersen 2007), allows us to assess the influence of detrital inputs on the magnetic susceptibility signal (Riquier et al. 2010; Da Silva et al. 2012, 2013). The correlation is high (the link between magnetic susceptibility (MS) and detrital proxy elements on 45 samples: for Al2O3, r ¼ 0.84; for TiO2, r ¼ 0.83; for Rb, r ¼ 0.79; for Th, r ¼ 0.81; and for Zr, r ¼ 0.72), indicating a major impact of detrital inputs on the MS signal. As mentioned earlier, the magnetic susceptibility of shale layers is systematically higher than the adjacent carbonate beds, a link not always observed. For example, Bertola et al. (2013) have shown that the magnetic susceptibility values were fairly similar in the carbonates and in the adjacent shale beds in two Tournaisian sections from Belgium (Bertola et al. 2013, fig. 10). These results on Tournaisian Belgian rocks also suggest that the input of MS-carriers stayed roughly constant during the deposition of the limestone shale alternation and that these carriers were probably ferromagnetic. In the Blankenheim section, magnetic susceptibility appears more as a proxy for shale proportion (paramagnetic minerals), which can be noted from outcrop visual observation. However, Figure 9 displays a plot with separate magnetic susceptibility curves for carbonates v. shales. Both curves show relatively similar trends, with relatively high peaks in the lower part of the section and decreasing magnetic susceptibility towards the top, which could be related to a decrease in detrital inputs, in relation with the observed transgressive trend. Ellwood et al. (1999) proposed that during transgression, sea level increases and the portion of landscape exposed to erosion decreases, leading to a decrease in detrital input. This link was actually observed in various carbonate platform examples (Hladil 2002; Racki et al. 2002; Da Silva et al. 2010; Whalen & Day 2010). This interpretation fits well with the observed deepening trend reflected by the facies. Geochemical proxies Carbon and oxygen isotopes The carbon isotope values of bulk rock samples (Fig. 12) fall within a range of 22.1 and 1.7‰. A positive excursion of d13Ccarb values from 20.1‰ (BL 12-34B) to 1.5‰ (BL 12-31) is observed at the base of the section. The interval between BL 12-31 and BL 12-22 is characterized by a gradual decrease in d13C from 1.5 to 20.2‰. The second positive excursion lasts from BL 12-22 until BL 12-16, peaking at 1.2 ‰. Further, two positive peaks of d13Ccarb values are recorded within the 5–6 and 6–7 m of section separated by low values at BL 12-15 (0.9‰) and BL 12-9A (0.1‰). The values decrease between BL 12-7 and BL 12-3B, ranging from 1.6 to 22.1‰ (lowest d13Ccarb value obtained across the entire section). Two positive peaks occur in BL 12-6A and around BL 12-S4. Carbon isotope values of prasinophytes fluctuate between a minimum of 229.7‰ and a maximum of 225.7‰. The trend of d13Corg within the lowermost 1 m of the section corresponds to the trend of d13Ccarb of the same interval. The highest d13Corg value is recorded in BL 12-14. A negative shift of 23.6‰ is observed from BL 12-14 to BL 12-12. Below this negative shift, the values are relatively higher than those above the shift and show an amplitude fluctuation ranging from 228 to 225.7‰ and from 229.7 to 228‰, respectively. The lowest value Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT of d13Corg, as well as d13Ccarb, is observed around 7.5–8 m of the section. The value increases towards the top of the section from 229.7‰ (BL 12-3) to 229.1 ‰ (BL 2-1). The increase of values in the uppermost part of the section is also observed in MS (carbonate) and d13Ccarb. Although we provide oxygen isotopes of bulk sedimentary rocks, we are aware that d18O analyses from bulk samples do not produce reliable palaeoenvironmental results compared with data from the calcite of brachiopod shells or the PO4 group of biogenic phosphate. In this paper, d18O values are plotted to compare variations through the section with other geochemical and geophysical proxies. d18O values of bulk-rock samples range between 26.7 and 23.1‰. TOC and sulphur content Total organic carbon is below 0.7% throughout the entire section (Fig. 12), with a minimum of 0.12% (BL 12-20) and a maximum of 0.69% (BL 12-S6). Sample BL 12-34B from the base of the section shows a TOC concentration of 0.24%. The content decreases to 0.16% and then increases to 0.36% approximately 1 m above the section base (BL 12-31). A second positive peak (0.51%) is observed at BL 12-S24, followed by decreasing values to approximately 3.5 m. Another positive shift to 0.43% is observed between BL 12-22 and BL 12-21. The following drop to 0.12% (BL 12-20) equates to the lowest value measured across the entire section. Two minor positive spikes are recorded at BL 12-18 (0.34%) and BL 12-14 (0.43%) between BL 12-20 and BL 12-7, followed by a low-amplitude fluctuation to approximately 7 m above the section base. Sulphur content varies between 0.02 (BL 12-S5) and 1.54% (BL 12-31). The significant positive shift near the base of the section peaks at the same level, with the first maximum in the d13Ccarb and the TOC curves. Sulphur content then decreases to 0.36% (BL 12-29C), followed by a minor positive excursion (maximum 0.73%: BL 12-29B). Thereafter, values decrease to 0.04% (BL 12-S24), followed by an interval of low-amplitude fluctuation through to sample level BL 12-13. Above this interval, although values increase to 0.26% (BL 12-11), and to 0.33 –0.41% between BL 12-S8 and BL 12-7, they consistently show less fluctuation than the section below. Estimated palaeotemperature In recent years, the oxygen isotope composition of conodont apatite has been used to estimate the palaeotemperature (e.g. Joachimski et al. 2004, 2009; Trotter et al. 2008). In this study, we measured exclusively icriodontid platform elements of two samples (BL 12-29c and BL 12-22). d18Oapatite values of both samples show a ratio of 19.2‰ (1 SD ¼ +0.2‰). Assuming an oxygen isotope composition of 21‰ for Middle Devonian seawater, a palaeotemperature of 29.7 8C is calculated using the temperature equation provided by Pucéat et al. (2010). According to Joachimski et al. (2009), who summarized d18Oapatite data of conodonts from Germany, France, the Czech Republic and the United States for the Middle Devonian using a value of 22.6‰ for NBS120c, the d18Oapatite values ranging from 19 to 21‰ (VSMOW) gave palaeotemperatures from 22 to 30 8C. Major, trace and REE analysis Principal component analysis (PCA) of normalized whole-rock values was conducted using PAST software (Hammer et al. 2001). The first two principal components contribute meaningful information for the interpretation of geochemical patterns throughout the section (Fig. 13). PC-1 values are interpreted as a detrital signal due to the positive and negative loadings of siliciclastic detrital indicators v. carbonate indicators (Fig. 14). Conversely, PC-2 values are consistent with the signals indicated by d13C and TOC (Fig. 15), even if they are not proxies for anoxia. Plotting individual signals against stratigraphy provides a clearer signal of detrital input (Fig. 16) and events (Figs 17 & 18). Given the tectonic complexity of the area, we analysed samples to determine sediment provenance. Trace elements, such as Hf, Nb, Sc, Ta, Th, U, Y, Yb and Zr, are found in the titanium-bearing detrital fraction of sedimentary rocks and can be used to distinguish the maturity (Carpentier et al. 2013) or the tectonic environment (Wood 1980; Pearce et al. 1984; Bhatia & Crook 1986) of the source rock. Analysis of Th/U ratios in sediments in the Blankenheim section suggests that the sediment source is juvenile, with Th/U ,3 (Fig. 19). In addition, trace-element geochemical signatures suggest that the sedimentary provenance for the detrital material is consistent with a continental island arc environment (Fig. 20). This interpretation supports earlier research that the Blankenheim section was part of the newly amalgamated Avalonian microcontinent (Franke & Oncken 1995; Franke 2000), as well as recently published data of the eastern part of the Rheinisches Schiefergebirge (e.g. Eckelmann et al. 2014). Detrital signatures are seen between 61 and 148 cm (around samples BL 12-32 and BL 12-29B) above the base of the section, with two major pulses of detrital input higher up in the section at 738 and 759 cm (around samples BL 12-S5 –BL 12-3C: Fig. 16). Based on increases of chalcophile elements, such as Zn, Pb, Cu and S Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 13. Principal component analysis (PCA) of normalized whole-rock values showing PC1 v. PC2. PC1 is interpreted as a detrital signal, with PC2 showing an anoxia signal. (Fig. 17), a signature of anoxia is suggested at 89 cm above the base of the section (samples BL 12-31 and BL 12-30), although redox proxies such as authigenic U and Ce anomalies do not support anoxia in this interval (Fig. 18). A second event at the top of the section (from 738 to 779 cm) possibly indicates anoxia through a set of staggered negative d13C excursions, a Ce anomaly and increases in TOC (Fig. 18). This upper event begins with the pulse of detrital input at 738 cm and continues through 779 cm. Although there is a spike in excess Ba, excess P and V/Cr at 365 cm (Fig. 18), we are reluctant to assign a meaning to this interval as V/Cr is a proxy for anoxia, while excess Ba and P are proxies Fig. 14. PC1 loadings showing the detrital signal in contrast (positive values) to the carbonate signal (negative values). Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 15. PC2 loadings showing the anoxia signals, with positive loadings in chalcophile elements and P2O5. Fig. 16. Stratigraphic distribution of detrital proxies in addition to d13C and d18O signatures. There is a major excursion at 738 cm and at 759 cm, which we interpret as a major sediment influx. Note that TOC and d18O are correlated with this pulse of detrital input. An earlier flux of material can be seen at 89 cm above the base of the section. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 17. Stratigraphic distribution of chalcophile element proxies in addition to d13C and d18O signatures. There is an excursion of the chalcophile elements at 89 cm above the base of the section. for primary productivity, which are not generally seen in the same interval in known anoxia events (Dymond et al. 1992; Pujol et al. 2006), and because the use of Ba as an indicator is dependent on a variety of local water-column conditions (Von Breymann et al. 1992; Paytan et al. 2007). In addition, these signals are not replicated by the other proxies used in this study. While TOC has been used as a proxy for anoxia in many studies (Ingall et al. 1993; Algeo & Maynard 2004; Pujol et al. 2006; Marynowski & Filipiak 2007; D.P. Bond et al. 2013), it may not be an appropriate proxy to use in this case, as it has a positive correlation with detrital input (Fig. 15). Interpretation In the lower event (89 cm), elevations in chalcophile element concentrations are not correlated with redox-sensitive element anomalies in comparison to the upper potentially anoxic horizon (738 – 779 cm) (Fig. 17). This may possibly be due to the differences in sediment supply and flux to the environment; the lower horizon shows a gradual increase in detrital fraction elements above the increases of chalcophile elements. The reason for this discrepancy between anoxia proxies is unclear, and may represent bacterial sulphate reduction in buried sediments rather than an anoxic event at the sediment –water interface. At the top of the section, however, the sharp spike in redox-sensitive elemental anomalies and a negative excursion in d13C concurrent with, and subsequent to, two large sediment pulses may indicate anoxia at the sediment –water interface. Discussion and conclusions The entire section is composed of shallow subtidal to moderately deep subtidal mixed carbonates and siltstones. According to Struve (1990), the stratigraphical extent of the so-called ‘Great Gap’ period lasting from the lower Eifelian into the lower Givetian is characterized by sedimentary gaps and not full-marine sediments, and may also have been recognized in the Couvin area, Belgium (Bultynck & Hollevoet 1999). Phosphorite intraclasts and ironoxide crusts around some bioclasts have been found Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 18. Stratigraphic distribution of anoxia (V/Cr) and productivity proxies (normalized Ba, excess P2O5) in addition to d13C and d18O signatures. There is a major excursion at 365 cm, which we are reluctant to classify as an anoxic event or an event correlated with changes in sedimentation. in some thin sections, which may be an indication of submarine weathering and does not necessarily mean that an area further in the north underwent subaerial exposure, as suggested by Winter (1977). Thus, we favour a shallow-marine environment for the entire section, which shows an overall slightly transgressive trend. The latter is also confirmed by magnetic susceptibility data. According to DeSantis & Brett (2011), magnetic susceptibility data suggest a signature for a regressive phase in eastern North America for deposits of the Stony Hollow, followed by the transgressive succession of the Hurley–Cherry Valley deposits (Brett et al. 2011). Magnetic susceptibility data of the Blankenheim section show similar patterns in the same stratigraphic position and are also comparable to those described from the GSSP Mech Irdane, Tafilalt, Morocco (Crick et al. 2000). It might be possible that the lower event at the base of the Blankenheim section is based on regional variations in sediment supply and/or sea-level changes, or may be correlated with the Stony Hollow Event associated with probable warming and incursion of tropical species into the subtropical to temperate shelf region of eastern North America (DeSantis & Brett 2011). As shown earlier, geochemical proxies of the Blankenheim section exhibit two major signals: an increase in chalcophile elements that occurs at the base of the section at 89 cm within the kockelianus conodont Biozone; and a second peak that occurs from 738 to 779 cm from the base of the section. Above the alluvial sediment influx at 738 cm, there is a large negative excursion in d13C and a Ce anomaly ,20.1 (Fig. 15), the latter of which has been seen in other sections with anoxia (Morad & Felitsyn 2001; Pujol et al. 2006; Carmichael et al. 2014). Although spikes in V/Cr are not readily apparent in this interval, they do show increases from a secular low (Fig. 17). Any potential authigenic U signatures are likely to have been overprinted by high levels of detrital Th in this part of the section, invalidating authigenic U as a potential anoxia tracer in this particular location. Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. Fig. 19. (a) Th/U ratios in basinal sediments can be used to determine the sediment source, with Th/U values for mature, cratonic sediments with highly variable ratios (average of c. 6), and juvenile sediment sources with Th/U values clustered around 3 (modified from Carpentier et al. 2013). (b) The source of sediments in the Blankenheim section has a clustered Th/U value of approximately 2, suggesting a juvenile source. While there is a phase offset between the d13C and Ce anomaly at 779 cm in comparison with the detrital pulses at 738 cm and 759 cm, this could be due to bacterial reduction of the organic matter in the sediments, which will lead to carbon isotope fractionation (Kump et al. 1999). This mechanism is consistent with our observation that TOC in this section is often correlated with detrital sedimentation, and the negative d13C excursion seen may be due to bacterial reduction, fractionation and mobilization of accumulated organic carbon in the detrital sediments below. While this explanation for a negative excursion is highly dependent on local conditions, negative excursions in d13C have been seen immediately prior to the Kačák Event both in Ontario, Canada (van Hengstum & Gröcke 2008) and in Morocco (Ellwood et al. 2003). The d13C excursion and Ce anomaly are, therefore, interpreted to be the local representation of the beginning of the ‘true’ Kačák Event Interval, which is also consistent with the conodont and microfacies analyses presented above. The major conclusions of our multidisciplinary approach are as follows: † Based on micro- and macrofossils, the Blankenheim section exposes the Junkerberg and Freilingen formations, and reaches the uppermost Eifelian (kockelianus and ensensis conodont biozones). Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 SHALLOW-WATER FACIES AROUND THE KAČÁK EVENT Fig. 20. Tectonomagmatic discrimination diagrams showing sediment-source signatures in detrital Ti-bearing minerals from (a) basalts (modified from Wood 1980), (b) greywacke sediments (modified from Bhatia & Crook 1986) and (c) & (d) granitic sources (modified from Pearce et al. 1984). Regardless of the signature suite used, the Blankenheim sediments clearly cluster as arc-volcanic signatures, consistent with deposition along the newly amalgamating Avalonia microcontinent. † The entire section is composed of shallow- to deep-subtidal mixed carbonates and siltstones. The beginning of the Kačák Event Interval can be recognized even in the shallow-marine environment (open-marine organisms, such as tentaculites, cephalopods and/or pelagic ostracodes are absent). The overall section shows a slightly transgressive trend. † From the magnetic measurements, we can deduce that we have a magnetic signal, which is mostly carried by paramagnetic minerals. The magnetic susceptibility of shale layers is systematically higher than the adjacent carbonate beds. Furthermore, hysteresis plots point to preserved primary coarse (MD) detrital magnetite, which is classically not the case in most of the remagnetized Rhenohercynian zone. Magnetic susceptibility (MS) shows a generally decreasing trend in the section, which could be related to a decrease in detrital input, consistent with the observed transgressive trend. † The lower part of the sequence (89 cm above the base of the section, around samples BL 12-31 and BL 12-30) exhibits a correlation of the occurrence of chalcophile elements with the occurrence of quartz-rich calcisiltite. This event may be a result of local variations in sediment supply and/or sea-level changes or may be Downloaded from http://sp.lyellcollection.org/ by guest on June 10, 2015 P. KÖNIGSHOF ET AL. associated with the pre-Kačák Stony Hollow Event. The d13C excursion and Ce anomaly in the upper part of the section (738 –779 cm) are interpreted to be the local representation of the beginning of the ‘real’ Kačák Event Interval, which is also consistent with the conodont and microfacies analyses presented herein. † The conodont apatite oxygen isotope record from our samples is interpreted as reflecting the palaeotemperature of the Devonian tropical and subtropical sea surface. The estimated palaeotemperature of our samples is 29.7 8C. † In terms of plate tectonics, the analysed trace elements, such as Hf, Nb, Sc, Ta, Th, U, Y, Yb and Zr, suggest that the sediment source is juvenile and consistent with the Avalonian microcontinent. 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