The Devonian, Carboniferous, and Permian periods in time scales 1989 – 2006. 

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... Group is composed of four formations. All are dominantly siliciclastics, in contrast to the earlier carbonate dominance in the Onondaga Formation. This change in deposition was caused by the second pulse of the Acadian Orogeny, which created an extensive highland area east of New York State. The Marcellus Formation is composed mainly of dark shales and siltstones which represent the initial influx produced by the Acadian Orogeny. Its upper part is marked by anoxic dark shales which indicate the Ka č ák Event (cf. 3.3.1). Above this level, in central New York State the Skaneateles Formation undergoes a transition from basinal shales to a sandy shallow shelf, succeeded by the Ludlowville and Moscow formations with their succession of fossiliferous mudstone, siltstone and sandstone. The Hamilton Group is succeeded by the Tully Limestone, the age of which has been controversial, but now known to be within the upper portion of the Polygnathus varcus Zone (cf. Rickard, 1975; Kirchgasser, 2000). It thus represents the Taghanic Event, which is regarded as the last distinct global event of the Middle Devonian (e.g. Aboussalam, 2003). The Tully Formation with its shallow marine carbonates was abruptly covered by thick black shales of the Genesee Formation. This major change of post-event facies is regarded as a widespread transgression, and thus marks the onset of the Acadian Tectophase III sensu Ettensohn. The overlying Genesee Group extends into the Late Devonian. The GSSP of the Late Devonian Epoch, and the coincident base of the Frasnian Stage, is at Puech de la Suque, Montagne Noire, France, within the lower Clapassous Formation (Klapper et al., 1987; House et al., 2000). The boundary is indexed by the conodont Ancyrodella rotundiloba early morphotype sensu Klapper. The FAD of this taxon, which is present in both pelagic and neritic facies, almost coincides with the base of the Frasnes Group in Frasnes, Belgium. The level of the boundary is substantially higher than that of previous opinions on the Late Devonian base, which approximated to the base of the Assise de Fromelennes Formation which underlies the Assise de Frasne in Belgium. There were, however, prolonged controversies on the taxonomy of the index conodont and therefore on its position within the conodont zonation (particularly by Ziegler and Sandberg, 1996). The latter authors finally placed the boundary within the Early M. falsiovalis Zone, i.e. at the top of the Polygnathus norrisi Zone (see Weddige, 1996). For accompanying fossil groups, the boundary is marked by a late stage of radiation of the pharciceratid ammonoids. A new goniatite record, Neopharciceras , together with conodont index fossils, permits world-wide correlation. In contrast, a short anoxic pulse, the so-called Frasnes Event, which lies slightly above the base, is of less global significance. The GSSP for the base of the Famennian Stage was placed within the upper Coumiac Formation slightly above the Upper Coumiac Quarry, near Cessenon, Montagne Noire, France (Klapper et al., 1993; House et al., 2000). The stratigraphic level is very close to that of the base of the Famennian as formerly used in the Famenne area in Belgium (Bultynck and Martin, 1995). The boundary is indexed by the entry of Palmatolepis triangularis . The taxon is derived from the Palmatolepis praetriangularis stock demonstrating a proven phylogenetic lineage. Thus, the taxonomy of the boundary index is generally accepted, whereas the lithologic position of the GSSP is still under review. The GSSP, with respect to the FOD of P. triangularis , is immediately above a global mass extinction event, i.e. the last event of the Kellwasser Events, when, for example, the goniatites Manticoceras and Beloceras become extinct (e.g. Schindler, 1990; House et al., 2000). A haematite crust then, immediately above the GSSP spike, obviously represents a hardground with condensation, and it is suspected that the GSSP spike itself seems to mark nothing else than a gap in sedimentation, i.e. of documentation (e.g. Ziegler and Sandberg, 1996). Geochemical, magnetostratigraphic and other investigations are still in progress. The numerical ages for the Late Devonian stages of the DCP 2003 were selected as young as possible using ID-TIMS ages of Tucker et al. (1998) and their error bars (cf. Weddige et al., 2005: Taf. IV). This is necessary to get a balance with the Early Carboniferous SHRIMP ages of Roberts et al. (1995) which are too young because of the heterogenity of the standard SL13 zircon used to determine them. For over 10 years the Famennian has been regarded as considerably longer than the Frasnian Stage, not least because of its division into a greater number of conodont and ammonoid zones. However, this time-relation has yet to be confirmed by a reliable isotopic age determination. Only in the time scale of Compston (2000), using SHRIMP and selected/ reinterpreted MSID ages, is the Frasnian longer than the Famennian Stage (Fig. 1). Since the work of Tucker et al. (1998) the Famennian has a duration of about 15 my (Table 2) and, therefore is one of the longest Phanerozoic stages. In the Rhenish Massif, the regional historical subdivision of the Late Devonian into Adorf-, Nehden-, Hemberg-, Dasberg- and Wocklum-Stufe, was mainly introduced by Wedekind (1913). This regional subdivision has been used extensively by survey geologists for intensive mapping for about one century. Moreover, pioneering conodont studies started in the Rhenish Massif, where the Late Devonian pelagic limestone succession contain abundant conodont faunas, which have supported and promoted the usage of the regional stratigraphy. The discrepancy with the new global (GSSP) stages, however, is actually not serious, because the basal boundaries of the global Frasnian and the historical Adorf-Stufe are nearly coeval, but the latter is slightly younger (cf. Fig. 2). In contrast, the global Famennian corresponds to four regional historical stages which now could be taken as substages of a four-partite Famennian. The Late Devonian of the Sauerland starts, locally, immediately above the “ Roteisenstein-Horizont ” , which is derived from volcanic exhalative activities (Fig. 2). Then follow the reefal Adorf-Massenkalk formations and bioherms, which coincide with cephalopod limestones, or inter-reefal siliciclastic Adorf- Bänderschiefer (Fig. 2). These facies variants demonstrate a topographic wave pattern with small rises and troughs. Towards the end of the Frasnian, the reef growth generally slowed and finally stopped because of subsidence, too rapid for reef builders to match. All reefal development was terminated by two anoxic pulses, which are globally recognizable as the Early (Lower) and Late (Upper) Kellwasser events (Schindler, 1990). The latter is characterized by a mass- extinction (cf. Section 3.4.1). Later, slow pelagic sedimentation during the Famennian gradually levelled the Frasnian rise-trough relief. Red or yellow shales, ostracod shales or fine-grained silt- and sandstones or Plattenkalke filled up the troughs and thin cephalopod limestones, with enriched pelagic fossils, covered the rises (Fig. 2: Nehden-, Hemberg- and Dasberg- Schichten). In the Sauerland – as in the Harz Mountains and in Thuringia, too – this Famennian sedimentation pattern was interrupted twice by the semiglobal Annulata Event, Hemberg-Stufe and the global Hang- enberg Event, late Wocklum-Stufe, anoxic pulses (Fig. 2; cf. Walliser, 1996). The marine Frasnian, dominantly carbonates, covers most of the East-European Platform. The lower boundary of the Frasnian, which is fixed by the FAD of the conodont A. rotundiloba, is difficult to trace because this level is in a clastic part of the succession. However, in the Timan area (north-eastern platform) the joint occurrence of A. rotundiloba and spore assemblage of the Timanian and Sargayan regional substages indicates that the base of the Frasnian lies within or at the base of the Timanian. Besides these two substages, the other 5 Frasnian substages contain abundant shallow-water conodont polygnathid assemblages which have been correlated to the standard zonation by Ziegler et al. (2000). In the Ural Mts. the FAD of the conodont A. rotundiloba in limestone of the Sargaevo Formation, marks the lowermost Frasnian Stage. The overlying formations are composed of shallow water limestones, occasional reefs, biostromes, and thin beds of chert, argillites, marls, and limestone. This succession contains the Polygnathus timanicus – Ancyrognathus triangularis, Palmatolepis gigas, P. rhenana and P. linguiformis zones. It has long been known that a hiatus exists between the latest Frasnian (Livnian) and the earliest Famennian (Zadonskian). This gap spans mainly the basal Famennian P. triangularis Zone. But, rocks of this age (Volgogradian) occur in some restricted deep depressions located in the eastern part of the platform (Galushin and Kononova, 2004). The Famennian is represented mainly by carbonates, often with evaporitic intervals in its upper part. It includes 9 regional substages (Fig. 2). The Famennian conodont assemblages are dominated by shallow-water endemic taxa, but spore zonation gives some arguments for the correlation with the West European succession (Avkhimovich et al., 1993). In the Ural Mts. the Famennian Stage is mostly represented by a shallow-water limestone succession defined by the P. triangularis, P. crepida, P. rhomboi- expansa , and Siphonodella praesulcata conodont zones. Foraminifers, brachiopods, and ammonoids are also used to divide and correlate the Famennian. The short stratigraphic break between the Famennian and Tournaisian stages lies at an undetermined position within the succession. Late Devonian sections in the Zeravshan Ridge are represented by bedded and massive shallow-water limestones. They are characterized by diverse conodont associations, with almost all zonal index-species (Kim et al., 1984, ...

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... guide fossils and by interfacial correlations. Thus, the accompanying conodont Latericriodus steinachensis appears just below the GSSP boundary. The dacryoconarids Nowakia sororcula and Nowakia acuaria enter shortly above the boundary, and chitinozoans are useful because they can be linked to the spore zonation. Lithologically, the early Pragian rocks become gradually lighter in colour, although their overall facies remains similar to that of the Lochkovian (Chlupá č , 2000). A relatively minor faunal change in the early Pragian marks the rather insignif- icant Lochkovian – Pragian Event. The GSSP for the Emsian Stage is within the Khodzha – Kurgan Formation of the Zinzilban Gorge of the Kitab National Park in Uzbekistan (Yolkin et al., 1997; Yolkin et al., 2000). The boundary is indexed by the conodont Eocostapolygnathus kitabicus which is derived, and well differentiated, from its ancestor Eoctenopolygnathus pireneae , and thus represents a proven biostratigraphic time marker (Yolkin et al., 1994). The level of this phylogenetic transition is lower than the entry of the somewhat controversial taxon Eolinguipolygnathus dehiscens , and the entries of both taxa are lower than the top of the regional historical Pragian Stage. The rarity of conodonts in the late Pragian of the type area, however, does not permit an exact correlation with the Emsian GSSP stratotype in Uzbekistan. When Yolkin et al. (1994) proposed such a low Emsian base at the expense of the historical Pragian Stage, the authors intended to introduce a further stage for the latest Early Devonian. The early interval of the global Emsian corresponds with a part of the late original Pragian Stage in its Czech type region, and with the so-called Zinzilbanian in Uzbekistan, Central Asia ( Fig. 2). For the middle and late intervals of the global Emsian, the Czech terms Zlichovian and Dalejan are currently used informally when open marine to pelagic guide fossils like conodonts, dacryoconarids, trilobites, and goniatites are available. But, when the latter are lacking and only shallow water forms of brachiopods and trilobites are available, the terms of the neritic Rhenish facies, Early and Late Emsian substages, have to be used. The historical base of the original Early Emsian has been correlated, a century later, with the base of the Ulmen Substage i.e., the Eckfeld Schichten (Fig. 2), which may be slightly younger than the base of the Zlichovian (cf. Weddige, 1996). The base of the Late Emsian, in contrast, is more or less coeval with the base of the Dalejan. With regard to the event stratigraphy, this subdivision of the global Emsian is supported by the Basal Zlichov Event (regressive pulse) and the (basal) Daleje Event (transgressive pulse). Both events are now regarded as more important than the basal Pragian Event, because they are recognized globally. The numerical age for the base of the Lochkovian Stage (base of the Devonian Period) was estimated by Tucker et al. (1998) at 418 Ma using three ID-TIMS ages from the Lochkovian and Pragian stages and the Kalkberg age of 417.6 ± 1.9 Ma from the upper half of the Icriodus woschmidti conodont Zone (earliest Lochkovian; Kaufmann, 2006). Using the same ages as Tucker et al. (1998), and part of their error bars, an age of 417.5 Ma has been derived for the DCP 2003. Kaufmann (2006) favours an age of 418.1 ± 3.0 Ma which may have been influenced at least from the Early Emsian Bundenbach ID-TIMS age of 407.7 ± 0.7 Ma of Kaufmann et al. (2005) (cf. Fig. 1). House and Gradstein (2004: Fig. 14.5; GTS 2004) estimated an age of 416 Ma using the Ludfordian Upper Whitcliffe ID- TIMS age of 420.2 ± 3.9 Ma (Tucker et al., 1998) as tie point, giving the other ages a lower weight. Compston (2000) suggested a SHRIMP age of 409.9 ± 1.1 Ma for the base of the Devonian, after he reprocessed original data, allowing for the heterogeneity in the zircon SL13 standard. His entire Devonian time scale, except for the upper limit, is in strong contrast to the other scales based mainly on ID-TIMS ages (Fig. 1). Since the work of Tucker et al. (1998), the Emsian Stage has been thought to have a duration between 15 Ma and 17 Ma, except for the House and Gradstein (2004: Fig. 14.5; GTS, 2004) scale. The latter combines three isotopic ages and the “ House scale ” which “ juggles ” the high-resolution ammonoid and conodont (sub) zonations, such that shortest segments are fairly close to equal duration ” (House and Gradstein, 2004: 215). This methodology appears to be flawed by the exclusion of some very significant ID-TIMS ages in its calibration procedure. Thus, the age of the upper boundary of the Emsian Stage at 397.5 ± 2.7 Ma may be too old and its duration of 9.5 Ma is probably too short. Most probably, the Emsian Stage is longer than the combined ages of the Lochkovian and Pragian stages (Fig. 1). Two thirds of the German part of the Rhenish Massif is formed by Early Devonian rocks, which consist of mainly siliciclastic sediments in a neritic shallow-water facies ( “ Rhenish Facies “ ). The fauna is dominated by mostly brachiopods and trilobites, whereas (hemi) pelagic global index fossils such as conodonts, goniatites, tentaculites, and graptolites are lacking. Thus, the global stages Lochkovian, Pragian, and Emsian are rarely recognizable, and the regional historical stage subdivision of Gedinne, Siegen, and Ems has to be maintained. The siliciclastic material originated from the NW, from the “ Old Red Continent ” (part of the northern supercontinent Laurussia), and was accumulated in the Rhenish trough. Accumulation rate and crustal subsidence of the shelf area were both high suggesting a constant palaeobathymetry so that the entire Early Devonian succession became up to 10,000 m thick. The strong subsidence initiated crustal movements, and volcanics sporadically intercalated in the sediments, e.g. the Hauptkeratophyr (Fig. 2). Facies shifts can be observed in sedimentary sequences from lacustrine, brackish, limnic-fluviatile to marine shallow-water (Fig. 2: Bunte Ebbe-Schichten and Siesel-Schichten). The stratigraphy of these Early Devonian successions is still under review. It is suspected, however, that the successions are only local developments, and thus do not reflect extra-regional causes like climatic changes. Even at the end of the Early Devonian, the late Emsian strata are generally transgressive, obviously related to a widespread deepening of the shelf area and/or even with a global rise of the sea level and climatic warming (Fig. 2: Wiltz Schichten and Heisdorf Schichten). During the Early Devonian most of the East European (or Russian) Platform was an area of sub- aerial erosion, and only on its southwestern margin (Lvov-Volyn, Ukraine) carbonate marine sedimentation took place (Alekseev et al., 1996). There, the base of the Lochkovian coincides with the lower boundary of the Borshevian Horizon with the FAD of the conodont I. woschmidti together with M. uniformis angustidens (Drygant, 1984). The Late Lochkovian, Pragian, and lower Emsian (Kemerian) are represented by a lagoonal and terrestrial succession in the Lvov-Volyn and Baltic area (Latvia and Lithuania). Their ages are based on evidence from thelodont and acanthodian fish zonation and spore assemblages (Avkhimovich et al., 1993). The first Devonian marine transgression is of late Emsian age. Its carbonate and terrigenous sediments are known mainly on the eastern part of the East European Platform (Volga-Ural, western slope of the Ural Mts.). They contain conodont assemblages of the E. dehiscens, Eocostapolygnathus gronbergi , Eucostapolygnathus inversus – Linguipolygnathus serotinus , and Eucostapolygnathus patulus zones. The ages of the Takata and Vanyshkinsk formations are defined by miospores and brachiopods. The Koivinsk and Biysk formations are composed by shallow water argillites, marls, and limestone (Koivinsk Formation) and essentially shallow water limestones (Biysk Formation). Their ages are defined by conodonts, corals, and brachiopods. Early Devonian deposits of the Zeravshan Ridge (South Tien Shan) are represented by shallow water reef and marly limestones with diverse benthic corals, brachiopods, trilobites, and other faunal groups, as well as remains of pelagic organisms: graptolites, dacryoconarids, and conodonts (Kim et al., 1984, 1988; Yolkin et al., 2000). The base of the Lochkovian coincides with sharp lithological changes at the lower boundary of the Bursykhirmanian and is defined by the appearance of M. uniformis , together with the trilobite W. rugulosa , and the brachiopods Protathyris praecur- sor and Lanceomyonia borealiformis . The Pragian massive limestones of the Khukarian are characterized by benthic Koneprusy-type faunas and conodont assemblages of the E. sulcatus, Pseudogondwania kindlei , and E. pireneae zones. The GSSP for the basal Emsian boundary is defined at the base of the E. kitabicus Zone (Yolkin et al., 1989, 1994). This boundary is aligned with the base of the Kitabian that embraces ...

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... guide fossils and by interfacial correlations. Thus, the accompanying conodont Latericriodus steinachensis appears just below the GSSP boundary. The dacryoconarids Nowakia sororcula and Nowakia acuaria enter shortly above the boundary, and chitinozoans are useful because they can be linked to the spore zonation. Lithologically, the early Pragian rocks become gradually lighter in colour, although their overall facies remains similar to that of the Lochkovian (Chlupá č , 2000). A relatively minor faunal change in the early Pragian marks the rather insignif- icant Lochkovian – Pragian Event. The GSSP for the Emsian Stage is within the Khodzha – Kurgan Formation of the Zinzilban Gorge of the Kitab National Park in Uzbekistan (Yolkin et al., 1997; Yolkin et al., 2000). The boundary is indexed by the conodont Eocostapolygnathus kitabicus which is derived, and well differentiated, from its ancestor Eoctenopolygnathus pireneae , and thus represents a proven biostratigraphic time marker (Yolkin et al., 1994). The level of this phylogenetic transition is lower than the entry of the somewhat controversial taxon Eolinguipolygnathus dehiscens , and the entries of both taxa are lower than the top of the regional historical Pragian Stage. The rarity of conodonts in the late Pragian of the type area, however, does not permit an exact correlation with the Emsian GSSP stratotype in Uzbekistan. When Yolkin et al. (1994) proposed such a low Emsian base at the expense of the historical Pragian Stage, the authors intended to introduce a further stage for the latest Early Devonian. The early interval of the global Emsian corresponds with a part of the late original Pragian Stage in its Czech type region, and with the so-called Zinzilbanian in Uzbekistan, Central Asia ( Fig. 2). For the middle and late intervals of the global Emsian, the Czech terms Zlichovian and Dalejan are currently used informally when open marine to pelagic guide fossils like conodonts, dacryoconarids, trilobites, and goniatites are available. But, when the latter are lacking and only shallow water forms of brachiopods and trilobites are available, the terms of the neritic Rhenish facies, Early and Late Emsian substages, have to be used. The historical base of the original Early Emsian has been correlated, a century later, with the base of the Ulmen Substage i.e., the Eckfeld Schichten (Fig. 2), which may be slightly younger than the base of the Zlichovian (cf. Weddige, 1996). The base of the Late Emsian, in contrast, is more or less coeval with the base of the Dalejan. With regard to the event stratigraphy, this subdivision of the global Emsian is supported by the Basal Zlichov Event (regressive pulse) and the (basal) Daleje Event (transgressive pulse). Both events are now regarded as more important than the basal Pragian Event, because they are recognized globally. The numerical age for the base of the Lochkovian Stage (base of the Devonian Period) was estimated by Tucker et al. (1998) at 418 Ma using three ID-TIMS ages from the Lochkovian and Pragian stages and the Kalkberg age of 417.6 ± 1.9 Ma from the upper half of the Icriodus woschmidti conodont Zone (earliest Lochkovian; Kaufmann, 2006). Using the same ages as Tucker et al. (1998), and part of their error bars, an age of 417.5 Ma has been derived for the DCP 2003. Kaufmann (2006) favours an age of 418.1 ± 3.0 Ma which may have been influenced at least from the Early Emsian Bundenbach ID-TIMS age of 407.7 ± 0.7 Ma of Kaufmann et al. (2005) (cf. Fig. 1). House and Gradstein (2004: Fig. 14.5; GTS 2004) estimated an age of 416 Ma using the Ludfordian Upper Whitcliffe ID- TIMS age of 420.2 ± 3.9 Ma (Tucker et al., 1998) as tie point, giving the other ages a lower weight. Compston (2000) suggested a SHRIMP age of 409.9 ± 1.1 Ma for the base of the Devonian, after he reprocessed original data, allowing for the heterogeneity in the zircon SL13 standard. His entire Devonian time scale, except for the upper limit, is in strong contrast to the other scales based mainly on ID-TIMS ages (Fig. 1). Since the work of Tucker et al. (1998), the Emsian Stage has been thought to have a duration between 15 Ma and 17 Ma, except for the House and Gradstein (2004: Fig. 14.5; GTS, 2004) scale. The latter combines three isotopic ages and the “ House scale ” which “ juggles ” the high-resolution ammonoid and conodont (sub) zonations, such that shortest segments are fairly close to equal duration ” (House and Gradstein, 2004: 215). This methodology appears to be flawed by the exclusion of some very significant ID-TIMS ages in its calibration procedure. Thus, the age of the upper boundary of the Emsian Stage at 397.5 ± 2.7 Ma may be too old and its duration of 9.5 Ma is probably too short. Most probably, the Emsian Stage is longer than the combined ages of the Lochkovian and Pragian stages (Fig. 1). Two thirds of the German part of the Rhenish Massif is formed by Early Devonian rocks, which consist of mainly siliciclastic sediments in a neritic shallow-water facies ( “ Rhenish Facies “ ). The fauna is dominated by mostly brachiopods and trilobites, whereas (hemi) pelagic global index fossils such as conodonts, goniatites, tentaculites, and graptolites are lacking. Thus, the global stages Lochkovian, Pragian, and Emsian are rarely recognizable, and the regional historical stage subdivision of Gedinne, Siegen, and Ems has to be maintained. The siliciclastic material originated from the NW, from the “ Old Red Continent ” (part of the northern supercontinent Laurussia), and was accumulated in the Rhenish trough. Accumulation rate and crustal subsidence of the shelf area were both high suggesting a constant palaeobathymetry so that the entire Early Devonian succession became up to 10,000 m thick. The strong subsidence initiated crustal movements, and volcanics sporadically intercalated in the sediments, e.g. the Hauptkeratophyr (Fig. 2). Facies shifts can be observed in sedimentary sequences from lacustrine, brackish, limnic-fluviatile to marine shallow-water (Fig. 2: Bunte Ebbe-Schichten and Siesel-Schichten). The stratigraphy of these Early Devonian successions is still under review. It is suspected, however, that the successions are only local developments, and thus do not reflect extra-regional causes like climatic changes. Even at the end of the Early Devonian, the late Emsian strata are generally transgressive, obviously related to a widespread deepening of the shelf area and/or even with a global rise of the sea level and climatic warming (Fig. 2: Wiltz Schichten and Heisdorf Schichten). During the Early Devonian most of the East European (or Russian) Platform was an area of sub- aerial erosion, and only on its southwestern margin (Lvov-Volyn, Ukraine) carbonate marine sedimentation took place (Alekseev et al., 1996). There, the base of the Lochkovian coincides with the lower boundary of the Borshevian Horizon with the FAD of the conodont I. woschmidti together with M. uniformis angustidens (Drygant, 1984). The Late Lochkovian, Pragian, and lower Emsian (Kemerian) are represented by a lagoonal and terrestrial succession in the Lvov-Volyn and Baltic area (Latvia and Lithuania). Their ages are based on evidence from thelodont and acanthodian fish zonation and spore assemblages (Avkhimovich et al., 1993). The first Devonian marine transgression is of late Emsian age. Its carbonate and terrigenous sediments are known mainly on the eastern part of the East European Platform (Volga-Ural, western slope of the Ural Mts.). They contain conodont assemblages of the E. dehiscens, Eocostapolygnathus gronbergi , Eucostapolygnathus inversus – Linguipolygnathus serotinus , and Eucostapolygnathus patulus zones. The ages of the Takata and Vanyshkinsk formations are defined by miospores and brachiopods. The Koivinsk and Biysk formations are composed by shallow water argillites, marls, and limestone (Koivinsk Formation) and essentially shallow water limestones (Biysk Formation). Their ages are defined by conodonts, corals, and brachiopods. Early Devonian deposits of the Zeravshan Ridge (South Tien Shan) are represented by shallow water reef and marly limestones with diverse benthic corals, brachiopods, trilobites, and other faunal groups, as well as remains of pelagic organisms: graptolites, dacryoconarids, and conodonts (Kim et al., 1984, 1988; Yolkin et al., 2000). The base of the Lochkovian coincides with sharp lithological changes at the lower boundary of the Bursykhirmanian and is ...

Context 4

... guide fossils and by interfacial correlations. Thus, the accompanying conodont Latericriodus steinachensis appears just below the GSSP boundary. The dacryoconarids Nowakia sororcula and Nowakia acuaria enter shortly above the boundary, and chitinozoans are useful because they can be linked to the spore zonation. Lithologically, the early Pragian rocks become gradually lighter in colour, although their overall facies remains similar to that of the Lochkovian (Chlupá č , 2000). A relatively minor faunal change in the early Pragian marks the rather insignif- icant Lochkovian – Pragian Event. The GSSP for the Emsian Stage is within the Khodzha – Kurgan Formation of the Zinzilban Gorge of the Kitab National Park in Uzbekistan (Yolkin et al., 1997; Yolkin et al., 2000). The boundary is indexed by the conodont Eocostapolygnathus kitabicus which is derived, and well differentiated, from its ancestor Eoctenopolygnathus pireneae , and thus represents a proven biostratigraphic time marker (Yolkin et al., 1994). The level of this phylogenetic transition is lower than the entry of the somewhat controversial taxon Eolinguipolygnathus dehiscens , and the entries of both taxa are lower than the top of the regional historical Pragian Stage. The rarity of conodonts in the late Pragian of the type area, however, does not permit an exact correlation with the Emsian GSSP stratotype in Uzbekistan. When Yolkin et al. (1994) proposed such a low Emsian base at the expense of the historical Pragian Stage, the authors intended to introduce a further stage for the latest Early Devonian. The early interval of the global Emsian corresponds with a part of the late original Pragian Stage in its Czech type region, and with the so-called Zinzilbanian in Uzbekistan, Central Asia ( Fig. 2). For the middle and late intervals of the global Emsian, the Czech terms Zlichovian and Dalejan are currently used informally when open marine to pelagic guide fossils like conodonts, dacryoconarids, trilobites, and goniatites are available. But, when the latter are lacking and only shallow water forms of brachiopods and trilobites are available, the terms of the neritic Rhenish facies, Early and Late Emsian substages, have to be used. The historical base of the original Early Emsian has been correlated, a century later, with the base of the Ulmen Substage i.e., the Eckfeld Schichten (Fig. 2), which may be slightly younger than the base of the Zlichovian (cf. Weddige, 1996). The base of the Late Emsian, in contrast, is more or less coeval with the base of the Dalejan. With regard to the event stratigraphy, this subdivision of the global Emsian is supported by the Basal Zlichov Event (regressive pulse) and the (basal) Daleje Event (transgressive pulse). Both events are now regarded as more important than the basal Pragian Event, because they are recognized globally. The numerical age for the base of the Lochkovian Stage (base of the Devonian Period) was estimated by Tucker et al. (1998) at 418 Ma using three ID-TIMS ages from the Lochkovian and Pragian stages and the Kalkberg age of 417.6 ± 1.9 Ma from the upper half of the Icriodus woschmidti conodont Zone (earliest Lochkovian; Kaufmann, 2006). Using the same ages as Tucker et al. (1998), and part of their error bars, an age of 417.5 Ma has been derived for the DCP 2003. Kaufmann (2006) favours an age of 418.1 ± 3.0 Ma which may have been influenced at least from the Early Emsian Bundenbach ID-TIMS age of 407.7 ± 0.7 Ma of Kaufmann et al. (2005) (cf. Fig. 1). House and Gradstein (2004: Fig. 14.5; GTS 2004) estimated an age of 416 Ma using the Ludfordian Upper Whitcliffe ID- TIMS age of 420.2 ± 3.9 Ma (Tucker et al., 1998) as tie point, giving the other ages a lower weight. Compston (2000) suggested a SHRIMP age of 409.9 ± 1.1 Ma for the base of the Devonian, after he reprocessed original data, allowing for the heterogeneity in the zircon SL13 standard. His entire Devonian time scale, except for the upper limit, is in strong contrast to the other scales based mainly on ID-TIMS ages (Fig. 1). Since the work of Tucker et al. (1998), the Emsian Stage has been thought to have a duration between 15 Ma and 17 Ma, except for the House and Gradstein (2004: Fig. 14.5; GTS, 2004) scale. The latter combines three isotopic ages and the “ House scale ” which “ juggles ” the high-resolution ammonoid and conodont (sub) zonations, such that shortest segments are fairly close to equal duration ” (House and Gradstein, 2004: 215). This methodology appears to be flawed by the exclusion of some very significant ID-TIMS ages in its calibration procedure. Thus, the age of the upper boundary of the Emsian Stage at 397.5 ± 2.7 Ma may be too old and its duration of 9.5 Ma is probably too short. Most probably, the Emsian Stage is longer than the combined ages of the Lochkovian and Pragian stages (Fig. 1). Two thirds of the German part of the Rhenish Massif is formed by Early Devonian rocks, which consist of mainly siliciclastic sediments in a neritic shallow-water facies ( “ Rhenish Facies “ ). The fauna is dominated by mostly brachiopods and trilobites, whereas (hemi) pelagic global index fossils such as conodonts, goniatites, tentaculites, and graptolites are lacking. Thus, the global stages Lochkovian, Pragian, and Emsian are rarely recognizable, and the regional historical stage subdivision of Gedinne, Siegen, and Ems has to be maintained. The siliciclastic material originated from the NW, from the “ Old Red Continent ” (part of the northern supercontinent Laurussia), and was accumulated in the Rhenish trough. Accumulation rate and crustal subsidence of the shelf area were both high suggesting a constant palaeobathymetry so that the entire Early Devonian succession became up to 10,000 m thick. The strong subsidence initiated crustal movements, and volcanics sporadically intercalated in the sediments, e.g. the Hauptkeratophyr (Fig. 2). Facies shifts can be observed in sedimentary sequences from lacustrine, brackish, limnic-fluviatile to marine shallow-water (Fig. 2: Bunte Ebbe-Schichten and Siesel-Schichten). The stratigraphy of these Early Devonian successions is still under review. It is suspected, however, that the successions are only local developments, and thus do not reflect extra-regional causes like climatic changes. Even at the end of the Early Devonian, the late Emsian strata are generally transgressive, obviously related to a widespread deepening of the shelf area and/or even with a global rise of the sea level and climatic warming (Fig. 2: Wiltz Schichten and Heisdorf Schichten). During the Early Devonian most of the East European (or Russian) Platform was an area of sub- aerial erosion, and only on its southwestern margin (Lvov-Volyn, Ukraine) carbonate marine sedimentation took place (Alekseev et al., 1996). There, the base of the Lochkovian coincides with the lower boundary of the Borshevian Horizon with the FAD of the conodont I. woschmidti together with M. uniformis angustidens (Drygant, 1984). The Late Lochkovian, Pragian, and lower Emsian (Kemerian) are represented by a lagoonal and terrestrial succession in the Lvov-Volyn and Baltic area (Latvia and Lithuania). Their ages are based on evidence from thelodont and acanthodian fish zonation and spore assemblages (Avkhimovich et al., 1993). The first Devonian marine transgression is of late Emsian age. Its carbonate and terrigenous sediments are known mainly on the eastern part of the East European Platform (Volga-Ural, western slope of the Ural Mts.). They contain conodont assemblages of the E. dehiscens, Eocostapolygnathus gronbergi , Eucostapolygnathus inversus – Linguipolygnathus serotinus , and Eucostapolygnathus patulus zones. The ages of the Takata and Vanyshkinsk formations are defined by miospores and brachiopods. The Koivinsk and Biysk formations are composed by shallow water argillites, marls, and limestone (Koivinsk Formation) and essentially shallow water limestones (Biysk Formation). Their ages are defined by conodonts, corals, and brachiopods. Early Devonian deposits of the Zeravshan Ridge (South Tien Shan) are represented by shallow water reef and marly limestones with diverse benthic corals, brachiopods, trilobites, and other faunal groups, as well as remains of pelagic organisms: graptolites, dacryoconarids, and conodonts (Kim et al., 1984, 1988; Yolkin et al., 2000). The base of the Lochkovian coincides with sharp lithological changes at the lower boundary of the Bursykhirmanian and is defined by the appearance of M. uniformis , together with the trilobite W. rugulosa , and the brachiopods Protathyris praecur- sor and Lanceomyonia borealiformis . The Pragian massive limestones of the Khukarian are characterized by benthic Koneprusy-type faunas and conodont assemblages of the E. sulcatus, Pseudogondwania kindlei , and E. pireneae zones. The GSSP for the basal Emsian boundary is defined at the base of the E. kitabicus Zone (Yolkin et al., 1989, 1994). This boundary is aligned with the base of the Kitabian that embraces the Emsian E. kitabicus , E. excavatus , serotinus , and E. patulus conodont zones and the Eifelian E. partitus Zone. The Early Devonian of South Fergana is represented by more deep water carbonates and fine clastics that include such index graptolite species as M. uniformis , M . praehercynicus , and M. hercynicus (Kim et al., 1988). The New York State succession is regarded as the classical standard of the North American Devonian, and of the method of sequence stratigraphy (e.g. Brett and Ver Straeten, 1997). The Rondout Formation, which includes the Silurian – Devonian boundary, in eastern New York represents a subtidal nearshore marine environment. It forms the base of the Helderberg Group in which carbonates of intertidal facies dominate. Then, a transgressive period followed with the ...

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... the Filippovian and Irenian horizons), by the inclusion of the former Artinskian Saranian Horizon and the Ufimian, Solikamskian, and Sheshmian horizons. Thus, in contrast to the GTS 2004, the global Kungurian may be the longest global Permian stage (cf. Fig. 4). Numerical ages of 290 to 299 ± 1.0 Ma are given for the base of the Asselian Stage (Fig. 1). An age of 296 Ma (Menning, 1989; ff.) is used in the DCP 2003/ STD 2002, which is derived from ages from Central Europe (Lippolt and Hess, 1983; Lippolt et al., 1984; Lippolt and Hess, 1989; Lippolt et al., 1989). This age range encompasses those previously suggested viz., ca. 292 Ma (Chuvashov et al., 1996, SHRIMP; cf. Menning et ...

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... with the Cherry Canyon Formation (Guadalupe Mts.) as shown by Dunbar et al. (1960), but about 1.5 my earlier (this work, Fig. 4; cf. Dunbar et al., 1960: chart-columns 69, 68). With respect to c): The three GSSPs for the global Middle Permian were ratified in 2001. Their stratotypes are in the Guadalupe Mts. (Delaware Basin) of West Texas (Ogg, 2004; cf. 5.3.1, Fig. 4). The base of the Roadian Stage is defined in the upper part of the original Bone Spring Formation (now the Cutoff Formation). In the Glass Mts. this level is within the lower part of the original Word Formation, now Road Canyon Formation, the former “ First Limestone ” of King (1931), which was included in the original Guadalupian Series by Dunbar et al. (1960). The base of the Wordian Stage is in the Getaway Limestone Member of the new Cherry Canyon Formation, which correlates with parts of the original middle and upper Word Formation of the Glass Mts. The base of the Capitanian Stage is within the Pinery Limestone Member of the Bell Canyon Formation which correlates to the main part of the Capitan Limestone (Dunbar et al., 1960, chart-column: American Standard Section). The name of the Lopingian Series is derived from the Loping Coal-bearing “ series ” in South China, named by von Richthofen in 1884 (Jin et al., 2003). Huang (1932) later established it as a formal series to include all Permian deposits that overlie the Maokou Limestone. The Lopingian is subdivided into the Wuchiapingian and Changhsingian stages. Their names are derived from the Wuchiaping and Changhsing (Changxing) formations (Furnish and Glenister, 1970; Kanmera and Nakazawa, 1973). The Lopingian finishes at the base of the Triassic (Yin et al., 2001), which is defined by the FAD of the conodont Hindeodus parvus . The GSSP for the basal boundary of the Wuchiapingian Stage, and also for the Lopingian Series/Epoch is placed at the FAD of the conodont Clarkina postbitteri postbitteri within an evolutionary lineage from C. postbitteri hongshuiensis to C. dukouensi s in Bed 6k at the Penglaitan Section, Laibin area, Guangxi Province, South China (Jin et al., 2001). The FAD of C. postbitteri sensu lato could also be used to approximate this boundary, as it is only 20 cm below the defining point at the base of Bed 6i upper at the Penglaitan Section. Historically, this boundary was intended to coincide with a global regression and has been documented as a level coincident with an important mass extinction event. This event is marked by a rapid change from Jinogondolella ( Mesogondolella ) to Clarkina for off- shore conodonts, and Sweetognathus to Iranognathus for nearshore shallow water conodonts, and the sudden extinction of verbeekinid fusulinids (Jin et al., 1998a). Wuchiapingian fusulinids are dominated by the fusulinids Codonofusiella and Reichelina . The stage is composed of the Laibinian and Laoshanian substages bounded by the base of the Clarkina leveni Zone. This level is approximately correlated with the base of the Anderssonoceras – Prototoceras ammonoid Zone and coincides with the surface of the Lopingian transgression. The Laibinian Substage contains the C . postbitteri postbitteri , C. dukouensis , and C . asymmetrica conodont zones, and the Laoshanian Substage contains the C . leveni , C. guangyuanensis , and Clarkina orientalis conodont zones as well as the Anderssonoceras – Prototoceras , Araxoceras – Konglingites , and Sanyangites ammonoid zones. The Changhsingian Stage was formally proposed as the last stage of the Palaeozoic Era with its stratotype in the Section D at Meishan, Changxing County, Zhejiang Province of China (Zhao et al., 1981) with a basal boundary between the C. orientalis and Clarkina subcarinata zones. Recently, the boundary GSSP was ratified by the IUGS at the FAD of the conodont Clarkina wangi within the lineage from Clarkina longicuspidata to C. wangi at a point 88 cm above the base of the Changhsing Formation in the lower part of Bed 4 (base of 4a-2). This point is just above the flooding surface of the second parasequence in the Changhsing Limestone (Jin et al., 2004). The basal part of this stage is also marked by the occurrence of advanced forms of Palaeofusulina , and the tapashani- tid and pseudotirolitid ammonoids. The Baoqingian and the Meishanian substages have been suggested, bounded by the FOD of Clarkina changxingensis within Bed 10. The Changhsingian – Indusian boundary GSSP is close to the top of the Changhsing Formation within the lowermost Yinkeng Formation (Fig. 4). The numerical age of 260.5 Ma for the Capitanian – Wuchiapingian boundary in the STD 2002/DCP 2003 is an estimation which is based on a) the 265.3 ± 0.2 Ma ID-TIMS age (Bowring et al., 1998) for a bentonite bed just below the base of the Capitanian Stage in its boundary stratotype at the Nipple Hill near the Guadalupe Mts., West Texas, b) the numerous ID- TIMS ages from the sections Meishan and Shangsi around the PTB (Bowring et al., 1998; cf. Menning, 2001: Fig. 1) as well as the SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan (Claoué- Long et al., 1991) which is only slightly older than the PTB in Bed 27c (Yin et al., 2001), and c) geological time indications (cf. Section 2.1). The latter suggest that the age of 253.4 ± 0.2 Ma from just above the base of the Changhsingian Stage (Bowring et al., 1998) may be little too young. In the DCP 2003 are allocated durations of ∼ 4 my to the Changhsingian, of ∼ 5.5 Ma to the Wuchiapingian, and of ∼ 4.5 my to the Capitanian stages (Figs. 1 and 4). When using the age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of ∼ 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The allocation of ∼ 0.4 my to each of the 25 Lopingian parasequences (17.3 in the Wuchiapingian and 7.7 in the Changhsingian) of Chen et al. (1998: Fig. 9) results in durations of ∼ 10 my for the Lopingian Epoch, of ∼ 6.9 my for the Wuchiapingian Stage and of ∼ 3.1 my for the Changhsingian Stage (Menning et al., 2005; Fig. 1). Thus, the age for the Capitanian – Wuchiapingian boundary is ∼ 261 Ma according to an age of ∼ 251 Ma for the PTB (STD 2002; GTS 2004) or ∼ 262.5 Ma according to an age of ∼ 252.5 Ma for the PTB (Menning et al., 2005). The corresponding ages for the base of the Changhsingian Stage are 254.1 Ma and 255.6 Ma respectively, whereas in the STD 2002/DCP 2003 an age of ∼ 255 Ma is suggested. According to the orbital interpretation of the 25 Lopingian parasequences of Chen et al. (1998: Fig. 9) the duration of the Wuchiapingian is slightly underestimated and that of the Changhsingian is slightly overestimated. The ages of ∼ 261 Ma and ∼ 262.5 Ma for the base of the Wuchiapingian Stage (Menning et al., 2005) are consistent with a zircon age of 259.3 ± 3 Ma from the Emeishan volcanics occurring around the Guadalupian – Lopingian boundary (Zhou et al., 2002), and approximates with the age of 260.8 ± 0.8 Ma from the Shangsi Section (Mundil et al., 2004). The latter age is not Early Wuchiapingian as suggested, but Late Wuchiapingian because Bed 7 contains conodonts of the C. orientalis Zone (Jin, this work). The age of 257.3 ± 0.3 Ma (Mundil et al., 2004) for the Early Changhsingian is significantly older than the orbital age of 255.6 Ma derived by Menning et al. (2005), which is based on the 252.6 ± 0.2 Ma age for bed 25 in Meishan (Mundil et al., 2004). In Central and West Europe the stratigraphic range of the succession equivalent in age to the Lopingian Epoch (STD 2002/DCP 2003: ∼ 260.5 Ma to 251 Ma) is under discussion. Whereas sediments of that age are rarely developed in the intra-Variscan basins, a complete succession is represented in the foreland basin between Central England and East Poland. There is no evidence of a break in time of ≥ 0.1 my in the migrating basin centre in North Germany during the 266 Ma to 229 Ma interval (STD 2002; Menning et al., 2005), but Wardlaw et al. (2004: Fig. 16.2) show a gap between the top of the Zechstein at 259.8 Ma and the base of the Buntsandstein at 251.0 Ma. In the Central European Basin late Rotliegend sediments consists of fluvial, aeolian, playa, and sabkha sediments, which are up to 2000 m thick (Deutsche Stratigraphische Kommission, 1995; Schröder et al., 1995). These sediments are covered by Zechstein and younger sediments (STD 2002), which are some kilometres thick and penetrated by more than 1000 wells. The mainly brownish red clastics of the Havel Subgroup include aeolian sandstones with significant resources of natural gas. The clastic – evaporitic Elbe Subgroup covers a much larger area and is better subdivided and locally correlated than the Havel Subgroup. In NW Germany it accumulated mainly in a large perennial salt lake (Gast, 1991; Legler et al., 2005) and its orbitally estimated duration is ca. 5.6 my (cf. Section 5.3.2; STD 2002/DCP 2003: ∼ 4 Ma). This duration is consistent with the age for the Rotliegend – Zechstein boundary at ∼ 258 Ma in the STD 2002/DCP 2003, as well as the Re – Os age of 257.3 ± 1.7 Ma of the Kupferschiefer (Copper Shale) at the base of the Zechstein (Brauns et al., 2003). The marine – lagoonal – terrestrial Zechstein, which is subdivided and correlated in detail using marker horizons (Richter-Bernburg, 1955), consists of fine grained clastics, carbonates, sulfates, and chlorides in the central basin and terrestrial conglomerates, sandstones, and claystones in marginal areas. As in the late Rotliegend, the cyclic sedimentation is related to climatic variations. According to magnetostratigraphic data (Menning et al., 1988), correlations (Menning, 1986, 2001) and geological time indications, the uppermost part of the Rotliegend Group and the lower part of the Zechstein Group belong to the Wuchiapingian Stage, whereas the main part of the Zechstein Group belongs to the Changhsingian Stage (Fig. 4; STD 2002). The base of the Wuchiapingian Stage is within the upper Rotliegend Elbe Subgroup (Fig. 4; STD ...

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... et al., 1969) for defining the base of the Visean was inappropriate. The SCCS Working Group on the Tournaisian/Visean boundary followed the proposal of Hance and Muchez (1995) to search for a criterion within the evolutionary lineage of the genus Eoparastaffella . Work by its members, summarized in Devuyst et al. (2003) led to the proposal that the first appearance of Eoparastaffella simplex in the lineage E . ovalis → E . simplex should be used in the choice of the base of the Visean. This proposal was formally approved by the Voting Members of the SCCS in 2002 (Work, 2002) (cf. Fig. 3). The proposed new GSSP for the base of the Visean at Pengchong (Guangxi) southern China (Devuyst et al., 2003) is in the process of ratification. The section exposes a unit of allochthonous mass-flow deposits (Pengchong Member, Luzhai Formation) which accumulated in a starved, intra-platform basin. The evolution of Eoparastaffella is exceptionally well documented. The base of the Visean is defined at the base of bed 85, in which E. simplex first appears. The conodont Gnathodus homopunctatus , a useful, but cryptogenic, guide for the base of the Visean, enters in bed 86, about 1 m above the boundary (see Section 4.2.5). This new Tournaisian – Visean boundary is significantly younger than some historical (pre-1969) interpretations. Certain commonly used index macrofossils, formerly regarded as earliest Visean, on microfossil evidence, now indicate latest Tournaisian, e.g. the productid brachiopod Levitusia humerosa in the shallow water facies of Western Europe was thought to indicate earliest Visean, in contrast to the long controversial, but now accepted, Russian dating of this biozone as late Late Tournaisian in the Kosvian Horizon of the Ural Mts. Similarly, microfossil evidence indicates the Ammonellipsites / Merocanites ammonoid community in Europe, North Africa, Middle Asia, and North America should be removed from the basal Visean and restricted to an expanded Late Ivorian. The Serpukhovian was proposed in 1890 and officially included into the stratigraphic scale of the USSR in 1971 as an equivalent of the Early Namurian. Its type area is at the south limb of the Moscow Basin about 100 km south of Moscow, in the vicinity of Serpukhov. Originally it was defined on the basis of brachiopod assemblages and later subdivided into several foraminiferal zones. The conodont and foraminiferal assemblages of the early Serpukhovian are very similar to those of the late Visean (Gibshman, 2001; Nikolaeva et al., 2002). The late Serpukhovian was redefined mainly using the more complete succession of the Donets Basin (late Arnsbergian Zapaltyubian Horizon). The boundary between the Visean and Serpukhovian stages has been traditionally recognized on the basis of goniatites, and placed at the base of the Cravenoceras leion Zone. Korn (1996) proposed Edmooroceras pseudocoronula as a better alternative, because it is geographically less restricted. However, although Mississippian goniatites provide high-resolution regional subdivisions, some endemic taxa are of limited use for effective interregional and intercontinental correlation. Conodonts have also been proposed for the identification of the Visean – Serpukhovian boundary in Europe (Skompski et al., 1995). A SCCS Task Group led by B. Richards is currently working on finding a suitable criterion and GSSP for the base of the Serpukhovian. The lineage Lochriea nodosa – Lochriea ziegleri , which is recorded and controlled by ammonoids in the Rhenish Slate Mts., Germany, has been selected for this purpose. The FAD of L. ziegleri (Nemirovskaya et al., 1994) has been proposed to define the Visean – Serpukhovian boundary, because of its wide geographic distribution (Nemyrovska, 2005). The species is used to recognize the Visean upper boundary in China (Wang and Qi, 2003) and the Ural Mts. (Nikolaeva et al., 2002). On the other hand Gibshman and Baranova (in press) propose the use of the foraminifer “ Millerella ” tortula in the lineage “ Endostaffella ” tortula (part) –“ M. ” tortula (part), and of Janischewskina delicata in the lineage J. typica – J. delicata . According to the authors, the recognition of “ M. ” tortula at the base of the Serpukhovian in the Moscow Region and the Pericaspian Basin offers the possibility of a correlation with the Middle Chesterian of North America. The numerical age for the base of the Tournaisian Stage (base of the Carboniferous Period) is between 354 Ma (Young and Laurie, 1996) and 362 Ma (Tucker et al., 1998) (Fig. 1). The 354 Ma age was derived from an Early Carboniferous SHRIMP age (Claoué-Long et al., 1993), which is now known to be too young because of the inherent problem in the SL13 zircon standard used for the determination. Nevertheless, the 354 Ma age was applied as a tie point in the Phanerozoic time scales of Young and Laurie (1996) and Gradstein and Ogg (1996), and in the Carboniferous time scales of Jones (1995) and Menning et al. (2000) (Fig. 1), not least, because the dated sample was collected from the S. sulcata conodont Zone in the well known Hasselbach section in the Rhenish Slate Mts. The older age of 362 Ma was derived mainly from the average U – Pb ID-TIMS age of 363.6 ± 1.6 Ma from the latest Famennian Piskahegan Group of New Brunswick, Canada (Tucker et al., 1998) and, less significantly, from a Rb – Sr isochron age of 361.0 ± 4.1 Ma for the S. sulcata Zone in Nanbiancun, South China (Yang et al., 1988). Because of the 8 my discrepancy between the 354 Ma and 362 Ma ages, a compromise age of 358 Ma is used for the STD 2002 and the DCP 2003. Recently, Trapp et al. (2004) derived an age of 360.7 Ma for the DCB from two tuffs from the Hasselbach section, Rhenish Slate Mts. Their age of 360.5 ± 0.8 Ma and the age of 353.7 ± 4.2 Ma (Claoué-Long et al., 1993; renormalized by Claoué-Long et al., 1995) are both from the same tuff in bed 79. Using the ages of 360.5 ± 0.8 Ma and 360.2 ± 0.7 Ma of Trapp et al. (2004) and the thicknesses of the Hasselbach section instead of the thicknesses of the Lali section (South China, Trapp et al., 2004), the age of the DCB is estimated at 361.4 Ma (Menning, this work, Fig. 1). A Re – Os age of 361.3 ± 2.4 Ma of a black shale from the Exshaw Formation, Canada (Selby and Creaser, 2005) is very close to it. The duration of the Tournaisian Stage in many charts is shown to vary between 10 my and 14 my (Table 2). However, U – Pb ID-TIMS ages from the Rhenish Slate Mts., which yielded an age of ca. 361 Ma for its base (Trapp et al., 2004) and ca. 342 Ma for its top (Trapp, pers. com.), suggest a longer duration of slightly less than 20 my. Thus, these ages suggest the Tournaisian is approximately as long as the Visean Stage, which to date, is, together with the Norian Stage, the longest of the Phanerozoic Eon (STD 2002; Menning et al., 2005). The age for the base of the Serpukhovian Stage has been changed from 333 Ma (Harland et al., 1990) to 327 – 325 Ma, but there are no reliable isotopic ages about this boundary. The DCP age of 326.5 Ma is based on the Pendleian/Arnsbergian coal-tonstein 40 Ar/ 39 Ar age of 324.6 Ma (Lippolt et al., 1984). If preference is given to the middle Arnsbergian age of 319.5 Ma (Lippolt et al., 1984) in lieu of the Pendleian/ Arnsbergian age of 324.6 Ma, an age of 325 Ma or even less, is more probable for the base of the Namurian/ Serpukhovian (cf. Menning et al., 2000: Fig. 7, Scale B). The DCB S. sulcata criterion is not applicable in the shallow carbonate platform facies of the Franco-Belgian Basin. Here, the DCB is traditionally placed at the top of the basal (1 m thick) bed ( “ Tn 1b α ” ) of the Hastière Formation, a coarse grained bioclastic grainstone with lithoclasts and ooids and a reworked Devonian fauna. This bed is included between “ known Devonian ” and “ known Carboniferous ” (Van Steenwinkel, 1992). The Early Tournaisian (mid-Hastarian) of Western Europe includes a remarkable anoxic black shale, just above the basal Siphonodella crenulata Zone. It is seen both in basinal facies (Lower Alum Shale and Ru β schiefer of Germany) and in shallow water facies (Schistes du Pont d'Arcole of Belgium, Malevkian on the East European Platform). Cephalopod limestones ( Gattendorfia Genozone) and ostracod shales (with “ finger print ” entomozoids) below this semiglobal event are reminiscent of the Frasnian/Famennian dysphotic/ aphotic facies development. The historical stratotype for the base of the Visean is located in Belgium in the Bastion section. The boundary lies within the uppermost part of the Leffe Formation which contains thin grainstone layers derived from shallower areas (Lees, 1997). The first Eoparastaffella sp. (a broken specimen unidentifiable at the level of species) is found less than 1 m below the entry of the conodont G. homopunctatus (Conil et al., 1989). The utilization of the original British stages of the Visean (George et al., 1976) has been greatly enhanced by the critical and detailed review of Riley (1993). In the British Chadian stratotype section (Chatburn, Craven Basin, NW England), Riley (1990, 1995) showed that the genus Eoparastaffella enters about 300 m higher than originally reported (George et al., 1976) and precedes the incoming of G. homopunctatus . Riley (1990) therefore introduced the term ‘ late Chadian ’ for that part of the Chadian recognized by Eoparastaffella and accessory taxa, such as G. homopunctatus . Revisions of the widely used Brigantian, Pendleian and Arnsbergian ammonoid zonations of western Europe (Korn, 1996; Korn and Horn, 1997) demonstrate a greater degree of provinciality than previously estimated, when comparisons are made with Mediterranean and Russian ammonoid faunas. The Tournaisian and Visean and the Belgian substages, Hastarian to Warnantian, are revised by Dejonghe (2006). On the East European Platform the Early Carboniferous deposits are composed of mainly marine carbonates with minor terrestrial intervals. ...

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... 1963; Ross, 1979). The base of the Atokan, operationally defined with the FADs of the fusulinids Eoschubertella and Pseudostaffella (Sutherland and Manger, 1984; Groves, 1986) is significantly older than the base of the Moscovian proposed above. Groves et al. (1999) have suggested that the base of the Atokan lies within the Late Bashkirian, either within the late Tashastian, or at the base of the Asatauian Horizon (South Ural Mts.). The Bashkirian – Moscovian boundary coincides with either the base of the Winslowoceras – Diaboloceras ammonoid Genozone of the East European Platform and Ural Mts. (Ruzhentsev and Bogoslovskaya, 1978), or with the Eowellerites Genozone in western Europe and North America (Ramsbottom and Saunders, 1984) which may be slightly younger (Popov, 1979). The Kasimovian Stage is designated as the third stage of the Pennsylvanian Epoch. This interval was separated from the upper part of the Moscovian in 1926 as the “ Teguliferina Horizon ” , a name changed later to Kasimovian Horizon (Dan'shin, 1947) and finally accepted as Kasimovian Stage (Teodorovich, 1949). Its recently nominated type area is in lower reaches of the Moskva River, about 80 km southeast of Moscow, but its name comes from the town of Kasimov in the Ryazan Region. For additional regional information see 4.3.3. The lower boundary was first defined on brachiopod evidence, and in 1962 it was selected at the base of the Protriticites pseudomontiparus – Obsoletes obsoletus foraminiferal Zone. In order to define the base of the Kasimovian in the DCP 2003, and to suggest index fossils for global correlation, the FODs of P. pseudomontiparus, O. obsoletus and the conodont Swadelina subexcelsa (Goreva and Alekseev, 2001) are chosen for the DCP 2003. The FAD of the above- mentioned fusulinids may be in a time span of ≤ 0.5 my. However, in recent years both fusulinid and conodont workers have suggested selecting a higher level for the base of the Kasimovian, at the FODs of the fusulinid Montiparus and of the conodonts Idiognathodus sagittalis (Alekseev and Goreva, 2002; Villa and Task Group, 2004). Because their use would require a redefinition of both the Moscovian and Kasimovian stages, Davydov (2002) proposed the selection of an index species within the evolutionary chronocline of Protriticites fusulinids. This fauna is widely distributed throughout the Tethyan and Boreal realms and has also been found in the Great Basin within middle to late Desmoinesian successions (Wahlman et al., 1997; Davydov et al., 1999). The Gzhelian Stage is designated as the fourth stage of the Pennsylvanian Epoch. It was introduced by Nikitin (1890) defined by a specific brachiopod assemblage, and with a stratotype about 60 km east of Moscow at the village Gzhel. For additional regional information see 4.3.3. The task of defining the base of the Gzhelian was accepted by the Moscovian/Kasimo- vian Boundary Task Group in 2002. From fusulinid data (Rauzer-Chernousova, 1941) its lower boundary coincides with the base of the Rauserites rossicus – Rauserites stuckenbergi Zone. Also the FAD of the fusulinids Daixina , Jigulites , and Rugosofusulina has been proposed as an operational index for the lower boundary of the Gzhelian Stage (Rozovskaya, 1975; Rauzer-Chernousova and Shchegolev, 1979; Davydov, 1990). These genera, however, do not occur in North America (except for the Canadian Arctic). The FOD of the conodont Idiognathodus simulator s. s. is an useful marker because it is easily traced in Europe, Asia and North America (Boardman and Work, 2004; cf. 4.3.6). The Shumardites – Vidrioceras ammonoid Genozone has been conventionally placed at the base of the Gzhelian Stage (Bogoslovskaya et al., 1999). In the southern Ural Mts., advanced species of Shumardites occur in the late Gzhelian (Popov et al., 1985; Davydov, 2001). In order to define its base for the DCP 2003, and to suggest index fossils for global correlation, the conodont I. simulator s.s. and the foraminifer R. rossicus are chosen in this paper. The FAD of both species may be in a time span of <0.5 my. The numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U – Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A (Riley et al., 1993; cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member (Lippolt et al., 1984). This is allocated to approximately the Pendleian – Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). When the 40 Ar/ 39 Ar age of 319.5 Ma ± 7.8 Ma of the Poruba Member (Lippolt et al., 1984) is used instead of the Jaklovec age, an age of 318 Ma or even younger is more plausible. Such an age of 318.1 ± 1.3 Ma has been used by Davydov et al. (2004: Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. (2000). Ages of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. 1). They are based on all the 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian age of Central Europe (Lippolt et al., 1984). In the DCP 2003/STD 2002 an age of 312 Ma is used. An age of younger than 310.5 Ma can be excluded in the case of the base of the Moscovian being within the Westphalian B (Duckmantian Substage). Ages of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe (Lippolt et al., 1984; Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma (Davydov et al. in the GTS, 2004) is an approach to make the largest possible timespan for the Kasimovian and Gzhelian stages, after the maximum age of the top of the Gzhelian was regarded as 299 Ma (Ramezani et al., 2003, ID-TIMS; cf. Section 5.2.1). Ages of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. (marine, GSSP area) and Central/West Europe (continental, the current source of most of the isotopic age data). In the DCP 2003/STD 2002 an age of ca. 302 Ma is used. It is based on the ages of ca. 305 Ma and ca. 296 Ma for the bases of the Kasimovian and Asselian stages (Menning, 1989; Menning et al., 2000; Menning, 2001b) using Davydov's (1992) data that from geological indications, the Gzhelian should be more than twice duration of the Kasimovian. The classical Late Carboniferous successions of Central and Northwest Europe are mainly paralic, with an upward increase of terrigenous material culminating in exclusively continental beds. The stratigraphy is based on very detailed lithostratigraphy in coal basins (with marker horizons as marine bands, coal seams, volcanic ash layers), on macrofloral and miospore zonations, and only to a lesser degree on marine and non-marine faunal zones. Shallow water carbonate faunas (foraminiferids) are totally absent. The most suitable ammonoid faunas are restricted to the early Bashkirian, where they have an extremely fine biozonal resolution from Chokierian to Marsdenian substages (Bisat and Hudson, 1943; Ramsbottom et al., 1979; Ramsbottom and Saunders, 1984). The last Carboniferous conodont assemblage occurs in the early Moscovian (at the base of Bolsovian; =Aegir [Mansfield] marine band). It provides an inter-regional correlation across Europe. Thus, there is an approximate correlation between the base of the Vereian (earliest Moscovian) and the base of the Westphalian C (Bolsovian) (Nemyrovska, herein). However, in Fig. 3 the lower boundary of the Vereian is shown within the upper Westphalian B (Duckmantian), as based on Menning et al. (2001b). Therefore, a downward pointing arrow is shown at the base of the Westphalian C (Fig. 3). Above this level, the macroflora (Remy and Remy, 1977; Kerp, 1988; Josten, 1991), palynomorphs (Clayton et al., 1977; Peppers, 1996), and terrestrial faunas (Haubold, 1970, 1973; Holub and Kozur, 1981; Schneider, 1982; Martens, 1983a,b, 1984; Schneider, 1985, 1996; Boy, 1987; Werneburg, 1989, 2001; Boy and Martens, 1991; Schneider and Werneburg, 1993; Hampe, 1994; Voigt, 2005) are useful for local and regional correlation, but these fossil groups rarely allow detailed intercontinental correlation. However, the carbonate successions of the Cantabrian Mts. (Spain) are correlated by fusulinids, conodonts, and ammonoids with those of eastern Europe (cf. Villa and Task Group, 2004). The Moscovian – Kasimovian boundary is within the Cantabrian (Fig. 3, arrow down at the base of the Cantabrian). In the Subvariscan Foredeep, between Upper Silesia (South Poland) and South Wales (UK), the average cumulative thickness is > 5000 m for rocks of Namurian and Westphalian age, compared with > 3000 m for those of Bashkirian and Moscovian age (Drozdzewski, in Menning et al., 2000: Table 3). In Central Europe, coal seams are concentrated in rocks of Namurian C (Yeadonian Substage) to Westphalian C (Bolsovian Substage) age, e.g. the Ruhr district (Fig. 3), but they start also in the Namurian A (latest Mississippian) in Upper Silesia, and finish also in the early Rotliegend (Gzhelian/Asselian) in the intra-Variscan Saar-Nahe Basin and Saale Basin (STD 2002). In several depocenters, paralic sediments accumulated continuously with minimal gaps, of less than 0.5 my. In the Saar-Nahe Basin, however, with lacustrine, deltaic, and fluvial sediments, some 4 km thick, an “ Asturian ” gap of unknown, but in maximum 2.5 my, ...

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... the boundary (see Section 4.2.5). This new Tournaisian – Visean boundary is significantly younger than some historical (pre-1969) interpretations. Certain commonly used index macrofossils, formerly regarded as earliest Visean, on microfossil evidence, now indicate latest Tournaisian, e.g. the productid brachiopod Levitusia humerosa in the shallow water facies of Western Europe was thought to indicate earliest Visean, in contrast to the long controversial, but now accepted, Russian dating of this biozone as late Late Tournaisian in the Kosvian Horizon of the Ural Mts. Similarly, microfossil evidence indicates the Ammonellipsites / Merocanites ammonoid community in Europe, North Africa, Middle Asia, and North America should be removed from the basal Visean and restricted to an expanded Late Ivorian. The Serpukhovian was proposed in 1890 and officially included into the stratigraphic scale of the USSR in 1971 as an equivalent of the Early Namurian. Its type area is at the south limb of the Moscow Basin about 100 km south of Moscow, in the vicinity of Serpukhov. Originally it was defined on the basis of brachiopod assemblages and later subdivided into several foraminiferal zones. The conodont and foraminiferal assemblages of the early Serpukhovian are very similar to those of the late Visean (Gibshman, 2001; Nikolaeva et al., 2002). The late Serpukhovian was redefined mainly using the more complete succession of the Donets Basin (late Arnsbergian Zapaltyubian Horizon). The boundary between the Visean and Serpukhovian stages has been traditionally recognized on the basis of goniatites, and placed at the base of the Cravenoceras leion Zone. Korn (1996) proposed Edmooroceras pseudocoronula as a better alternative, because it is geographically less restricted. However, although Mississippian goniatites provide high-resolution regional subdivisions, some endemic taxa are of limited use for effective interregional and intercontinental correlation. Conodonts have also been proposed for the identification of the Visean – Serpukhovian boundary in Europe (Skompski et al., 1995). A SCCS Task Group led by B. Richards is currently working on finding a suitable criterion and GSSP for the base of the Serpukhovian. The lineage Lochriea nodosa – Lochriea ziegleri , which is recorded and controlled by ammonoids in the Rhenish Slate Mts., Germany, has been selected for this purpose. The FAD of L. ziegleri (Nemirovskaya et al., 1994) has been proposed to define the Visean – Serpukhovian boundary, because of its wide geographic distribution (Nemyrovska, 2005). The species is used to recognize the Visean upper boundary in China (Wang and Qi, 2003) and the Ural Mts. (Nikolaeva et al., 2002). On the other hand Gibshman and Baranova (in press) propose the use of the foraminifer “ Millerella ” tortula in the lineage “ Endostaffella ” tortula (part) –“ M. ” tortula (part), and of Janischewskina delicata in the lineage J. typica – J. delicata . According to the authors, the recognition of “ M. ” tortula at the base of the Serpukhovian in the Moscow Region and the Pericaspian Basin offers the possibility of a correlation with the Middle Chesterian of North America. The numerical age for the base of the Tournaisian Stage (base of the Carboniferous Period) is between 354 Ma (Young and Laurie, 1996) and 362 Ma (Tucker et al., 1998) (Fig. 1). The 354 Ma age was derived from an Early Carboniferous SHRIMP age (Claoué-Long et al., 1993), which is now known to be too young because of the inherent problem in the SL13 zircon standard used for the determination. Nevertheless, the 354 Ma age was applied as a tie point in the Phanerozoic time scales of Young and Laurie (1996) and Gradstein and Ogg (1996), and in the Carboniferous time scales of Jones (1995) and Menning et al. (2000) (Fig. 1), not least, because the dated sample was collected from the S. sulcata conodont Zone in the well known Hasselbach section in the Rhenish Slate Mts. The older age of 362 Ma was derived mainly from the average U – Pb ID-TIMS age of 363.6 ± 1.6 Ma from the latest Famennian Piskahegan Group of New Brunswick, Canada (Tucker et al., 1998) and, less significantly, from a Rb – Sr isochron age of 361.0 ± 4.1 Ma for the S. sulcata Zone in Nanbiancun, South China (Yang et al., 1988). Because of the 8 my discrepancy between the 354 Ma and 362 Ma ages, a compromise age of 358 Ma is used for the STD 2002 and the DCP 2003. Recently, Trapp et al. (2004) derived an age of 360.7 Ma for the DCB from two tuffs from the Hasselbach section, Rhenish Slate Mts. Their age of 360.5 ± 0.8 Ma and the age of 353.7 ± 4.2 Ma (Claoué-Long et al., 1993; renormalized by Claoué-Long et al., 1995) are both from the same tuff in bed 79. Using the ages of 360.5 ± 0.8 Ma and 360.2 ± 0.7 Ma of Trapp et al. (2004) and the thicknesses of the Hasselbach section instead of the thicknesses of the Lali section (South China, Trapp et al., 2004), the age of the DCB is estimated at 361.4 Ma (Menning, this work, Fig. 1). A Re – Os age of 361.3 ± 2.4 Ma of a black shale from the Exshaw Formation, Canada (Selby and Creaser, 2005) is very close to it. The duration of the Tournaisian Stage in many charts is shown to vary between 10 my and 14 my (Table 2). However, U – Pb ID-TIMS ages from the Rhenish Slate Mts., which yielded an age of ca. 361 Ma for its base (Trapp et al., 2004) and ca. 342 Ma for its top (Trapp, pers. com.), suggest a longer duration of slightly less than 20 my. Thus, these ages suggest the Tournaisian is approximately as long as the Visean Stage, which to date, is, together with the Norian Stage, the longest of the Phanerozoic Eon (STD 2002; Menning et al., 2005). The age for the base of the Serpukhovian Stage has been changed from 333 Ma (Harland et al., 1990) to 327 – 325 Ma, but there are no reliable isotopic ages about this boundary. The DCP age of 326.5 Ma is based on the Pendleian/Arnsbergian coal-tonstein 40 Ar/ 39 Ar age of 324.6 Ma (Lippolt et al., 1984). If preference is given to the middle Arnsbergian age of 319.5 Ma (Lippolt et al., 1984) in lieu of the Pendleian/ Arnsbergian age of 324.6 Ma, an age of 325 Ma or even less, is more probable for the base of the Namurian/ Serpukhovian (cf. Menning et al., 2000: Fig. 7, Scale B). The DCB S. sulcata criterion is not applicable in the shallow carbonate platform facies of the Franco-Belgian Basin. Here, the DCB is traditionally placed at the top of the basal (1 m thick) bed ( “ Tn 1b α ” ) of the Hastière Formation, a coarse grained bioclastic grainstone with lithoclasts and ooids and a reworked Devonian fauna. This bed is included between “ known Devonian ” and “ known Carboniferous ” (Van Steenwinkel, 1992). The Early Tournaisian (mid-Hastarian) of Western Europe includes a remarkable anoxic black shale, just above the basal Siphonodella crenulata Zone. It is seen both in basinal facies (Lower Alum Shale and Ru β schiefer of Germany) and in shallow water facies (Schistes du Pont d'Arcole of Belgium, Malevkian on the East European Platform). Cephalopod limestones ( Gattendorfia Genozone) and ostracod shales (with “ finger print ” entomozoids) below this semiglobal event are reminiscent of the Frasnian/Famennian dysphotic/ aphotic facies development. The historical stratotype for the base of the Visean is located in Belgium in the Bastion section. The boundary lies within the uppermost part of the Leffe Formation which contains thin grainstone layers derived from shallower areas (Lees, 1997). The first Eoparastaffella sp. (a broken specimen unidentifiable at the level of species) is found less than 1 m below the entry of the conodont G. homopunctatus (Conil et al., 1989). The utilization of the original British stages of the Visean (George et al., 1976) has been greatly enhanced by the critical and detailed review of Riley (1993). In the British Chadian stratotype section (Chatburn, Craven Basin, NW England), Riley (1990, 1995) showed that the genus Eoparastaffella enters about 300 m higher than originally reported (George et al., 1976) and precedes the incoming of G. homopunctatus . Riley (1990) therefore introduced the term ‘ late Chadian ’ for that part of the Chadian recognized by Eoparastaffella and accessory taxa, such as G. homopunctatus . Revisions of the widely used Brigantian, Pendleian and Arnsbergian ammonoid zonations of western Europe (Korn, 1996; Korn and Horn, 1997) demonstrate a greater degree of provinciality than previously estimated, when comparisons are made with Mediterranean and Russian ammonoid faunas. The Tournaisian and Visean and the Belgian substages, Hastarian to Warnantian, are revised by Dejonghe (2006). On the East European Platform the Early Carboniferous deposits are composed of mainly marine carbonates with minor terrestrial intervals. Rocks of the basal Tournaisian Stage are absent in the central part of the platform, and the marker conodont species S. sulcata occurs only in sections of the western Ural Mts. Before 1986 (approved officially in Kagarmanov, 1998) the uppermost Famennian was considered as earliest Tournaisian, comparable with the Etroeungtian ( “ Tournai 1a ” ) in Central/West Europe, which now also belongs to the Devonian. The Tournaisian includes 6 regional substages Malevkian ( Siphonodella duplicata ), Upian ( S. kononovae ), Karakubian, Cherepetian ( S. quadruplicata ), Kizelian ( Siphonodella isosticha , Dollymae hassi ), and Kosvian. However, before 2002 the Kosvian substage (late Gnathodus typicus and Scaliognathus anchoralis zones) was regarded as the lowest unit of the Visean (Fig. 3). Whereas the coeval foraminiferal zones correspond approximately to each conodont zone, the brachiopod zones have a coarser stratigraphic resolution. The foraminiferal fauna from the Kosvian correlates well with the latest Tournaisian of Belgium (Brenckle, 1997). The late Tournaisian sediments of the Kizelian and Kosvian regional substages are absent in the central part of the platform, and are ...

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... River Ufa (Chuvashov et al., 2002c). The Artinskian is characterized by clastic-rich deposits. Ammonoids, best represented in the Pre-Uralian Foredeep of the southern Ural Mts. (Ruzhentsev, 1954), and fusulinids, best known from the eastern margin of the East European Platform (Rauzer-Chernousova, 1949), are provincial and, therefore, conodonts are the favoured GSSP fossils. The Sakmarian – Artinskian boundary deposits are well represented in the Dalny Tyulkas (= Tulkus) section which is suggested as stratotype (GSSP candidate) for the base of the Artinskian. A potential GSSP is located within a few meters of the FAD of Sweetognathus whitei , which is sometimes associated with the conodont Mesogondolella bisselli . Several levels yielded radiolarians of the Enactinosphaera crassicalthrata – Quin- queremis arundinea Zone (Chuvashov et al., 2002c). According to Kozur (this work), the Trilonche (not Enactinosphaera ) crassicalthrata – Q. arundinea Zone is a junior synonym of the Patrickella cathaphracta Zone (Kozur and Mostler, 1989). The definitive chronomorphocline of Sweetognathus binodosus to S. whitei also can be recognized in the lower Great Bear Cape Formation on southwestern Ellesmere Island, in the Sverdrup Basin, Canadian Arctic (Henderson, 1988; Beauchamp and Henderson, 1994; Mei et al., 2002), and in the Schroyer to Florence limestones of the Chase Group in Kansas (Wardlaw et al., in press) (Fig. 4). Chuvashov and Chernykh (2000) proposed to define the base of the Kungurian Stage at the FAD of Neostreptognathodus pnevi within a chronomorphocline from advanced Neostreptognathodus pequopensis . This was later accepted (Resolutions, 2006). The proposed stratotype section (Chuvashov et al., 1993, 2002c) is exposed along at the Yuryuzan River downstream of the Mechetlino settlement. Beds 13 – 17 of the Sargian Horizon Gabdrashitovo Formation (1 – 18) include typical Artinskian ammonoids, fusulinids, and conodonts like Neostreptognathodus kamajensis , N. pequopensis , and N . aff. ruzhencevi . The basal part of bed 19 (earliest Saranian Horizon) yield Neostreptognathodus clinei (Kozur, this work: pseudoclinei , is partly assigned to the Roadian N. clinei ), N. pnevi , N. kamajensis , N. pequopensis , and Stepanovites sp. The defining chronomorphocline can also be recognized in the upper Great Bear Cape Formation and upper Trappers Cove Formation on southwestern Ellesmere Island, Sverdrup Basin, Canadian Arctic (Henderson, 1988; Beauchamp and Henderson, 1994; Mei et al., 2002). The proposed global Kungurian is significantly extended, in relation to the supraregional Kungurian (restricted on the Filippovian and Irenian horizons), by the inclusion of the former Artinskian Saranian Horizon and the Ufimian, Solikamskian, and Sheshmian horizons. Thus, in contrast to the GTS 2004, the global Kungurian may be the longest global Permian stage (cf. Fig. 4). Numerical ages of 290 to 299 ± 1.0 Ma are given for the base of the Asselian Stage (Fig. 1). An age of 296 Ma (Menning, 1989; ff.) is used in the DCP 2003/ STD 2002, which is derived from ages from Central Europe (Lippolt and Hess, 1983; Lippolt et al., 1984; Lippolt and Hess, 1989; Lippolt et al., 1989). This age range encompasses those previously suggested viz., ca. 292 Ma (Chuvashov et al., 1996, SHRIMP; cf. Menning et al., 2000: Fig. 7), 298 Ma (Jones, 1995, Central European isotopic ages), and the 299 Ma (Ramezani et al., 2003, ID-TIMS) recently used in the GTS 2004 (Wardlaw et al.) as a maximum age. On one hand, the 299 Ma reduces the timespan for the Kasimovian and Gzhelian stages to a maximum of 7.5 my (306.5 – 299 Ma; Section 4.3.1). On the other hand, it significantly extends the timespan for the Cisuralian Period to 36.6 – 40 my. Consequently, the Early Permian stages are also extended. There is more than sufficient time to insert the Early Permian successions into the Early Permian time scale. Numerical ages for the base of the Sakmarian, Artinskian, and Kungurian stages are estimated using mainly geological time indications, because there are no U – Pb ID-TIMS ages from the type area. Isotopic ages from Central and West Europe as well as from Australia (Roberts et al., 1996), are difficult to transpose into the East European Reference Scale, because of weak biostratigraphic correlations. The SHRIMP ages of Chuvashov et al. (1996) may be too young in relation to the “ ID-TIMS Time Scale ” which currently appears to be in greater use. ID-TIMS ages are needed to decide which time spans are more reliable, for the Asselian Stage [the ca. 6 my of the DCP 2003 or the 4.4 ± 0.8 my of the GTS 2004 (Wardlaw et al.)]; for the Sakmarian Stage [the ca. 6 my of the DCP 2003 or the 10.2 ± 0.8 my of the GTS 2004 (Wardlaw et al.)], for the Artinskian Stage [the ca. 4.5 my of the DCP 2003 or the 8.8 ± 0.7 my of the GTS 2004 (Wardlaw et al.)], and for the Kungurian Stage [the ca. 7 my of the DCP 2003 or the 5.0 ± 0.7 my of the GTS 2004 (Wardlaw et al.) (Table 2, Fig. 1). The DCP 2003/STD 2002 time scale is based on Menning (2001). The age of ca. 284 Ma for the Sakmarian – Artinskian boundary results from the SHRIMP ages of 280.3 ± 2.6 from the latest Sakmarian and of 280.3 ± 2.4 Ma from the earliest Artinskian (Chuvashov et al., 1996) and from a correction of these ages, according to Compston (2000), of ca. 1.3%. According to the STD 2002 the diachronous lower boundary of the Rotliegend Group starts at ca. 302 Ma, an age based on a 302 ± 3.0 Ma U – Pb SHRIMP-II zircon age from the Holzmühlenthal ignimbrite (Breitkreuz and Kennedy, 1999), which is allocated to the lowermost Rotliegend of the Central European Basin (Fig. 4). Also SHRIMP-II ages between 300 ± 3 Ma and 297 ± 3 Ma were found in Rotliegend volcanics from six deep bore holes in that basin. From the Saar-Nahe Basin an 40 Ar/ 39 Ar age of 300.0 ± 2.4 Ma is cited (but not published in detail) for the “ Stephanian/Rotliegend boundary ” (Burger et al., 1997). Earliest Rotliegend volcanics from the Thuringian Forest (Möhrenbach Fm., Georgenthal Fm.=lower Gehren Schichten) yielded variable data: 299.1 ± 5.1 Ma to 292.0 ± 2.4 Ma 40 Ar/ 39 Ar ages, an 299.8 ± 4.4 Ma Rb – Sr model age (Goll et al., 1996) and 40 Ar/ 39 Ar ages from 295 ± 3 Ma to 291.0 ± 2 Ma (Goll and Lippolt, 2001). Thus, the age of ca. 302 Ma for the base of the Möhrenbach Formation (Fig. 4, STD 2002), which was derived by also taking into account the above ages of 302 ± 3.0 Ma, 300 ± 3 Ma to 297 ± 3 Ma, and 300.0 ± 2.4 Ma, is slightly too old (Fig. 4). The “ Lower ” and “ Upper ” Rotliegend are mapping units in Central Europe. The stratigraphic position of the boundary between these units changes in different locations. Thus, it corresponds traditionally to the boundary between the Rotterode and Tambach formations (sometimes, recently, at the Oberhof – Rotterode boundary) in the Thuringian Forest (Fig. 4, ca. 285 – 284 Ma, STD 2002) and the Glan and Nahe subgroups (former Tholey/Grenzlager boundary) in the Saar-Nahe Basin (Fig. 4, ca. 292 Ma, STD 2002). Thus, the use of the terms Lower and Upper Rotliegend is under discussion, whereas there is no question about the use of the term Rotliegend and its well defined formations and subgroups (Fig. 4, STD 2002). “ Lower Rotliegend ” successions have their maximum duration in the Thuringian Forest ( ≤ 302 Ma to ∼ 285 Ma). Nevertheless, they span a significantly shorter time interval than the “ Upper Rotliegend ” , which starts earliest in the Saar-Nahe Basin at ∼ 292 Ma with the Nahe Subgroup and finishes in the Central European Basin with the top of the Elbe Subgroup at ∼ 258 Ma (Fig. 4). The overlying Zechstein Group starts with the Kupferschiefer (Copper Shale, Marl Slate), a most useful lithostratigraphic marker horizon between Lithuania in the east and the United Kingdom in the west. Furthermore the use of the terms “ Autunian ” , “ Saxonian ” , and “ Thuringian ” in a time sense is not advocated, because of their widely variable application and poor definition; otherwise, their citation should be accompanied by the name of the author and year of publication (Schneider, 2001). Even the definition of the base of the “ Autunian ” with the FOD of Autunia (former Callipteris ) conferta is flawed, because Doubinger (1956) and Broutin et al. (1990) have shown that in successions several hundred meters thick, alternations between “ Autunian ” rocks with A. conferta (meso- to xerophylous) and “ Stephanian ” rocks (hygro- to hydro- phylous) occur. The base of the “ Saxonian ” was first defined by Haubold and Katzung (1972), but the new ” index ” fossils are found only in one outcrop of the Tambach Formation of the Thuringian Forest (Fig. 4). Often the term “ Saxonian ” is synonymously used for “ Upper Rotliegend ” , meaning that the “ Saxonian ” extends up to the base of the Kupferschiefer. For more details see Kozur (1980). A “ Thuringian ” age is usually cited by palynostratigraphers, using palynomorphs, which are seldom found in the almost exclusively red coloured beds of monotonously thick successions of late Rotliegend. The macroflora, also, lacks the time sensitivity to serve as diagnostic age markers in the middle and late Permian of Central and West Europe. Therefore, the positions of the age proposed for the base of the Thuringian are highly variable, ranging from about the early Wellington Formation (early Leonardian; global early Kungurian; Remy, 1975) up to the base of the Zechstein Group (Wuchiapingian) (Fig. 4). The Rotliegend usually begins with thick deposits of volcanic rocks or coarse grained sediments. It consists of very variable successions of coarse to fine grained clastics of fluvial, lacustrine (including black shales and carbonates), aeolian, and sabkha origin. The oldest beds are often grey coloured, whereas the late Rotliegend is exclusively red coloured except for some thin limnic horizons in the Central European Basin (Southern Permian Basin). On one hand, the Rotliegend ...

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... Guadalupian Series by Dunbar et al. (1960). The base of the Wordian Stage is in the Getaway Limestone Member of the new Cherry Canyon Formation, which correlates with parts of the original middle and upper Word Formation of the Glass Mts. The base of the Capitanian Stage is within the Pinery Limestone Member of the Bell Canyon Formation which correlates to the main part of the Capitan Limestone (Dunbar et al., 1960, chart-column: American Standard Section). The name of the Lopingian Series is derived from the Loping Coal-bearing “ series ” in South China, named by von Richthofen in 1884 (Jin et al., 2003). Huang (1932) later established it as a formal series to include all Permian deposits that overlie the Maokou Limestone. The Lopingian is subdivided into the Wuchiapingian and Changhsingian stages. Their names are derived from the Wuchiaping and Changhsing (Changxing) formations (Furnish and Glenister, 1970; Kanmera and Nakazawa, 1973). The Lopingian finishes at the base of the Triassic (Yin et al., 2001), which is defined by the FAD of the conodont Hindeodus parvus . The GSSP for the basal boundary of the Wuchiapingian Stage, and also for the Lopingian Series/Epoch is placed at the FAD of the conodont Clarkina postbitteri postbitteri within an evolutionary lineage from C. postbitteri hongshuiensis to C. dukouensi s in Bed 6k at the Penglaitan Section, Laibin area, Guangxi Province, South China (Jin et al., 2001). The FAD of C. postbitteri sensu lato could also be used to approximate this boundary, as it is only 20 cm below the defining point at the base of Bed 6i upper at the Penglaitan Section. Historically, this boundary was intended to coincide with a global regression and has been documented as a level coincident with an important mass extinction event. This event is marked by a rapid change from Jinogondolella ( Mesogondolella ) to Clarkina for off- shore conodonts, and Sweetognathus to Iranognathus for nearshore shallow water conodonts, and the sudden extinction of verbeekinid fusulinids (Jin et al., 1998a). Wuchiapingian fusulinids are dominated by the fusulinids Codonofusiella and Reichelina . The stage is composed of the Laibinian and Laoshanian substages bounded by the base of the Clarkina leveni Zone. This level is approximately correlated with the base of the Anderssonoceras – Prototoceras ammonoid Zone and coincides with the surface of the Lopingian transgression. The Laibinian Substage contains the C . postbitteri postbitteri , C. dukouensis , and C . asymmetrica conodont zones, and the Laoshanian Substage contains the C . leveni , C. guangyuanensis , and Clarkina orientalis conodont zones as well as the Anderssonoceras – Prototoceras , Araxoceras – Konglingites , and Sanyangites ammonoid zones. The Changhsingian Stage was formally proposed as the last stage of the Palaeozoic Era with its stratotype in the Section D at Meishan, Changxing County, Zhejiang Province of China (Zhao et al., 1981) with a basal boundary between the C. orientalis and Clarkina subcarinata zones. Recently, the boundary GSSP was ratified by the IUGS at the FAD of the conodont Clarkina wangi within the lineage from Clarkina longicuspidata to C. wangi at a point 88 cm above the base of the Changhsing Formation in the lower part of Bed 4 (base of 4a-2). This point is just above the flooding surface of the second parasequence in the Changhsing Limestone (Jin et al., 2004). The basal part of this stage is also marked by the occurrence of advanced forms of Palaeofusulina , and the tapashani- tid and pseudotirolitid ammonoids. The Baoqingian and the Meishanian substages have been suggested, bounded by the FOD of Clarkina changxingensis within Bed 10. The Changhsingian – Indusian boundary GSSP is close to the top of the Changhsing Formation within the lowermost Yinkeng Formation (Fig. 4). The numerical age of 260.5 Ma for the Capitanian – Wuchiapingian boundary in the STD 2002/DCP 2003 is an estimation which is based on a) the 265.3 ± 0.2 Ma ID-TIMS age (Bowring et al., 1998) for a bentonite bed just below the base of the Capitanian Stage in its boundary stratotype at the Nipple Hill near the Guadalupe Mts., West Texas, b) the numerous ID- TIMS ages from the sections Meishan and Shangsi around the PTB (Bowring et al., 1998; cf. Menning, 2001: Fig. 1) as well as the SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan (Claoué- Long et al., 1991) which is only slightly older than the PTB in Bed 27c (Yin et al., 2001), and c) geological time indications (cf. Section 2.1). The latter suggest that the age of 253.4 ± 0.2 Ma from just above the base of the Changhsingian Stage (Bowring et al., 1998) may be little too young. In the DCP 2003 are allocated durations of ∼ 4 my to the Changhsingian, of ∼ 5.5 Ma to the Wuchiapingian, and of ∼ 4.5 my to the Capitanian stages (Figs. 1 and 4). When using the age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of ∼ 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The allocation of ∼ 0.4 my to each of the 25 Lopingian parasequences (17.3 in the Wuchiapingian and 7.7 in the Changhsingian) of Chen et al. (1998: Fig. 9) results in durations of ∼ 10 my for the Lopingian Epoch, of ∼ 6.9 my for the Wuchiapingian Stage and of ∼ 3.1 my for the Changhsingian Stage (Menning et al., 2005; Fig. 1). Thus, the age for the Capitanian – Wuchiapingian boundary is ∼ 261 Ma according to an age of ∼ 251 Ma for the PTB (STD 2002; GTS 2004) or ∼ 262.5 Ma according to an age of ∼ 252.5 Ma for the PTB (Menning et al., 2005). The corresponding ages for the base of the Changhsingian Stage are 254.1 Ma and 255.6 Ma respectively, whereas in the STD 2002/DCP 2003 an age of ∼ 255 Ma is suggested. According to the orbital interpretation of the 25 Lopingian parasequences of Chen et al. (1998: Fig. 9) the duration of the Wuchiapingian is slightly underestimated and that of the Changhsingian is slightly overestimated. The ages of ∼ 261 Ma and ∼ 262.5 Ma for the base of the Wuchiapingian Stage (Menning et al., 2005) are consistent with a zircon age of 259.3 ± 3 Ma from the Emeishan volcanics occurring around the Guadalupian – Lopingian boundary (Zhou et al., 2002), and approximates with the age of 260.8 ± 0.8 Ma from the Shangsi Section (Mundil et al., 2004). The latter age is not Early Wuchiapingian as suggested, but Late Wuchiapingian because Bed 7 contains conodonts of the C. orientalis Zone (Jin, this work). The age of 257.3 ± 0.3 Ma (Mundil et al., 2004) for the Early Changhsingian is significantly older than the orbital age of 255.6 Ma derived by Menning et al. (2005), which is based on the 252.6 ± 0.2 Ma age for bed 25 in Meishan (Mundil et al., 2004). In Central and West Europe the stratigraphic range of the succession equivalent in age to the Lopingian Epoch (STD 2002/DCP 2003: ∼ 260.5 Ma to 251 Ma) is under discussion. Whereas sediments of that age are rarely developed in the intra-Variscan basins, a complete succession is represented in the foreland basin between Central England and East Poland. There is no evidence of a break in time of ≥ 0.1 my in the migrating basin centre in North Germany during the 266 Ma to 229 Ma interval (STD 2002; Menning et al., 2005), but Wardlaw et al. (2004: Fig. 16.2) show a gap between the top of the Zechstein at 259.8 Ma and the base of the Buntsandstein at 251.0 Ma. In the Central European Basin late Rotliegend sediments consists of fluvial, aeolian, playa, and sabkha sediments, which are up to 2000 m thick (Deutsche Stratigraphische Kommission, 1995; Schröder et al., 1995). These sediments are covered by Zechstein and younger sediments (STD 2002), which are some kilometres thick and penetrated by more than 1000 wells. The mainly brownish red clastics of the Havel Subgroup include aeolian sandstones with significant resources of natural gas. The clastic – evaporitic Elbe Subgroup covers a much larger area and is better subdivided and locally correlated than the Havel Subgroup. In NW Germany it accumulated mainly in a large perennial salt lake (Gast, 1991; Legler et al., 2005) and its orbitally estimated duration is ca. 5.6 my (cf. Section 5.3.2; STD 2002/DCP 2003: ∼ 4 Ma). This duration is consistent with the age for the Rotliegend – Zechstein boundary at ∼ 258 Ma in the STD 2002/DCP 2003, as well as the Re – Os age of 257.3 ± 1.7 Ma of the Kupferschiefer (Copper Shale) at the base of the Zechstein (Brauns et al., 2003). The marine – lagoonal – terrestrial Zechstein, which is subdivided and correlated in detail using marker horizons (Richter-Bernburg, 1955), consists of fine grained clastics, carbonates, sulfates, and chlorides in the central basin and terrestrial conglomerates, sandstones, and claystones in marginal areas. As in the late Rotliegend, the cyclic sedimentation is related to climatic variations. According to magnetostratigraphic data (Menning et al., 1988), correlations (Menning, 1986, 2001) and geological time indications, the uppermost part of the Rotliegend Group and the lower part of the Zechstein Group belong to the Wuchiapingian Stage, whereas the main part of the Zechstein Group belongs to the Changhsingian Stage (Fig. 4; STD 2002). The base of the Wuchiapingian Stage is within the upper Rotliegend Elbe Subgroup (Fig. 4; STD 2002) or Havel Subgroup (Menning et al., 2005: Plate IX). In contrast, the latest occurrence of the conodont Merrillina divergens in the Early Dzhulfian (Early Wuchiapingian; Kozur, 1994) and its occurrence in the Zechsteinkalk (Bender and Stoppel, 1965), early Zechstein 1, Werra Folge, is evidence for a slightly older base for the Zechstein. The occurrence of the conodont Mesogondolella britannica in the earliest Zechstein of England (Kozur, 1998), the Kupferschiefer of the German North Sea (Legler et al., 2005), and the Dzhulfian of the Salt Range (Kozur, 1998) is consistent with Fig. 4. According to bio- and chemostratigraphic evidence (Kozur, 1999) and magnetostratigraphic ...

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... section for a possible change in the future. The base of the Visean was first formally defined in the Dinant Basin (SCCS, 1969). However, subsequent work revealed that the section chosen (Bastion section, near Dinant) was unsatisfactory, not only for long distance correlation, but also at the local scale. Entries of stratigraphically significant foraminifer taxa are facies controlled. In addition, the use of unspecified Eoparastaffella (Conil et al., 1969) for defining the base of the Visean was inappropriate. The SCCS Working Group on the Tournaisian/Visean boundary followed the proposal of Hance and Muchez (1995) to search for a criterion within the evolutionary lineage of the genus Eoparastaffella . Work by its members, summarized in Devuyst et al. (2003) led to the proposal that the first appearance of Eoparastaffella simplex in the lineage E . ovalis → E . simplex should be used in the choice of the base of the Visean. This proposal was formally approved by the Voting Members of the SCCS in 2002 (Work, 2002) (cf. Fig. 3). The proposed new GSSP for the base of the Visean at Pengchong (Guangxi) southern China (Devuyst et al., 2003) is in the process of ratification. The section exposes a unit of allochthonous mass-flow deposits (Pengchong Member, Luzhai Formation) which accumulated in a starved, intra-platform basin. The evolution of Eoparastaffella is exceptionally well documented. The base of the Visean is defined at the base of bed 85, in which E. simplex first appears. The conodont Gnathodus homopunctatus , a useful, but cryptogenic, guide for the base of the Visean, enters in bed 86, about 1 m above the boundary (see Section 4.2.5). This new Tournaisian – Visean boundary is significantly younger than some historical (pre-1969) interpretations. Certain commonly used index macrofossils, formerly regarded as earliest Visean, on microfossil evidence, now indicate latest Tournaisian, e.g. the productid brachiopod Levitusia humerosa in the shallow water facies of Western Europe was thought to indicate earliest Visean, in contrast to the long controversial, but now accepted, Russian dating of this biozone as late Late Tournaisian in the Kosvian Horizon of the Ural Mts. Similarly, microfossil evidence indicates the Ammonellipsites / Merocanites ammonoid community in Europe, North Africa, Middle Asia, and North America should be removed from the basal Visean and restricted to an expanded Late Ivorian. The Serpukhovian was proposed in 1890 and officially included into the stratigraphic scale of the USSR in 1971 as an equivalent of the Early Namurian. Its type area is at the south limb of the Moscow Basin about 100 km south of Moscow, in the vicinity of Serpukhov. Originally it was defined on the basis of brachiopod assemblages and later subdivided into several foraminiferal zones. The conodont and foraminiferal assemblages of the early Serpukhovian are very similar to those of the late Visean (Gibshman, 2001; Nikolaeva et al., 2002). The late Serpukhovian was redefined mainly using the more complete succession of the Donets Basin (late Arnsbergian Zapaltyubian Horizon). The boundary between the Visean and Serpukhovian stages has been traditionally recognized on the basis of goniatites, and placed at the base of the Cravenoceras leion Zone. Korn (1996) proposed Edmooroceras pseudocoronula as a better alternative, because it is geographically less restricted. However, although Mississippian goniatites provide high-resolution regional subdivisions, some endemic taxa are of limited use for effective interregional and intercontinental correlation. Conodonts have also been proposed for the identification of the Visean – Serpukhovian boundary in Europe (Skompski et al., 1995). A SCCS Task Group led by B. Richards is currently working on finding a suitable criterion and GSSP for the base of the Serpukhovian. The lineage Lochriea nodosa – Lochriea ziegleri , which is recorded and controlled by ammonoids in the Rhenish Slate Mts., Germany, has been selected for this purpose. The FAD of L. ziegleri (Nemirovskaya et al., 1994) has been proposed to define the Visean – Serpukhovian boundary, because of its wide geographic distribution (Nemyrovska, 2005). The species is used to recognize the Visean upper boundary in China (Wang and Qi, 2003) and the Ural Mts. (Nikolaeva et al., 2002). On the other hand Gibshman and Baranova (in press) propose the use of the foraminifer “ Millerella ” tortula in the lineage “ Endostaffella ” tortula (part) –“ M. ” tortula (part), and of Janischewskina delicata in the lineage J. typica – J. delicata . According to the authors, the recognition of “ M. ” tortula at the base of the Serpukhovian in the Moscow Region and the Pericaspian Basin offers the possibility of a correlation with the Middle Chesterian of North America. The numerical age for the base of the Tournaisian Stage (base of the Carboniferous Period) is between 354 Ma (Young and Laurie, 1996) and 362 Ma (Tucker et al., 1998) (Fig. 1). The 354 Ma age was derived from an Early Carboniferous SHRIMP age (Claoué-Long et al., 1993), which is now known to be too young because of the inherent problem in the SL13 zircon standard used for the determination. Nevertheless, the 354 Ma age was applied as a tie point in the Phanerozoic time scales of Young and Laurie (1996) and Gradstein and Ogg (1996), and in the Carboniferous time scales of Jones (1995) and Menning et al. (2000) (Fig. 1), not least, because the dated sample was collected from the S. sulcata conodont Zone in the well known Hasselbach section in the Rhenish Slate Mts. The older age of 362 Ma was derived mainly from the average U – Pb ID-TIMS age of 363.6 ± 1.6 Ma from the latest Famennian Piskahegan Group of New Brunswick, Canada (Tucker et al., 1998) and, less significantly, from a Rb – Sr isochron age of 361.0 ± 4.1 Ma for the S. sulcata Zone in Nanbiancun, South China (Yang et al., 1988). Because of the 8 my discrepancy between the 354 Ma and 362 Ma ages, a compromise age of 358 Ma is used for the STD 2002 and the DCP 2003. Recently, Trapp et al. (2004) derived an age of 360.7 Ma for the DCB from two tuffs from the Hasselbach section, Rhenish Slate Mts. Their age of 360.5 ± 0.8 Ma and the age of 353.7 ± 4.2 Ma (Claoué-Long et al., 1993; renormalized by Claoué-Long et al., 1995) are both from the same tuff in bed 79. Using the ages of 360.5 ± 0.8 Ma and 360.2 ± 0.7 Ma of Trapp et al. (2004) and the thicknesses of the Hasselbach section instead of the thicknesses of the Lali section (South China, Trapp et al., 2004), the age of the DCB is estimated at 361.4 Ma (Menning, this work, Fig. 1). A Re – Os age of 361.3 ± 2.4 Ma of a black shale from the Exshaw Formation, Canada (Selby and Creaser, 2005) is very close to it. The duration of the Tournaisian Stage in many charts is shown to vary between 10 my and 14 my (Table 2). However, U – Pb ID-TIMS ages from the Rhenish Slate Mts., which yielded an age of ca. 361 Ma for its base (Trapp et al., 2004) and ca. 342 Ma for its top (Trapp, pers. com.), suggest a longer duration of slightly less than 20 my. Thus, these ages suggest the Tournaisian is approximately as long as the Visean Stage, which to date, is, together with the Norian Stage, the longest of the Phanerozoic Eon (STD 2002; Menning et al., 2005). The age for the base of the Serpukhovian Stage has been changed from 333 Ma (Harland et al., 1990) to 327 – 325 Ma, but there are no reliable isotopic ages about this boundary. The DCP age of 326.5 Ma is based on the Pendleian/Arnsbergian coal-tonstein 40 Ar/ 39 Ar age of 324.6 Ma (Lippolt et al., 1984). If preference is given to the middle Arnsbergian age of 319.5 Ma (Lippolt et al., 1984) in lieu of the Pendleian/ Arnsbergian age of 324.6 Ma, an age of 325 Ma or even less, is more probable for the base of the Namurian/ Serpukhovian (cf. Menning et al., 2000: Fig. 7, Scale B). The DCB S. sulcata criterion is not applicable in the shallow carbonate platform facies of the Franco-Belgian Basin. Here, the DCB is traditionally placed at the top of the basal (1 m thick) bed ( “ Tn 1b α ” ) of the Hastière Formation, a coarse grained bioclastic grainstone with lithoclasts and ooids and a reworked Devonian fauna. This bed is included between “ known Devonian ” and “ known Carboniferous ” (Van Steenwinkel, 1992). The Early Tournaisian (mid-Hastarian) of Western Europe includes a remarkable anoxic black shale, just above the basal Siphonodella crenulata Zone. It is seen both in basinal facies (Lower Alum Shale and Ru β schiefer of Germany) and in shallow water facies (Schistes du Pont d'Arcole of Belgium, Malevkian on the East European Platform). Cephalopod limestones ( Gattendorfia Genozone) and ostracod shales (with “ finger print ” entomozoids) below this semiglobal event are reminiscent of the Frasnian/Famennian dysphotic/ aphotic facies development. The historical stratotype for the base of the Visean is located in Belgium in the Bastion section. The boundary lies within the uppermost part of the Leffe Formation which contains thin grainstone layers derived from shallower areas (Lees, 1997). The first Eoparastaffella sp. (a broken specimen unidentifiable at the level of species) is found less than 1 m below the entry of the conodont G. homopunctatus (Conil et al., 1989). The utilization of the original British stages of the Visean (George et al., 1976) has been greatly enhanced by the critical and detailed review of Riley (1993). In the British Chadian stratotype section (Chatburn, Craven Basin, NW England), Riley (1990, 1995) showed that the genus Eoparastaffella enters about 300 m higher than originally reported (George et al., 1976) and precedes the incoming of G. homopunctatus . Riley (1990) therefore introduced the term ‘ late Chadian ’ for that part of the Chadian recognized by Eoparastaffella and accessory taxa, such as G. homopunctatus . Revisions of the widely used Brigantian, Pendleian and Arnsbergian ammonoid zonations ...

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... in western Europe and North America (Ramsbottom and Saunders, 1984) which may be slightly younger (Popov, 1979). The Kasimovian Stage is designated as the third stage of the Pennsylvanian Epoch. This interval was separated from the upper part of the Moscovian in 1926 as the “ Teguliferina Horizon ” , a name changed later to Kasimovian Horizon (Dan'shin, 1947) and finally accepted as Kasimovian Stage (Teodorovich, 1949). Its recently nominated type area is in lower reaches of the Moskva River, about 80 km southeast of Moscow, but its name comes from the town of Kasimov in the Ryazan Region. For additional regional information see 4.3.3. The lower boundary was first defined on brachiopod evidence, and in 1962 it was selected at the base of the Protriticites pseudomontiparus – Obsoletes obsoletus foraminiferal Zone. In order to define the base of the Kasimovian in the DCP 2003, and to suggest index fossils for global correlation, the FODs of P. pseudomontiparus, O. obsoletus and the conodont Swadelina subexcelsa (Goreva and Alekseev, 2001) are chosen for the DCP 2003. The FAD of the above- mentioned fusulinids may be in a time span of ≤ 0.5 my. However, in recent years both fusulinid and conodont workers have suggested selecting a higher level for the base of the Kasimovian, at the FODs of the fusulinid Montiparus and of the conodonts Idiognathodus sagittalis (Alekseev and Goreva, 2002; Villa and Task Group, 2004). Because their use would require a redefinition of both the Moscovian and Kasimovian stages, Davydov (2002) proposed the selection of an index species within the evolutionary chronocline of Protriticites fusulinids. This fauna is widely distributed throughout the Tethyan and Boreal realms and has also been found in the Great Basin within middle to late Desmoinesian successions (Wahlman et al., 1997; Davydov et al., 1999). The Gzhelian Stage is designated as the fourth stage of the Pennsylvanian Epoch. It was introduced by Nikitin (1890) defined by a specific brachiopod assemblage, and with a stratotype about 60 km east of Moscow at the village Gzhel. For additional regional information see 4.3.3. The task of defining the base of the Gzhelian was accepted by the Moscovian/Kasimo- vian Boundary Task Group in 2002. From fusulinid data (Rauzer-Chernousova, 1941) its lower boundary coincides with the base of the Rauserites rossicus – Rauserites stuckenbergi Zone. Also the FAD of the fusulinids Daixina , Jigulites , and Rugosofusulina has been proposed as an operational index for the lower boundary of the Gzhelian Stage (Rozovskaya, 1975; Rauzer-Chernousova and Shchegolev, 1979; Davydov, 1990). These genera, however, do not occur in North America (except for the Canadian Arctic). The FOD of the conodont Idiognathodus simulator s. s. is an useful marker because it is easily traced in Europe, Asia and North America (Boardman and Work, 2004; cf. 4.3.6). The Shumardites – Vidrioceras ammonoid Genozone has been conventionally placed at the base of the Gzhelian Stage (Bogoslovskaya et al., 1999). In the southern Ural Mts., advanced species of Shumardites occur in the late Gzhelian (Popov et al., 1985; Davydov, 2001). In order to define its base for the DCP 2003, and to suggest index fossils for global correlation, the conodont I. simulator s.s. and the foraminifer R. rossicus are chosen in this paper. The FAD of both species may be in a time span of <0.5 my. The numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U – Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A (Riley et al., 1993; cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member (Lippolt et al., 1984). This is allocated to approximately the Pendleian – Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). When the 40 Ar/ 39 Ar age of 319.5 Ma ± 7.8 Ma of the Poruba Member (Lippolt et al., 1984) is used instead of the Jaklovec age, an age of 318 Ma or even younger is more plausible. Such an age of 318.1 ± 1.3 Ma has been used by Davydov et al. (2004: Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. (2000). Ages of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. 1). They are based on all the 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian age of Central Europe (Lippolt et al., 1984). In the DCP 2003/STD 2002 an age of 312 Ma is used. An age of younger than 310.5 Ma can be excluded in the case of the base of the Moscovian being within the Westphalian B (Duckmantian Substage). Ages of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe (Lippolt et al., 1984; Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma (Davydov et al. in the GTS, 2004) is an approach to make the largest possible timespan for the Kasimovian and Gzhelian stages, after the maximum age of the top of the Gzhelian was regarded as 299 Ma (Ramezani et al., 2003, ID-TIMS; cf. Section 5.2.1). Ages of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. (marine, GSSP area) and Central/West Europe (continental, the current source of most of the isotopic age data). In the DCP 2003/STD 2002 an age of ca. 302 Ma is used. It is based on the ages of ca. 305 Ma and ca. 296 Ma for the bases of the Kasimovian and Asselian stages (Menning, 1989; Menning et al., 2000; Menning, 2001b) using Davydov's (1992) data that from geological indications, the Gzhelian should be more than twice duration of the Kasimovian. The classical Late Carboniferous successions of Central and Northwest Europe are mainly paralic, with an upward increase of terrigenous material culminating in exclusively continental beds. The stratigraphy is based on very detailed lithostratigraphy in coal basins (with marker horizons as marine bands, coal seams, volcanic ash layers), on macrofloral and miospore zonations, and only to a lesser degree on marine and non-marine faunal zones. Shallow water carbonate faunas (foraminiferids) are totally absent. The most suitable ammonoid faunas are restricted to the early Bashkirian, where they have an extremely fine biozonal resolution from Chokierian to Marsdenian substages (Bisat and Hudson, 1943; Ramsbottom et al., 1979; Ramsbottom and Saunders, 1984). The last Carboniferous conodont assemblage occurs in the early Moscovian (at the base of Bolsovian; =Aegir [Mansfield] marine band). It provides an inter-regional correlation across Europe. Thus, there is an approximate correlation between the base of the Vereian (earliest Moscovian) and the base of the Westphalian C (Bolsovian) (Nemyrovska, herein). However, in Fig. 3 the lower boundary of the Vereian is shown within the upper Westphalian B (Duckmantian), as based on Menning et al. (2001b). Therefore, a downward pointing arrow is shown at the base of the Westphalian C (Fig. 3). Above this level, the macroflora (Remy and Remy, 1977; Kerp, 1988; Josten, 1991), palynomorphs (Clayton et al., 1977; Peppers, 1996), and terrestrial faunas (Haubold, 1970, 1973; Holub and Kozur, 1981; Schneider, 1982; Martens, 1983a,b, 1984; Schneider, 1985, 1996; Boy, 1987; Werneburg, 1989, 2001; Boy and Martens, 1991; Schneider and Werneburg, 1993; Hampe, 1994; Voigt, 2005) are useful for local and regional correlation, but these fossil groups rarely allow detailed intercontinental correlation. However, the carbonate successions of the Cantabrian Mts. (Spain) are correlated by fusulinids, conodonts, and ammonoids with those of eastern Europe (cf. Villa and Task Group, 2004). The Moscovian – Kasimovian boundary is within the Cantabrian (Fig. 3, arrow down at the base of the Cantabrian). In the Subvariscan Foredeep, between Upper Silesia (South Poland) and South Wales (UK), the average cumulative thickness is > 5000 m for rocks of Namurian and Westphalian age, compared with > 3000 m for those of Bashkirian and Moscovian age (Drozdzewski, in Menning et al., 2000: Table 3). In Central Europe, coal seams are concentrated in rocks of Namurian C (Yeadonian Substage) to Westphalian C (Bolsovian Substage) age, e.g. the Ruhr district (Fig. 3), but they start also in the Namurian A (latest Mississippian) in Upper Silesia, and finish also in the early Rotliegend (Gzhelian/Asselian) in the intra-Variscan Saar-Nahe Basin and Saale Basin (STD 2002). In several depocenters, paralic sediments accumulated continuously with minimal gaps, of less than 0.5 my. In the Saar-Nahe Basin, however, with lacustrine, deltaic, and fluvial sediments, some 4 km thick, an “ Asturian ” gap of unknown, but in maximum 2.5 my, duration (Menning et al., 2005) separates Westphalian D from Stephanian A (Barruelian Substage). This gap could approximately correspond to the regional Cantabrian Substage (Stage, Wagner, 1972). The Bashkirian Stage was proposed in the South Ural Mts. (Semikhatova, 1934), because it represented an interval of strata that was missing on the central East European Platform. Its base was defined, originally with the brachiopod Choristites bisulcatiformus , and later by the foraminifers Plectostaffella bogdanovkensis / Pseudostaffella antiqua (Semikhatova et al., 1979; Sinitsyna and Sinitsyn, 1987). After the GSSP for the MCB was defined in Nevada (Lane et al., 1999), the ...

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... is used for correlations throughout the northern and eastern margins of Pangaea (Solovieva, 1986), but its potential for this purpose is restricted by fusulinid provinciality. Thus, the species composition in Eurasia and America is quite different (Solovieva, 1977; Ross and Ross, 1988; Davydov, 1996). Instead, the conodonts D. donetzianus and/or I. postsulcatus are proposed as index fossils because of their wide distribution in the Moscow Basin, Ural Mts., NW Europe, Spain, Canadian Arctic, Alaska, South America, and Japan. Moreover, their evolution is clearly recognized (Nemyrovska, 1999; Goreva and Alekseev, 2001). This conodont based definition of the Bashkirian – Moscovian boundary does not coincide with the base of the Atokan Stage in North America as suggested by fusulinid studies (Miklukho-Maclay, 1963; Ross, 1979). The base of the Atokan, operationally defined with the FADs of the fusulinids Eoschubertella and Pseudostaffella (Sutherland and Manger, 1984; Groves, 1986) is significantly older than the base of the Moscovian proposed above. Groves et al. (1999) have suggested that the base of the Atokan lies within the Late Bashkirian, either within the late Tashastian, or at the base of the Asatauian Horizon (South Ural Mts.). The Bashkirian – Moscovian boundary coincides with either the base of the Winslowoceras – Diaboloceras ammonoid Genozone of the East European Platform and Ural Mts. (Ruzhentsev and Bogoslovskaya, 1978), or with the Eowellerites Genozone in western Europe and North America (Ramsbottom and Saunders, 1984) which may be slightly younger (Popov, 1979). The Kasimovian Stage is designated as the third stage of the Pennsylvanian Epoch. This interval was separated from the upper part of the Moscovian in 1926 as the “ Teguliferina Horizon ” , a name changed later to Kasimovian Horizon (Dan'shin, 1947) and finally accepted as Kasimovian Stage (Teodorovich, 1949). Its recently nominated type area is in lower reaches of the Moskva River, about 80 km southeast of Moscow, but its name comes from the town of Kasimov in the Ryazan Region. For additional regional information see 4.3.3. The lower boundary was first defined on brachiopod evidence, and in 1962 it was selected at the base of the Protriticites pseudomontiparus – Obsoletes obsoletus foraminiferal Zone. In order to define the base of the Kasimovian in the DCP 2003, and to suggest index fossils for global correlation, the FODs of P. pseudomontiparus, O. obsoletus and the conodont Swadelina subexcelsa (Goreva and Alekseev, 2001) are chosen for the DCP 2003. The FAD of the above- mentioned fusulinids may be in a time span of ≤ 0.5 my. However, in recent years both fusulinid and conodont workers have suggested selecting a higher level for the base of the Kasimovian, at the FODs of the fusulinid Montiparus and of the conodonts Idiognathodus sagittalis (Alekseev and Goreva, 2002; Villa and Task Group, 2004). Because their use would require a redefinition of both the Moscovian and Kasimovian stages, Davydov (2002) proposed the selection of an index species within the evolutionary chronocline of Protriticites fusulinids. This fauna is widely distributed throughout the Tethyan and Boreal realms and has also been found in the Great Basin within middle to late Desmoinesian successions (Wahlman et al., 1997; Davydov et al., 1999). The Gzhelian Stage is designated as the fourth stage of the Pennsylvanian Epoch. It was introduced by Nikitin (1890) defined by a specific brachiopod assemblage, and with a stratotype about 60 km east of Moscow at the village Gzhel. For additional regional information see 4.3.3. The task of defining the base of the Gzhelian was accepted by the Moscovian/Kasimo- vian Boundary Task Group in 2002. From fusulinid data (Rauzer-Chernousova, 1941) its lower boundary coincides with the base of the Rauserites rossicus – Rauserites stuckenbergi Zone. Also the FAD of the fusulinids Daixina , Jigulites , and Rugosofusulina has been proposed as an operational index for the lower boundary of the Gzhelian Stage (Rozovskaya, 1975; Rauzer-Chernousova and Shchegolev, 1979; Davydov, 1990). These genera, however, do not occur in North America (except for the Canadian Arctic). The FOD of the conodont Idiognathodus simulator s. s. is an useful marker because it is easily traced in Europe, Asia and North America (Boardman and Work, 2004; cf. 4.3.6). The Shumardites – Vidrioceras ammonoid Genozone has been conventionally placed at the base of the Gzhelian Stage (Bogoslovskaya et al., 1999). In the southern Ural Mts., advanced species of Shumardites occur in the late Gzhelian (Popov et al., 1985; Davydov, 2001). In order to define its base for the DCP 2003, and to suggest index fossils for global correlation, the conodont I. simulator s.s. and the foraminifer R. rossicus are chosen in this paper. The FAD of both species may be in a time span of <0.5 my. The numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U – Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A (Riley et al., 1993; cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member (Lippolt et al., 1984). This is allocated to approximately the Pendleian – Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). When the 40 Ar/ 39 Ar age of 319.5 Ma ± 7.8 Ma of the Poruba Member (Lippolt et al., 1984) is used instead of the Jaklovec age, an age of 318 Ma or even younger is more plausible. Such an age of 318.1 ± 1.3 Ma has been used by Davydov et al. (2004: Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. (2000). Ages of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. 1). They are based on all the 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian age of Central Europe (Lippolt et al., 1984). In the DCP 2003/STD 2002 an age of 312 Ma is used. An age of younger than 310.5 Ma can be excluded in the case of the base of the Moscovian being within the Westphalian B (Duckmantian Substage). Ages of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe (Lippolt et al., 1984; Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma (Davydov et al. in the GTS, 2004) is an approach to make the largest possible timespan for the Kasimovian and Gzhelian stages, after the maximum age of the top of the Gzhelian was regarded as 299 Ma (Ramezani et al., 2003, ID-TIMS; cf. Section 5.2.1). Ages of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. (marine, GSSP area) and Central/West Europe (continental, the current source of most of the isotopic age data). In the DCP 2003/STD 2002 an age of ca. 302 Ma is used. It is based on the ages of ca. 305 Ma and ca. 296 Ma for the bases of the Kasimovian and Asselian stages (Menning, 1989; Menning et al., 2000; Menning, 2001b) using Davydov's (1992) data that from geological indications, the Gzhelian should be more than twice duration of the Kasimovian. The classical Late Carboniferous successions of Central and Northwest Europe are mainly paralic, with an upward increase of terrigenous material culminating in exclusively continental beds. The stratigraphy is based on very detailed lithostratigraphy in coal basins (with marker horizons as marine bands, coal seams, volcanic ash layers), on macrofloral and miospore zonations, and only to a lesser degree on marine and non-marine faunal zones. Shallow water carbonate faunas (foraminiferids) are totally absent. The most suitable ammonoid faunas are restricted to the early Bashkirian, where they have an extremely fine biozonal resolution from Chokierian to Marsdenian substages (Bisat and Hudson, 1943; Ramsbottom et al., 1979; Ramsbottom and Saunders, 1984). The last Carboniferous conodont assemblage occurs in the early Moscovian (at the base of Bolsovian; =Aegir [Mansfield] marine band). It provides an inter-regional correlation across Europe. Thus, there is an approximate correlation between the base of the Vereian (earliest Moscovian) and the base of the Westphalian C (Bolsovian) (Nemyrovska, herein). However, in Fig. 3 the lower boundary of the Vereian is shown within the upper Westphalian B (Duckmantian), as based on Menning et al. (2001b). Therefore, a downward pointing arrow is shown at the base of the Westphalian C (Fig. 3). Above this level, the macroflora (Remy and Remy, 1977; Kerp, 1988; Josten, 1991), palynomorphs (Clayton et al., 1977; Peppers, 1996), and terrestrial faunas (Haubold, 1970, 1973; Holub and Kozur, 1981; Schneider, 1982; Martens, 1983a,b, 1984; Schneider, 1985, 1996; Boy, 1987; Werneburg, 1989, 2001; Boy and Martens, 1991; Schneider and Werneburg, 1993; Hampe, 1994; Voigt, 2005) are useful for local and regional correlation, but these fossil groups rarely allow detailed intercontinental correlation. However, the carbonate successions of the Cantabrian Mts. (Spain) are correlated by fusulinids, conodonts, and ammonoids with those of eastern Europe (cf. Villa and Task Group, 2004). The Moscovian – Kasimovian boundary is within the Cantabrian (Fig. 3, arrow down at the base of the Cantabrian). In the Subvariscan Foredeep, between Upper Silesia (South Poland) and South Wales (UK), the average cumulative ...

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... correlates to the main part of the Capitan Limestone (Dunbar et al., 1960, chart-column: American Standard Section). The name of the Lopingian Series is derived from the Loping Coal-bearing “ series ” in South China, named by von Richthofen in 1884 (Jin et al., 2003). Huang (1932) later established it as a formal series to include all Permian deposits that overlie the Maokou Limestone. The Lopingian is subdivided into the Wuchiapingian and Changhsingian stages. Their names are derived from the Wuchiaping and Changhsing (Changxing) formations (Furnish and Glenister, 1970; Kanmera and Nakazawa, 1973). The Lopingian finishes at the base of the Triassic (Yin et al., 2001), which is defined by the FAD of the conodont Hindeodus parvus . The GSSP for the basal boundary of the Wuchiapingian Stage, and also for the Lopingian Series/Epoch is placed at the FAD of the conodont Clarkina postbitteri postbitteri within an evolutionary lineage from C. postbitteri hongshuiensis to C. dukouensi s in Bed 6k at the Penglaitan Section, Laibin area, Guangxi Province, South China (Jin et al., 2001). The FAD of C. postbitteri sensu lato could also be used to approximate this boundary, as it is only 20 cm below the defining point at the base of Bed 6i upper at the Penglaitan Section. Historically, this boundary was intended to coincide with a global regression and has been documented as a level coincident with an important mass extinction event. This event is marked by a rapid change from Jinogondolella ( Mesogondolella ) to Clarkina for off- shore conodonts, and Sweetognathus to Iranognathus for nearshore shallow water conodonts, and the sudden extinction of verbeekinid fusulinids (Jin et al., 1998a). Wuchiapingian fusulinids are dominated by the fusulinids Codonofusiella and Reichelina . The stage is composed of the Laibinian and Laoshanian substages bounded by the base of the Clarkina leveni Zone. This level is approximately correlated with the base of the Anderssonoceras – Prototoceras ammonoid Zone and coincides with the surface of the Lopingian transgression. The Laibinian Substage contains the C . postbitteri postbitteri , C. dukouensis , and C . asymmetrica conodont zones, and the Laoshanian Substage contains the C . leveni , C. guangyuanensis , and Clarkina orientalis conodont zones as well as the Anderssonoceras – Prototoceras , Araxoceras – Konglingites , and Sanyangites ammonoid zones. The Changhsingian Stage was formally proposed as the last stage of the Palaeozoic Era with its stratotype in the Section D at Meishan, Changxing County, Zhejiang Province of China (Zhao et al., 1981) with a basal boundary between the C. orientalis and Clarkina subcarinata zones. Recently, the boundary GSSP was ratified by the IUGS at the FAD of the conodont Clarkina wangi within the lineage from Clarkina longicuspidata to C. wangi at a point 88 cm above the base of the Changhsing Formation in the lower part of Bed 4 (base of 4a-2). This point is just above the flooding surface of the second parasequence in the Changhsing Limestone (Jin et al., 2004). The basal part of this stage is also marked by the occurrence of advanced forms of Palaeofusulina , and the tapashani- tid and pseudotirolitid ammonoids. The Baoqingian and the Meishanian substages have been suggested, bounded by the FOD of Clarkina changxingensis within Bed 10. The Changhsingian – Indusian boundary GSSP is close to the top of the Changhsing Formation within the lowermost Yinkeng Formation (Fig. 4). The numerical age of 260.5 Ma for the Capitanian – Wuchiapingian boundary in the STD 2002/DCP 2003 is an estimation which is based on a) the 265.3 ± 0.2 Ma ID-TIMS age (Bowring et al., 1998) for a bentonite bed just below the base of the Capitanian Stage in its boundary stratotype at the Nipple Hill near the Guadalupe Mts., West Texas, b) the numerous ID- TIMS ages from the sections Meishan and Shangsi around the PTB (Bowring et al., 1998; cf. Menning, 2001: Fig. 1) as well as the SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan (Claoué- Long et al., 1991) which is only slightly older than the PTB in Bed 27c (Yin et al., 2001), and c) geological time indications (cf. Section 2.1). The latter suggest that the age of 253.4 ± 0.2 Ma from just above the base of the Changhsingian Stage (Bowring et al., 1998) may be little too young. In the DCP 2003 are allocated durations of ∼ 4 my to the Changhsingian, of ∼ 5.5 Ma to the Wuchiapingian, and of ∼ 4.5 my to the Capitanian stages (Figs. 1 and 4). When using the age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of ∼ 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The allocation of ∼ 0.4 my to each of the 25 Lopingian parasequences (17.3 in the Wuchiapingian and 7.7 in the Changhsingian) of Chen et al. (1998: Fig. 9) results in durations of ∼ 10 my for the Lopingian Epoch, of ∼ 6.9 my for the Wuchiapingian Stage and of ∼ 3.1 my for the Changhsingian Stage (Menning et al., 2005; Fig. 1). Thus, the age for the Capitanian – Wuchiapingian boundary is ∼ 261 Ma according to an age of ∼ 251 Ma for the PTB (STD 2002; GTS 2004) or ∼ 262.5 Ma according to an age of ∼ 252.5 Ma for the PTB (Menning et al., 2005). The corresponding ages for the base of the Changhsingian Stage are 254.1 Ma and 255.6 Ma respectively, whereas in the STD 2002/DCP 2003 an age of ∼ 255 Ma is suggested. According to the orbital interpretation of the 25 Lopingian parasequences of Chen et al. (1998: Fig. 9) the duration of the Wuchiapingian is slightly underestimated and that of the Changhsingian is slightly overestimated. The ages of ∼ 261 Ma and ∼ 262.5 Ma for the base of the Wuchiapingian Stage (Menning et al., 2005) are consistent with a zircon age of 259.3 ± 3 Ma from the Emeishan volcanics occurring around the Guadalupian – Lopingian boundary (Zhou et al., 2002), and approximates with the age of 260.8 ± 0.8 Ma from the Shangsi Section (Mundil et al., 2004). The latter age is not Early Wuchiapingian as suggested, but Late Wuchiapingian because Bed 7 contains conodonts of the C. orientalis Zone (Jin, this work). The age of 257.3 ± 0.3 Ma (Mundil et al., 2004) for the Early Changhsingian is significantly older than the orbital age of 255.6 Ma derived by Menning et al. (2005), which is based on the 252.6 ± 0.2 Ma age for bed 25 in Meishan (Mundil et al., 2004). In Central and West Europe the stratigraphic range of the succession equivalent in age to the Lopingian Epoch (STD 2002/DCP 2003: ∼ 260.5 Ma to 251 Ma) is under discussion. Whereas sediments of that age are rarely developed in the intra-Variscan basins, a complete succession is represented in the foreland basin between Central England and East Poland. There is no evidence of a break in time of ≥ 0.1 my in the migrating basin centre in North Germany during the 266 Ma to 229 Ma interval (STD 2002; Menning et al., 2005), but Wardlaw et al. (2004: Fig. 16.2) show a gap between the top of the Zechstein at 259.8 Ma and the base of the Buntsandstein at 251.0 Ma. In the Central European Basin late Rotliegend sediments consists of fluvial, aeolian, playa, and sabkha sediments, which are up to 2000 m thick (Deutsche Stratigraphische Kommission, 1995; Schröder et al., 1995). These sediments are covered by Zechstein and younger sediments (STD 2002), which are some kilometres thick and penetrated by more than 1000 wells. The mainly brownish red clastics of the Havel Subgroup include aeolian sandstones with significant resources of natural gas. The clastic – evaporitic Elbe Subgroup covers a much larger area and is better subdivided and locally correlated than the Havel Subgroup. In NW Germany it accumulated mainly in a large perennial salt lake (Gast, 1991; Legler et al., 2005) and its orbitally estimated duration is ca. 5.6 my (cf. Section 5.3.2; STD 2002/DCP 2003: ∼ 4 Ma). This duration is consistent with the age for the Rotliegend – Zechstein boundary at ∼ 258 Ma in the STD 2002/DCP 2003, as well as the Re – Os age of 257.3 ± 1.7 Ma of the Kupferschiefer (Copper Shale) at the base of the Zechstein (Brauns et al., 2003). The marine – lagoonal – terrestrial Zechstein, which is subdivided and correlated in detail using marker horizons (Richter-Bernburg, 1955), consists of fine grained clastics, carbonates, sulfates, and chlorides in the central basin and terrestrial conglomerates, sandstones, and claystones in marginal areas. As in the late Rotliegend, the cyclic sedimentation is related to climatic variations. According to magnetostratigraphic data (Menning et al., 1988), correlations (Menning, 1986, 2001) and geological time indications, the uppermost part of the Rotliegend Group and the lower part of the Zechstein Group belong to the Wuchiapingian Stage, whereas the main part of the Zechstein Group belongs to the Changhsingian Stage (Fig. 4; STD 2002). The base of the Wuchiapingian Stage is within the upper Rotliegend Elbe Subgroup (Fig. 4; STD 2002) or Havel Subgroup (Menning et al., 2005: Plate IX). In contrast, the latest occurrence of the conodont Merrillina divergens in the Early Dzhulfian (Early Wuchiapingian; Kozur, 1994) and its occurrence in the Zechsteinkalk (Bender and Stoppel, 1965), early Zechstein 1, Werra Folge, is evidence for a slightly older base for the Zechstein. The occurrence of the conodont Mesogondolella britannica in the earliest Zechstein of England (Kozur, 1998), the Kupferschiefer of the German North Sea (Legler et al., 2005), and the Dzhulfian of the Salt Range (Kozur, 1998) is consistent with Fig. 4. According to bio- and chemostratigraphic evidence (Kozur, 1999) and magnetostratigraphic data of Li and Wang (1989) and Szurlies et al. (2003), the Zechstein – Buntsandstein boundary is within the latest Permian, ∼ 0.1 my older than the Changhsingian – Indusian boundary (Kozur, 2003; Menning et al., 2005). The two Lopingian Tethyan regional stages, the Dzhulfian and Dorashamian, are widely used in the literature, especially in the central ...

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... et al. (1999) have suggested that the base of the Atokan lies within the Late Bashkirian, either within the late Tashastian, or at the base of the Asatauian Horizon (South Ural Mts.). The Bashkirian – Moscovian boundary coincides with either the base of the Winslowoceras – Diaboloceras ammonoid Genozone of the East European Platform and Ural Mts. (Ruzhentsev and Bogoslovskaya, 1978), or with the Eowellerites Genozone in western Europe and North America (Ramsbottom and Saunders, 1984) which may be slightly younger (Popov, 1979). The Kasimovian Stage is designated as the third stage of the Pennsylvanian Epoch. This interval was separated from the upper part of the Moscovian in 1926 as the “ Teguliferina Horizon ” , a name changed later to Kasimovian Horizon (Dan'shin, 1947) and finally accepted as Kasimovian Stage (Teodorovich, 1949). Its recently nominated type area is in lower reaches of the Moskva River, about 80 km southeast of Moscow, but its name comes from the town of Kasimov in the Ryazan Region. For additional regional information see 4.3.3. The lower boundary was first defined on brachiopod evidence, and in 1962 it was selected at the base of the Protriticites pseudomontiparus – Obsoletes obsoletus foraminiferal Zone. In order to define the base of the Kasimovian in the DCP 2003, and to suggest index fossils for global correlation, the FODs of P. pseudomontiparus, O. obsoletus and the conodont Swadelina subexcelsa (Goreva and Alekseev, 2001) are chosen for the DCP 2003. The FAD of the above- mentioned fusulinids may be in a time span of ≤ 0.5 my. However, in recent years both fusulinid and conodont workers have suggested selecting a higher level for the base of the Kasimovian, at the FODs of the fusulinid Montiparus and of the conodonts Idiognathodus sagittalis (Alekseev and Goreva, 2002; Villa and Task Group, 2004). Because their use would require a redefinition of both the Moscovian and Kasimovian stages, Davydov (2002) proposed the selection of an index species within the evolutionary chronocline of Protriticites fusulinids. This fauna is widely distributed throughout the Tethyan and Boreal realms and has also been found in the Great Basin within middle to late Desmoinesian successions (Wahlman et al., 1997; Davydov et al., 1999). The Gzhelian Stage is designated as the fourth stage of the Pennsylvanian Epoch. It was introduced by Nikitin (1890) defined by a specific brachiopod assemblage, and with a stratotype about 60 km east of Moscow at the village Gzhel. For additional regional information see 4.3.3. The task of defining the base of the Gzhelian was accepted by the Moscovian/Kasimo- vian Boundary Task Group in 2002. From fusulinid data (Rauzer-Chernousova, 1941) its lower boundary coincides with the base of the Rauserites rossicus – Rauserites stuckenbergi Zone. Also the FAD of the fusulinids Daixina , Jigulites , and Rugosofusulina has been proposed as an operational index for the lower boundary of the Gzhelian Stage (Rozovskaya, 1975; Rauzer-Chernousova and Shchegolev, 1979; Davydov, 1990). These genera, however, do not occur in North America (except for the Canadian Arctic). The FOD of the conodont Idiognathodus simulator s. s. is an useful marker because it is easily traced in Europe, Asia and North America (Boardman and Work, 2004; cf. 4.3.6). The Shumardites – Vidrioceras ammonoid Genozone has been conventionally placed at the base of the Gzhelian Stage (Bogoslovskaya et al., 1999). In the southern Ural Mts., advanced species of Shumardites occur in the late Gzhelian (Popov et al., 1985; Davydov, 2001). In order to define its base for the DCP 2003, and to suggest index fossils for global correlation, the conodont I. simulator s.s. and the foraminifer R. rossicus are chosen in this paper. The FAD of both species may be in a time span of <0.5 my. The numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U – Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A (Riley et al., 1993; cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member (Lippolt et al., 1984). This is allocated to approximately the Pendleian – Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). When the 40 Ar/ 39 Ar age of 319.5 Ma ± 7.8 Ma of the Poruba Member (Lippolt et al., 1984) is used instead of the Jaklovec age, an age of 318 Ma or even younger is more plausible. Such an age of 318.1 ± 1.3 Ma has been used by Davydov et al. (2004: Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. (2000). Ages of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. 1). They are based on all the 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian age of Central Europe (Lippolt et al., 1984). In the DCP 2003/STD 2002 an age of 312 Ma is used. An age of younger than 310.5 Ma can be excluded in the case of the base of the Moscovian being within the Westphalian B (Duckmantian Substage). Ages of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe (Lippolt et al., 1984; Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma (Davydov et al. in the GTS, 2004) is an approach to make the largest possible timespan for the Kasimovian and Gzhelian stages, after the maximum age of the top of the Gzhelian was regarded as 299 Ma (Ramezani et al., 2003, ID-TIMS; cf. Section 5.2.1). Ages of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. (marine, GSSP area) and Central/West Europe (continental, the current source of most of the isotopic age data). In the DCP 2003/STD 2002 an age of ca. 302 Ma is used. It is based on the ages of ca. 305 Ma and ca. 296 Ma for the bases of the Kasimovian and Asselian stages (Menning, 1989; Menning et al., 2000; Menning, 2001b) using Davydov's (1992) data that from geological indications, the Gzhelian should be more than twice duration of the Kasimovian. The classical Late Carboniferous successions of Central and Northwest Europe are mainly paralic, with an upward increase of terrigenous material culminating in exclusively continental beds. The stratigraphy is based on very detailed lithostratigraphy in coal basins (with marker horizons as marine bands, coal seams, volcanic ash layers), on macrofloral and miospore zonations, and only to a lesser degree on marine and non-marine faunal zones. Shallow water carbonate faunas (foraminiferids) are totally absent. The most suitable ammonoid faunas are restricted to the early Bashkirian, where they have an extremely fine biozonal resolution from Chokierian to Marsdenian substages (Bisat and Hudson, 1943; Ramsbottom et al., 1979; Ramsbottom and Saunders, 1984). The last Carboniferous conodont assemblage occurs in the early Moscovian (at the base of Bolsovian; =Aegir [Mansfield] marine band). It provides an inter-regional correlation across Europe. Thus, there is an approximate correlation between the base of the Vereian (earliest Moscovian) and the base of the Westphalian C (Bolsovian) (Nemyrovska, herein). However, in Fig. 3 the lower boundary of the Vereian is shown within the upper Westphalian B (Duckmantian), as based on Menning et al. (2001b). Therefore, a downward pointing arrow is shown at the base of the Westphalian C (Fig. 3). Above this level, the macroflora (Remy and Remy, 1977; Kerp, 1988; Josten, 1991), palynomorphs (Clayton et al., 1977; Peppers, 1996), and terrestrial faunas (Haubold, 1970, 1973; Holub and Kozur, 1981; Schneider, 1982; Martens, 1983a,b, 1984; Schneider, 1985, 1996; Boy, 1987; Werneburg, 1989, 2001; Boy and Martens, 1991; Schneider and Werneburg, 1993; Hampe, 1994; Voigt, 2005) are useful for local and regional correlation, but these fossil groups rarely allow detailed intercontinental correlation. However, the carbonate successions of the Cantabrian Mts. (Spain) are correlated by fusulinids, conodonts, and ammonoids with those of eastern Europe (cf. Villa and Task Group, 2004). The Moscovian – Kasimovian boundary is within the Cantabrian (Fig. 3, arrow down at the base of the Cantabrian). In the Subvariscan Foredeep, between Upper Silesia (South Poland) and South Wales (UK), the average cumulative thickness is > 5000 m for rocks of Namurian and Westphalian age, compared with > 3000 m for those of Bashkirian and Moscovian age (Drozdzewski, in Menning et al., 2000: Table 3). In Central Europe, coal seams are concentrated in rocks of Namurian C (Yeadonian Substage) to Westphalian C (Bolsovian Substage) age, e.g. the Ruhr district (Fig. 3), but they start also in the Namurian A (latest Mississippian) in Upper Silesia, and finish also in the early Rotliegend (Gzhelian/Asselian) in the intra-Variscan Saar-Nahe Basin and Saale Basin (STD 2002). In several depocenters, paralic sediments accumulated continuously with minimal gaps, of less than 0.5 my. In the Saar-Nahe Basin, however, with lacustrine, deltaic, and fluvial sediments, some 4 km thick, an “ Asturian ” gap of unknown, but in maximum 2.5 my, duration (Menning et al., 2005) separates Westphalian D from Stephanian A (Barruelian Substage). This gap could approximately correspond to the regional Cantabrian Substage (Stage, Wagner, 1972). The Bashkirian Stage was proposed in the South Ural Mts. (Semikhatova, ...

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... appearance of the primitive fusulinids Eoschubertella and Pseudostaffella . This placed the Diaboloceras ammonoid Zone in the basal Atokan, which made the lower Atokan equivalent to the Duckmantian Substage (Westphalian B) in western Europe. The Atokan now encompasses three fusulinid genozones, in ascending order: Eoschubertella – Pseudostaffella, Profusulinella , and Fusulinella . Conodont faunas are not yet well enough known to define the base of the Atokan, but zones based on N. atokaensis and N. colombiensis (the latter zone including N. bothrops ) characterize the middle and late parts, respectively (Barrick et al., 2004). The overlying regional Desmoinesian Stage was named from central Iowa. It comprises ∼ 160 – 250 m of cyclic shale, coal, limestone, and sandstone across the northern Midcontinent shelf, thickening southward into the basin. It coincides with the Zone of Fusulina (now Beedeina ), with Wedekindellina also characterizing the lower half. Boardman et al. (1994) recognized it to comprise the Wellerites ammonoid Genozone overlain by the Eothalassoceras Genozone. Barrick et al. (2004) recognized three ascending conodont zones based on Neognathodus ( caudatus, asymmetricus, roundyi ), and five zones based on idiognathodids ( Idiognathodus neoshoensis, Sw. nodocarinata ). The base of the Desmoinesian has not yet been formally defined in the Midcontinent, but in the most complete succession in NE Oklahoma, I. praeobliquus followed by I. obliquus appear not far above the traditional boundary (Boardman et al., 2004). The two Swadelina zones at the top indicate correlation with the Krevyakinian Substage, which is currently the base of the global Kasimovian Stage. The early and middle Desmoinesian are equivalent to the late global Moscovian Stage. The Desmoinesian is also characterized by arborescent lycopods and their palynomorphs (Peppers, 1996), and is roughly equivalent to the latest substage (D, now named Asturian) of the Westphalian in W Europe. The overlying regional Missourian Stage was named from NW Missouri, and it comprises ∼ 150 – 200 m of cyclic limestone and shale (with some sandstone) across the northern Midcontinent shelf, thickening southward into the basin. It coincides with the lower part of the Zone of Triticites , although that genus first appears only in the fourth transgressive – regressive cycle above the base, above the first appearance of Eowaeringella in the third cycle. The two lower cycles lack large fusulinids (Heckel, 1999). Boardman et al. (1994) recognized in ascending order the Pennoceras, Preshumardites , and Pseudaktubites ammonoid genozones, with the latter zone extending above the top. Heckel et al. (2002) designated the base of the Missourian at the FAD of the conodont Idiognathodus eccentricus in the Exline cyclothem, and Barrick et al. (2004) recognized five ascending Missourian conodont zones ( I. eccentricus, St. cancellosus , S. confragus, S. gracilis , and I. aff. simulator ). Missourian palynomorphs suggest correlation with the Stephanian of western Europe, but the exact correlation of its basal Cantabrian Substage is not yet established. The regional Virgilian Stage at the top of the Pennsylvanian was named from SE Kansas, and it comprises ∼ 420 m of cyclic limestone, shale and sandstone across the northern Midcontinent shelf. It coincides with the upper part of the Triticites Zone, and it contains Waeringella near the base, Dunbar- inella through the middle, and Leptotriticites throughout the top (G.P. Wahlman, manuscript in review). It now includes strata that had been called “ Bursumian ” and the basal part of the Wolfcampian (Foraker Limestone) below the FAD of the conodont S. isolatus , which now defines the global base of the Permian System/Period. Boardman et al. (1994) referred the earliest Virgilian to the late part of the Pseudaktubites ammonoid Genozone and the remainder to the Shumardites Genozone. Heckel (1999) tentatively defined the base of the Virgilian at the first appearance of the conodont St. zethus in the base of the Cass cyclothem, and Barrick et al. (2004) recognized five ascending Virgilian conodont zones ( St. zethus, I. simulator, St. virgilicus, St. brownvillensis , and St. wabaunsensis ). The best position for defining the base of the global Gzhelian Stage appears to be at the first appearance of I. simulator , which coincides with the base of the Shumardites Vidrioceras global ammonoid Genozone (Boardman and Work, 2004), and is above the base of the regional Virgilian Stage. The Permian System/Period (Murchison, 1841) is traditionally subdivided into the Lower Series and Upper Series (Early Epoch and Late Epoch). In East Europe the Early Permian consists of the Asselian, Sakmarian, Artinskian, and Kungurian stages and the Late Permian, starting at the base of the Solikamskian Horizon, includes the Ufimian, Kazanian, and Tatarian stages (Resolutions, 1965). These seven (supra)regional stages were also used as global stages (Fig. 1), although the base of the Ufimian was not well defined in large areas of the East European Platform (Grunt, 2005). The Permian of Central/West Europe, the Dyas (Marcou, 1859; Geinitz, 1861 – 1862), comprises the groups Rotliegend and Zechstein. The Rotliegend starts in the latest Carboniferous Gzhelian Stage and continues into the Late Permian Wuchiapingian Stage (Fig. 4). Therefore, it cannot be considered as “ Lower Permian ” as widely mentioned in the literature. The Zechstein can be correlated with the late Wuchiapingian and most (> 95%) of the Changhsingian Stage. The overlying Buntsandstein Group of the Germanic Trias starts in the latest Changhsingian Stage (Kozur, 1999) (Fig. 4). The Permian of China currently consists of the Chuanshanian, Yangsingian, and Lopingian regional series/epochs (Sheng and Jin, 1994). The Chuanshanian Series plus the Chihsian Subseries correspond to the Cisuralian Series (Fig. 4). The Permian of North America consists of the Wolfcampian, Leonardian, Guadalupian, and Ochoan regional series (Dunbar et al., 1960; Fig. 4). Jin et al. (1994) proposed officially the tripartite Permian scale which is now the internationally accepted, global standard (Fig. 4). The Middle Permian (Guadalupian Epoch) of this scale starts within the Late Permian (Ufimian, Kazanian, Tatarian) of the bipartite scale. According to Grunt (this work), the bipartite scale corresponds better to the biotic evolution than the tripartite scale. Therefore, Leven (2003) proposed to reactivate the global bipartite scale, by classifying its parts as subsystems (Cisuralian and Tethyan) each divided into two series (Fig. 4). The same structure utilizing proper names for subsystems and series has been proposed for the East-European Permian Scale (Grunt, 2005; Fig. 4): the Lower Permian Subsystem (Cisuralian) includes the Uralian Series (Asselian and Sakmarian stages) and the Cistimanian Series (Artinskian and Kungurian regional stages). The Upper Permian (Vyatkian) Subsystem consists of the Biarmian Series with the regional Ufimian, Kazanian, and Urzhumian stages and the Tatarian Series with the Midian, Dzhulfian, and Changhsingian stages. However, the introduction of a new stratigraphic category (subsystem), the re-classification of e.g. the Cisuralian (from a series to a subsystem) as well as of the Tatarian (from a stage into a series), and the significantly reduced stratigraphic range of the Tatarian (excluding the Urzhumian Horizon) would in consequence lead to even more confusion than to more clarity (cf. Fig. 4, 5.3.3) (Menning and Kozur (this work). Moreover, the strong extinction event at the Guadalupian – Lopingian boundary evidenced in many fossil groups shows that the differences between the Guadalupian and Lopingian faunas are much stronger than those between the Cisuralian and Guadalupian faunas (e.g. conodonts, fusulinids). Therefore, the general pattern of biotic evolution corresponds better to the tripartite scale than to the bipartite scale (Kozur, this work). The Cisuralian Epoch/Early Permian Epoch/Lower Permian Series has been formally established by Jin et al. (1994). Its type area is the southern Ural Mountain region (Urals), comprising the Asselian, Sakmarian, Artinskian, and Kungurian stages. These stages were initially defined and widely recognized on ammonoid phylogenies (Karpinsky, 1891; Ruzhentsev, 1937, 1950, 1951, 1955, 1956). However, the boundaries and subdivisions of those stages were established using fusulinaceans, the most abundant, and one of the best studied Late Palaeozoic fossil groups of the southern Urals. It is expected that the boundary stratotypes (GSSP) for all Cisuralian stages will be formally established in this region. The GSSP for the base of the Asselian Stage, the co-incident base of the Permian Period, is established at Aidaralash Creek, Aktöbe (formerly Aktyubinsk) region, northern Kazakhstan. The position of the GSSP is at the FOD of the conodont S. isolatus (Davydov et al., 1998). The FOD of S. invaginatus and S. nodulinearis nearly coincide with the FOD of S. isolatus and therefore, can be used as accessory indicators for the boundary. The GSSP is 6.3 m below the traditional fusulinid boundary, i.e. the base of the Sphaeroschwagerina vulgaris aktjubensis – S. fusiformis Zone (Davydov et al., 1998). The latter can be widely correlated via Spitsbergen, the East European Platform, the Ural Mts., Central Asia, China and Japan. It is of practical value to identify the top of the informal Orenburgian Substage of the late Gzhelian Stage. The traditional ammonoid boundary, 26.8 m above the GSSP, includes the termination of the Prouddenites – Uddenites lineage and the introduction of the Permian taxa Svetlanoceras primore and Prostacheoceras principale . At the Usolka Section numerous volcanic ash beds offer great potential for radiometric dating (Davydov et al., 2002). The Asselian conodont succession of Aidaralash and Usolka is also displayed in the ...

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... The lower boundary was first defined on brachiopod evidence, and in 1962 it was selected at the base of the Protriticites pseudomontiparus – Obsoletes obsoletus foraminiferal Zone. In order to define the base of the Kasimovian in the DCP 2003, and to suggest index fossils for global correlation, the FODs of P. pseudomontiparus, O. obsoletus and the conodont Swadelina subexcelsa (Goreva and Alekseev, 2001) are chosen for the DCP 2003. The FAD of the above- mentioned fusulinids may be in a time span of ≤ 0.5 my. However, in recent years both fusulinid and conodont workers have suggested selecting a higher level for the base of the Kasimovian, at the FODs of the fusulinid Montiparus and of the conodonts Idiognathodus sagittalis (Alekseev and Goreva, 2002; Villa and Task Group, 2004). Because their use would require a redefinition of both the Moscovian and Kasimovian stages, Davydov (2002) proposed the selection of an index species within the evolutionary chronocline of Protriticites fusulinids. This fauna is widely distributed throughout the Tethyan and Boreal realms and has also been found in the Great Basin within middle to late Desmoinesian successions (Wahlman et al., 1997; Davydov et al., 1999). The Gzhelian Stage is designated as the fourth stage of the Pennsylvanian Epoch. It was introduced by Nikitin (1890) defined by a specific brachiopod assemblage, and with a stratotype about 60 km east of Moscow at the village Gzhel. For additional regional information see 4.3.3. The task of defining the base of the Gzhelian was accepted by the Moscovian/Kasimo- vian Boundary Task Group in 2002. From fusulinid data (Rauzer-Chernousova, 1941) its lower boundary coincides with the base of the Rauserites rossicus – Rauserites stuckenbergi Zone. Also the FAD of the fusulinids Daixina , Jigulites , and Rugosofusulina has been proposed as an operational index for the lower boundary of the Gzhelian Stage (Rozovskaya, 1975; Rauzer-Chernousova and Shchegolev, 1979; Davydov, 1990). These genera, however, do not occur in North America (except for the Canadian Arctic). The FOD of the conodont Idiognathodus simulator s. s. is an useful marker because it is easily traced in Europe, Asia and North America (Boardman and Work, 2004; cf. 4.3.6). The Shumardites – Vidrioceras ammonoid Genozone has been conventionally placed at the base of the Gzhelian Stage (Bogoslovskaya et al., 1999). In the southern Ural Mts., advanced species of Shumardites occur in the late Gzhelian (Popov et al., 1985; Davydov, 2001). In order to define its base for the DCP 2003, and to suggest index fossils for global correlation, the conodont I. simulator s.s. and the foraminifer R. rossicus are chosen in this paper. The FAD of both species may be in a time span of <0.5 my. The numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U – Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A (Riley et al., 1993; cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member (Lippolt et al., 1984). This is allocated to approximately the Pendleian – Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). When the 40 Ar/ 39 Ar age of 319.5 Ma ± 7.8 Ma of the Poruba Member (Lippolt et al., 1984) is used instead of the Jaklovec age, an age of 318 Ma or even younger is more plausible. Such an age of 318.1 ± 1.3 Ma has been used by Davydov et al. (2004: Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. (2000). Ages of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. 1). They are based on all the 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian age of Central Europe (Lippolt et al., 1984). In the DCP 2003/STD 2002 an age of 312 Ma is used. An age of younger than 310.5 Ma can be excluded in the case of the base of the Moscovian being within the Westphalian B (Duckmantian Substage). Ages of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe (Lippolt et al., 1984; Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma (Davydov et al. in the GTS, 2004) is an approach to make the largest possible timespan for the Kasimovian and Gzhelian stages, after the maximum age of the top of the Gzhelian was regarded as 299 Ma (Ramezani et al., 2003, ID-TIMS; cf. Section 5.2.1). Ages of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. (marine, GSSP area) and Central/West Europe (continental, the current source of most of the isotopic age data). In the DCP 2003/STD 2002 an age of ca. 302 Ma is used. It is based on the ages of ca. 305 Ma and ca. 296 Ma for the bases of the Kasimovian and Asselian stages (Menning, 1989; Menning et al., 2000; Menning, 2001b) using Davydov's (1992) data that from geological indications, the Gzhelian should be more than twice duration of the Kasimovian. The classical Late Carboniferous successions of Central and Northwest Europe are mainly paralic, with an upward increase of terrigenous material culminating in exclusively continental beds. The stratigraphy is based on very detailed lithostratigraphy in coal basins (with marker horizons as marine bands, coal seams, volcanic ash layers), on macrofloral and miospore zonations, and only to a lesser degree on marine and non-marine faunal zones. Shallow water carbonate faunas (foraminiferids) are totally absent. The most suitable ammonoid faunas are restricted to the early Bashkirian, where they have an extremely fine biozonal resolution from Chokierian to Marsdenian substages (Bisat and Hudson, 1943; Ramsbottom et al., 1979; Ramsbottom and Saunders, 1984). The last Carboniferous conodont assemblage occurs in the early Moscovian (at the base of Bolsovian; =Aegir [Mansfield] marine band). It provides an inter-regional correlation across Europe. Thus, there is an approximate correlation between the base of the Vereian (earliest Moscovian) and the base of the Westphalian C (Bolsovian) (Nemyrovska, herein). However, in Fig. 3 the lower boundary of the Vereian is shown within the upper Westphalian B (Duckmantian), as based on Menning et al. (2001b). Therefore, a downward pointing arrow is shown at the base of the Westphalian C (Fig. 3). Above this level, the macroflora (Remy and Remy, 1977; Kerp, 1988; Josten, 1991), palynomorphs (Clayton et al., 1977; Peppers, 1996), and terrestrial faunas (Haubold, 1970, 1973; Holub and Kozur, 1981; Schneider, 1982; Martens, 1983a,b, 1984; Schneider, 1985, 1996; Boy, 1987; Werneburg, 1989, 2001; Boy and Martens, 1991; Schneider and Werneburg, 1993; Hampe, 1994; Voigt, 2005) are useful for local and regional correlation, but these fossil groups rarely allow detailed intercontinental correlation. However, the carbonate successions of the Cantabrian Mts. (Spain) are correlated by fusulinids, conodonts, and ammonoids with those of eastern Europe (cf. Villa and Task Group, 2004). The Moscovian – Kasimovian boundary is within the Cantabrian (Fig. 3, arrow down at the base of the Cantabrian). In the Subvariscan Foredeep, between Upper Silesia (South Poland) and South Wales (UK), the average cumulative thickness is > 5000 m for rocks of Namurian and Westphalian age, compared with > 3000 m for those of Bashkirian and Moscovian age (Drozdzewski, in Menning et al., 2000: Table 3). In Central Europe, coal seams are concentrated in rocks of Namurian C (Yeadonian Substage) to Westphalian C (Bolsovian Substage) age, e.g. the Ruhr district (Fig. 3), but they start also in the Namurian A (latest Mississippian) in Upper Silesia, and finish also in the early Rotliegend (Gzhelian/Asselian) in the intra-Variscan Saar-Nahe Basin and Saale Basin (STD 2002). In several depocenters, paralic sediments accumulated continuously with minimal gaps, of less than 0.5 my. In the Saar-Nahe Basin, however, with lacustrine, deltaic, and fluvial sediments, some 4 km thick, an “ Asturian ” gap of unknown, but in maximum 2.5 my, duration (Menning et al., 2005) separates Westphalian D from Stephanian A (Barruelian Substage). This gap could approximately correspond to the regional Cantabrian Substage (Stage, Wagner, 1972). The Bashkirian Stage was proposed in the South Ural Mts. (Semikhatova, 1934), because it represented an interval of strata that was missing on the central East European Platform. Its base was defined, originally with the brachiopod Choristites bisulcatiformus , and later by the foraminifers Plectostaffella bogdanovkensis / Pseudostaffella antiqua (Semikhatova et al., 1979; Sinitsyna and Sinitsyn, 1987). After the GSSP for the MCB was defined in Nevada (Lane et al., 1999), the conodont D. noduliferus was selected also for the base of the Bashkirian in Russia. Thus, there is now total consis- tency between the base of the global Bashkirian Stage (Pennsylvanian Epoch) in North America and that of the regional Bashkirian Stage in East Europe. The Bashkirian contains two sets of regional substages: on the platform the Voznesenskian, Krasno- polyanian, Severokeltmenian, Prikamian, Cheremshan- kian, and Melekessian horizons, and in the foredeep of the southern Ural Mts., the type area, the Bogdanovkian ( D. noduliferus – Idiognathodus sinuatus ), Syuranian, Akavasian (upper part: I. sinuatus – Neognathodus askynensis ), ...

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... which since the introduction of the GSSP concept, would be of late Gzhelian age. Later, the boundary was proposed at the Neva Limestone (Baars et al., 1994) of Middle Asselian age. Recently, the CPB is precisely placed at the base of the Glenrock Limestone of the Red Eagle Cyclothem (Fig. 4) within the conodont succession S. wabaunsensis – S. isolatus at the FAD of the latter species (Chernykh et al., 1997; Boardman et al., 1998). Ross and Ross (1987) proposed the Bursumian as the new terminal stage of the Pennsylvanian. However, for several reasons (Davydov, 2001) this usage was never adopted. The Bursumian covers a late Gzhelian to middle Asselian time span. In its type section of Lenox Hill in the Glass Mts. the Lenoxian Stage contains typical Artinskian conodonts as Sw. whitei and Neostreptognathodus transitus and the Yakhtashian fusulinid Chalaroschwagerina (Wardlaw and Davydov, 2000; Mei et al., 2002). In its type section the Hessian Stage contains at its base N. pequopensis . At 17 m above the base, a plethora of species of Mesogondolella and Neostreptognathodus occur, including N. exsculptus which approximately marks the base of the global Kungurian. In the Pequope Mts. the later taxon occurs together with N. pnevi (Wardlaw et al., 1998), the proposed index-species for the base of the global Kungurian (cf. Section 5.2.1). Thus, the regional Hessian Stage starts in the latest Artinskian. The Hessian and Cathedralian stages as well as the Hess Member, the Cathedral Mountain Formation and the lower Road Canyon Formation correlate approximately with the global Kungurian (Fig. 4). In the Midcontinent the FAD of the conodont Sw. merrilli places the base of the Sakmarian quite precisely within the Eiss Limestone of the Council Grove Group and the FAD of the conodont Sw. whitei places the base of the Artinskian within the Florence Limestone of the Chase Group (Boardman et al., 1998; Mei et al., 2002; Fig. 4). The three GSSPs are located in the Guadalupe Mts., West Texas (for details: Section 5.3.6); the stages are defined by conodonts (Glenister et al., 1999). The term Roadian is derived from the Road Canyon Formation of the Glass Mts., the former lower Word Formation (for details cf. DCP 2003). The base of the Roadian Stage is defined in the El Centro Member, middle Cutoff Formation, by the FAD of Jinogondolella nankingensis . In the STD 2002/DCP 2003 its numerical age is taken at ca. 272.5 Ma (Menning and Deutsche Stratigraphische Kommission, 2002); it may be a little younger, but certainly older than 270 Ma in order to avoid an overstretching of the Early Permian. In contrast, other age estimates for this boundary are 269 Ma (Kozur, 2003) and 270.6 ± 0.7 Ma (Gradstein et al., 2004). The base of the Wordian Stage is defined by the FAD of Jinogondolella aserrata within the Getaway Limestone, lower Cherry Canyon Formation. The base of the Capitanian Stage is defined with the FAD of Jinogondolella postserrata in the Pinery Limestone, lower Bell Canyon Formation at the Nipple Hill. Its age is ca. 265.0 Ma, derived from the U – Pb ID-TIMS age of 265.3 ± 0.2 Ma (Bowring et al., 1998) of a tuff between the Hegler Limestone (below) and the Pinery Limestone. It confirms the age of ≤ 265 Ma which was derived using geological time indications (Menning, 1995a,b). To date, there are neither isotopic, nor cyclostratigraphic data on which to estimate the numerical age of the Roadian – Wordian boundary. The stratigraphic positions of the bases of the Roadian and Wordian stages within the terrestrial late Rotliegend are unknown. In the Central European Basin both boundaries are in a large gap between the Müritz Subgroup and the Havel Subgroup (Menning, 1995a,b; STD 2002: ca. 282 Ma to 266 Ma; Fig. 4). The position of the Illawarra Reversal has been recognized in the Parchim Formation (Menning et al., 1988) and in the latest Wordian of the Guadalupe Mts. (Menning et al., 2005), and thus, the base of the Capitanian Stage lies within that formation. The Rotliegend of the intra- Variscan basins is almost exclusively reversed magnetized. Therefore, it has a pre-Illawarra, pre-Capitanian age of ≥ 265 Ma (Menning et al., 1988, 2005) excluding the latest Rotliegend Eisleben, Cornberg, and Battenberg formations (STD 2002). The four formations of the Havel and Elbe subgroups are nearly isochronous time slices which have been termed “ Folgen ” . The Dethlingen Folge and the Hannover Folge consists on 7 cycles each. A cycle has a duration of ca. 0.4 my. Thus, the Elbe Subgroup has a duration of ca. 5.6 my (Gast and Menning, in Menning et al., 2005; Fig. 4: 4 my). Only the latest cycle, directly below the Zechstein, shows marine influence. The fossils of the late Rotliegend are indicators of facies, rather than of time (Schneider et al., 1995). The Late Permian of East Europe consisted for 40 years of the three regional stages Ufimian, Kazanian, and Tatarian (Resolutions, 1965), but these names were also used for stages of the GSS (Fig. 1). Their stratotype and reference sections are described in detail in a monograph by Esaulova et al. (1998). The introduction of the GSSP concept resulted in the controversial subdivision of the Late Permian into the Middle and Late Permian, the substitution of the East European stages by North American and Chinese ones including new defined stage boundaries. In order to improve the equivalence with the new Middle and Late Permian GSS (cf. Sections 5.1 and 5.3.1), the All-Russian Conference (2002, 2004) and Grunt (2005) introduced the term “ Biarmian ” in a redefinition, and reclassification of East European units (Fig. 4). Kazanian deposits are wide-spread in Povolzhe and Prikame. Marine clays, marls, and limestones inter- finger with terrestrial red-beds (muddy sandstones). On one hand, the Kazanian Stage is subdivided into the Sokian Substage (type locality: Sok River Basin, Baitugan Village; Baitugan, Kamyshla, and Barbashi formations; Forsh, 1955) and the Povolzhian Substage (type locality: Volga river, near the Pechishchi village; Esaulova et al., 1998). On the other hand, the Kazanian Stage is subdivided into the Kamian Substage (Baitugan, Kamyshla, and Krasnoyarsk formations, total thickness 75 – 85 m, and the Krasnovidovian Substage (Prikazan, Pechishchi, Verkhniya Uslonka, and Morkvashino formations, total thickness 48 – 50 m; Grunt, 2005) (Fig. 4). The base of the Kazanian Stage was defined at the base of “ Spiriferida ” ( Licharewia ) Beds. Recently, the base of Licharewia schrencki Beds in the Cheshskaya Bay Section of the Kanin Peninsula (Grunt et al., 2002) has been proposed as the lower boundary stratotype. The conodont Kamagnathus ( = Stepanovites) khalim- badzhae from the Baitugan Formation was discovered together with the Roadian index conodont J. nankingensis ( Mesogondolella serrata ) in the Phosphoria Formation of Utah (Chernykh et al., 2001; Chernykh and Silantiev, 2004). Moreover, the presence of the ammonoid Sverdrupites in the late Early Kazanian (Leonova et al., 2002) allows the approximate correlation of the early Kazanian Stage (RSS) with the early Roadian Stage (GSS). According to Chernykh, the entire Kazanian corresponds to the Kamagnathus Genozone and could be correlated with the Roadian Stage and therefore, the entire Ufimian Stage is probably older than the Roadian Stage. However, the suggestion of Kotlyar (2002) that the latest Ufimian Stage correlates with the earliest Roadian Stage, cannot be excluded (Fig. 4: arrow up at the upper boundary of the Ufimian Stage). The original Tatarian consists of the Early Tatarian Urzhumian Horizon and the Late Tatarian Severodvinian and Vyatkian horizons (Fig. 4). Recently, it was decided, to separate the Urzhumian from the Tatarian in order to restrict the latter to the former Late Tatarian, and to re-classify this reduced Tatarian as a regional Tatarian Series/Epoch corresponding to the global Capitanian, Wuchiapingian, and Changhsingian stages (Resolutions, 2006). The older, newly introduced regional Biarmian Series/Epoch, consists of the regional, Kazanian and Urzhumian stages, which correspond approximately to the global, Roadian and Wordian stages (Fig. 4). The boundary between the newly introduced supraregional Biarmian and Tatarian epochs/series approximates to the Illawarra Reversal, which separates the Carboniferous – Permian (Kiaman) Reversed polarized Superchrone/Megazone from the Permian – Triassic (Illawarra) Mixed polarized Superchrone/ Megazone. It is the most pronounced magnetostratigraphic marker in the Palaeozoic Era. The Urzhumian Horizon (Fredericks, 1918) is now re- classified as regional stage (Resolutions, 2006) which corresponds approximately to the global Wordian Stage. The upper boundaries of the Wordian and Urzhumian stages correlate exactly when using the Illawarra Reversal, which is positioned in the latest Urzhumian (Khramov, 1963) and the latest Wordian (Glenister et al., 1999; Burov et al., 2002; Menning et al., 2005). The Urzhumian is composed of interbedded sandstone-shale and rhythmic mudstone and limestone, and coincides with the Platysomus biarmicus – Kargalichthys efremovi ichthyological Zone. Its lower boundary is marked by the base of the Palaeodarwinula fragiliformis – Prasuchonella nasalis ostracod Zone. According to the Resolutions (2006), the new supraregional Tatarian Period/Series consists of the Severodvinian and Vyatkian stages, which were introduced by Ignatiev (1962) as horizons. The Severodvinian Horizon (Fig. 4) is composed of mainly alluvial – limnic sandstone – siltstone and variegated marly limestone. The basal boundary of its Early Substage is defined at the base of the Suchonellina inornata – P. nasalis ostracod Zone. Its Late Substage is fixed at the base of the S. inornata – Prasuchonella stelmachovi ostracod Zone, coinciding with the Deltavjatia vjatken- sis tetrapod Zone. The Vyatkian Horizon (Fig. 4) is composed of fluvial, alluvial plains, and limnic facies. ...

Context 20

... the beginning of Project 1054 of the Deutsche Forschungsgemeinschaft (DFG) “ The evolution of the Late Palaeozoic in the light of sedimentary geochemistry ” German geochemists were using time scales with significant age differences. Therefore, it was decided to use a uniform time scale and a stratigraphic correlation chart with sections from several continents. This was done to ensure variations in isotopic profiles of carbon, oxygen, sulfur, nitrogen, boron, and osmium would be directly comparable, and hence prove to be more reliable indicators of isochronous global events. During 2001 – 2002 the “ Stratigraphische Tabelle von Deutschland 2002 ” (herein abbreviated to STD 2002, and cited as STD, 2002 in references) was prepared by the Deutsche Stratigraphische Kommission (www.stra- tigraphie.de). At that time an updating of the global time scale was necessary in order to integrate isotopic age determinations that had become available over the previous decade since the publication of the previous Phanerozoic time scales (Harland et al., 1990; Odin, 1994; Gradstein and Ogg, 1996; Young and Laurie, 1996; IUGS, 2000), and the separate time scales for the Devonian (Tucker et al., 1998), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001) periods. The STD 2002 became the model for the Devonian – Carboniferous – Permian Correlation Chart 2003 (herein abbreviated to DCP 2003, and cited as DCP, 2003 in references), and in consequence, both charts have identical time scales. Version 2 and Version 4g of the DCP 2003 were presented as posters, respectively at the 15th International Congress on Carboniferous and Permian 2003 in Utrecht, and at the 32nd International Geological Congress 2004 in Florence. By August 2004, about 40 authors had contributed to more than 45 supraregional and regional columns, and to over 50 columns with marine and terrestrial biozonations from six continents. Work on several of these columns is still in progress. Both the DCP 2003 and the Explanations 2008 on the Devonian – Carboniferous – Permian Correlation Chart 2003 should be available in 2008. As a major part of the DCP 2003 its Global Stratigraphic Scale (GSS) , the numerical ages of the stages, and (part of) Regional Stratigraphic Scales ( RSS ; supraregional composite sections) of Central and West Europe, East Europe, Tethys, South China (eastern Tethys), and North America are presented here for the first time. It is most important to emphasize there are significant differences between homonymous global and supraregional stratigraphic units which must be taken into account, when time-related geochemical data and events of the DFG-Project 1054 are compared on a global scale. For the East European Platform, the DCP 2003 presents the Resolutions (1990a,b) ( “ Stratigraphic Guide ” of the Soviet Union 1990) for the Carboniferous and Permian periods, and additional terms where necessary. The data used in the DCP 2003 are published or will be published in near future. Arrows indicate questionable ages and positions of stratigraphic boundaries. Frequently used abbreviations are: The numerical calibration of the Global Stratigraphic Scale (GSS) as used in the DCP 2003 is identical to that of the STD 2002. The DCP 2003 time scale is composed of time scales for three periods: Devonian (Weddige et al., 2005), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001). These time scales are numerically calibrated to portray geological sequences and events in a linear time frame. Every endeavour has been made, as far as data allows, to remove unnecessary, artificial compression and expansion of time intervals, biozonations and depositional events. The numerical ages of the stages are mostly rounded to the nearest 0.5 Ma. The accepted and proposed Global Stratotype Section and Points (GSSP) and their index fossils are shown for the DCP 2003. For stages yet to be defined by a GSSP, traditional index fossils are used, as far as possible, in order to establish their boundaries for global correlation. An exception is the base of the Serpukhovian Stage, which is traditionally defined by foraminifers and ammonoids, but when defined by conodonts, lies stratigraphically slightly below the traditional boundary (i.e. the use of the FAD of Lochriea ziegleri leads to the referral of the latest Late “ Visean ” to the Serpukhovian). In the DCP 2003 time scale, as in most time scales, the Middle Devonian is considerably shorter than the Early and the Late Devonian epochs. The Emsian and Famennian are the longest of the Devonian stages. The DCP 2003 age of ca. 358 Ma for the Devonian – Carboniferous boundary (DCB) is a compromise between the ca. 354 Ma age, based on U – Pb SHRIMP dates (Claoué-Long et al., 1993), and the ca. 362 Ma age, based on U – Pb ID-TIMS dates (Tucker et al., 1998). Recently, an age of 360.7 ± 0.7 Ma has been derived for the DCB from U – Pb ID-TIMS dating of two metabentonites (360.2 ± 0.7 Ma for bed 70, Early Siphonodella duplicata Zone, and 360.5 ± 0.8 Ma for bed 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) (Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table 2). The 320 Ma age for the Mississippian – Pennsylvanian boundary is based on the 40 Ar/ 39 Ar date of 324.6 Ma for the Jaklovec Member in Upper Silesia (Lippolt et al., 1984), close to the Pendleian – Arnsbergian boundary (intra-Namurian A) (Menning et al., 2000: Fig. 6). The 40 Ar/ 39 Ar age of 319.5 Ma for the Poruba Member (Lippolt et al., 1984), within the Middle Arnsbergian (Menning et al., 2000: Fig. 6), is indicative of a younger age for this boundary, and a ca. 318 Ma is suggested (this work). The latter age is more consistent with the U – Pb ID-TIMS ages from the Visean of Trapp (pers. com.). The Mississippian/Early Carboniferous (Tournaisian – Serpukhovian) is considerably longer than the Pennsylvanian/Late Carboniferous (Bashkirian – Gzhelian). For a long time the duration of the Pennsylvanian was overestimated, mainly because of the tremendous thicknesses of coal-bearing deposits in the United States and Europe. In the DCP 2003 time scale, the time relationship of Mississippian to Pennsylvanian is 38 my to 24 my. When the boundary ages of ca. 361 Ma (Devonian – Carboniferous), 319 Ma (Mississippian – Pennsylvanian) and 299 Ma (Carboniferous – Permian) are used, the Mississippian (ca. 42 my) is about twice as long as the Pennsylvanian (ca. 20 my). The most significant shift of a numerical age of the Permian is at the top of the Kungurian Stage (top Cisuralian Epoch). The ages suggested for this boundary have swung from 255 – 256 Ma (Harland et al., 1982, 1990; Gradstein and Ogg, 1996) to 274 – 270 Ma (Menning, 1989, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004), despite the upward expansion of the Kungurian of the GSS at the expense of of the Ufimian Stage (cf. 5.1, 5.2.3). The age of the Permian – Triassic boundary (PTB) has been changed from 245 Ma (Harland et al., 1982) to 251 Ma (Menning, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004) based on the 40 Ar/ 39 Ar dating of the Grenzbitumen-Zone of 233 ± 9 Ma (Hellmann and Lippolt, 1981) and the U – Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB (Claoué-Long et al., 1991). Using the U – Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The STD 2002 (Menning and Deutsche Stratigraphische Kommission, 2002)/DCP 2003 time scales differ only slightly from the Geologic Time Scale 2004 (herein abbreviated as GTS 2004, and cited as GTS 2004; Gradstein et al., 2004). However, they all are in marked contrast to those of Haq and Van Eysinga (1987), Geologic Time Scale 1989 (herein abbreviated to GTS 1989, and cited as GTS, 1989 in references), (Harland et al., 1990), Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the more reliable, e.g. for the Sakmarian (6 my in DCP 2003, or 10.2 my in GTS 2004); for the Emsian (15 my in DCP 2003, 9.5 my in GTS 2004, or 17 my in Kaufmann, 2006). Since the publication of the GTS 1989 (Harland et al., 1990) time scale, the most significant changes in age for the Devonian – Permian time interval are for: the Silurian – Devonian boundary from 408 Ma to 416 – 418 Ma, the Dinantian – Silesian (Visean – Namurian) boundary from 332.9 Ma to 326.5 – 325 Ma, the Carboniferous – Permian boundary from 290 Ma to 296 – 299 Ma, the Early – Late/Middle Permian boundary from 256 – 258 Ma to 270 – 274 Ma, and for the Permian – Triassic boundary from 245 Ma to 251 – 252.5 Ma (cf. Fig. 1, Table 2; Menning, 1989). In the DCP 2003 questionable ages and positions of stratigraphic boundaries are marked by arrows, rather than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, however these is not well documented (Fig. 1). For example, comparing ages for the global Emsian – Eifelian boundary (FAD of Eucostapolygnathus partitus ) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma (Kaufmann, 2006) the error bars (± 2 σ ) do overlap hardly (Tables 2 and 3). The Devonian System was introduced by Sedgwick and Murchison (1839). There was, however, no ...

Context 21

... the beginning of Project 1054 of the Deutsche Forschungsgemeinschaft (DFG) “ The evolution of the Late Palaeozoic in the light of sedimentary geochemistry ” German geochemists were using time scales with significant age differences. Therefore, it was decided to use a uniform time scale and a stratigraphic correlation chart with sections from several continents. This was done to ensure variations in isotopic profiles of carbon, oxygen, sulfur, nitrogen, boron, and osmium would be directly comparable, and hence prove to be more reliable indicators of isochronous global events. During 2001 – 2002 the “ Stratigraphische Tabelle von Deutschland 2002 ” (herein abbreviated to STD 2002, and cited as STD, 2002 in references) was prepared by the Deutsche Stratigraphische Kommission (www.stra- tigraphie.de). At that time an updating of the global time scale was necessary in order to integrate isotopic age determinations that had become available over the previous decade since the publication of the previous Phanerozoic time scales (Harland et al., 1990; Odin, 1994; Gradstein and Ogg, 1996; Young and Laurie, 1996; IUGS, 2000), and the separate time scales for the Devonian (Tucker et al., 1998), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001) periods. The STD 2002 became the model for the Devonian – Carboniferous – Permian Correlation Chart 2003 (herein abbreviated to DCP 2003, and cited as DCP, 2003 in references), and in consequence, both charts have identical time scales. Version 2 and Version 4g of the DCP 2003 were presented as posters, respectively at the 15th International Congress on Carboniferous and Permian 2003 in Utrecht, and at the 32nd International Geological Congress 2004 in Florence. By August 2004, about 40 authors had contributed to more than 45 supraregional and regional columns, and to over 50 columns with marine and terrestrial biozonations from six continents. Work on several of these columns is still in progress. Both the DCP 2003 and the Explanations 2008 on the Devonian – Carboniferous – Permian Correlation Chart 2003 should be available in 2008. As a major part of the DCP 2003 its Global Stratigraphic Scale (GSS) , the numerical ages of the stages, and (part of) Regional Stratigraphic Scales ( RSS ; supraregional composite sections) of Central and West Europe, East Europe, Tethys, South China (eastern Tethys), and North America are presented here for the first time. It is most important to emphasize there are significant differences between homonymous global and supraregional stratigraphic units which must be taken into account, when time-related geochemical data and events of the DFG-Project 1054 are compared on a global scale. For the East European Platform, the DCP 2003 presents the Resolutions (1990a,b) ( “ Stratigraphic Guide ” of the Soviet Union 1990) for the Carboniferous and Permian periods, and additional terms where necessary. The data used in the DCP 2003 are published or will be published in near future. Arrows indicate questionable ages and positions of stratigraphic boundaries. Frequently used abbreviations are: The numerical calibration of the Global Stratigraphic Scale (GSS) as used in the DCP 2003 is identical to that of the STD 2002. The DCP 2003 time scale is composed of time scales for three periods: Devonian (Weddige et al., 2005), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001). These time scales are numerically calibrated to portray geological sequences and events in a linear time frame. Every endeavour has been made, as far as data allows, to remove unnecessary, artificial compression and expansion of time intervals, biozonations and depositional events. The numerical ages of the stages are mostly rounded to the nearest 0.5 Ma. The accepted and proposed Global Stratotype Section and Points (GSSP) and their index fossils are shown for the DCP 2003. For stages yet to be defined by a GSSP, traditional index fossils are used, as far as possible, in order to establish their boundaries for global correlation. An exception is the base of the Serpukhovian Stage, which is traditionally defined by foraminifers and ammonoids, but when defined by conodonts, lies stratigraphically slightly below the traditional boundary (i.e. the use of the FAD of Lochriea ziegleri leads to the referral of the latest Late “ Visean ” to the Serpukhovian). In the DCP 2003 time scale, as in most time scales, the Middle Devonian is considerably shorter than the Early and the Late Devonian epochs. The Emsian and Famennian are the longest of the Devonian stages. The DCP 2003 age of ca. 358 Ma for the Devonian – Carboniferous boundary (DCB) is a compromise between the ca. 354 Ma age, based on U – Pb SHRIMP dates (Claoué-Long et al., 1993), and the ca. 362 Ma age, based on U – Pb ID-TIMS dates (Tucker et al., 1998). Recently, an age of 360.7 ± 0.7 Ma has been derived for the DCB from U – Pb ID-TIMS dating of two metabentonites (360.2 ± 0.7 Ma for bed 70, Early Siphonodella duplicata Zone, and 360.5 ± 0.8 Ma for bed 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) (Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table 2). The 320 Ma age for the Mississippian – Pennsylvanian boundary is based on the 40 Ar/ 39 Ar date of 324.6 Ma for the Jaklovec Member in Upper Silesia (Lippolt et al., 1984), close to the Pendleian – Arnsbergian boundary (intra-Namurian A) (Menning et al., 2000: Fig. 6). The 40 Ar/ 39 Ar age of 319.5 Ma for the Poruba Member (Lippolt et al., 1984), within the Middle Arnsbergian (Menning et al., 2000: Fig. 6), is indicative of a younger age for this boundary, and a ca. 318 Ma is suggested (this work). The latter age is more consistent with the U – Pb ID-TIMS ages from the Visean of Trapp (pers. com.). The Mississippian/Early Carboniferous (Tournaisian – Serpukhovian) is considerably longer than the Pennsylvanian/Late Carboniferous (Bashkirian – Gzhelian). For a long time the duration of the Pennsylvanian was overestimated, mainly because of the tremendous thicknesses of coal-bearing deposits in the United States and Europe. In the DCP 2003 time scale, the time relationship of Mississippian to Pennsylvanian is 38 my to 24 my. When the boundary ages of ca. 361 Ma (Devonian – Carboniferous), 319 Ma (Mississippian – Pennsylvanian) and 299 Ma (Carboniferous – Permian) are used, the Mississippian (ca. 42 my) is about twice as long as the Pennsylvanian (ca. 20 my). The most significant shift of a numerical age of the Permian is at the top of the Kungurian Stage (top Cisuralian Epoch). The ages suggested for this boundary have swung from 255 – 256 Ma (Harland et al., 1982, 1990; Gradstein and Ogg, 1996) to 274 – 270 Ma (Menning, 1989, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004), despite the upward expansion of the Kungurian of the GSS at the expense of of the Ufimian Stage (cf. 5.1, 5.2.3). The age of the Permian – Triassic boundary (PTB) has been changed from 245 Ma (Harland et al., 1982) to 251 Ma (Menning, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004) based on the 40 Ar/ 39 Ar dating of the Grenzbitumen-Zone of 233 ± 9 Ma (Hellmann and Lippolt, 1981) and the U – Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB (Claoué-Long et al., 1991). Using the U – Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The STD 2002 (Menning and Deutsche Stratigraphische Kommission, 2002)/DCP 2003 time scales differ only slightly from the Geologic Time Scale 2004 (herein abbreviated as GTS 2004, and cited as GTS 2004; Gradstein et al., 2004). However, they all are in marked contrast to those of Haq and Van Eysinga (1987), Geologic Time Scale 1989 (herein abbreviated to GTS 1989, and cited as GTS, 1989 in references), (Harland et al., 1990), Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the more reliable, e.g. for the Sakmarian (6 my in DCP 2003, or 10.2 my in GTS 2004); for the Emsian (15 my in DCP 2003, 9.5 my in GTS 2004, or 17 my in Kaufmann, 2006). Since the publication of the GTS 1989 (Harland et al., 1990) time scale, the most significant changes in age for the Devonian – Permian time interval are for: the Silurian – Devonian boundary from 408 Ma to 416 – 418 Ma, the Dinantian – Silesian (Visean – Namurian) boundary from 332.9 Ma to 326.5 – 325 Ma, the Carboniferous – Permian boundary from 290 Ma to 296 – 299 Ma, the Early – Late/Middle Permian boundary from 256 – 258 Ma to 270 – 274 Ma, and for the Permian – Triassic boundary from 245 Ma to 251 – 252.5 Ma (cf. Fig. 1, Table 2; Menning, 1989). In the DCP 2003 questionable ages and positions of stratigraphic boundaries are marked by arrows, rather than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, however these is not well documented (Fig. 1). For example, comparing ages for the global Emsian – Eifelian boundary (FAD of Eucostapolygnathus partitus ) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma (Kaufmann, 2006) the error bars (± 2 σ ) do overlap hardly (Tables 2 and 3). The Devonian System was introduced by Sedgwick and Murchison (1839). There was, however, no ...

Context 22

... the beginning of Project 1054 of the Deutsche Forschungsgemeinschaft (DFG) “ The evolution of the Late Palaeozoic in the light of sedimentary geochemistry ” German geochemists were using time scales with significant age differences. Therefore, it was decided to use a uniform time scale and a stratigraphic correlation chart with sections from several continents. This was done to ensure variations in isotopic profiles of carbon, oxygen, sulfur, nitrogen, boron, and osmium would be directly comparable, and hence prove to be more reliable indicators of isochronous global events. During 2001 – 2002 the “ Stratigraphische Tabelle von Deutschland 2002 ” (herein abbreviated to STD 2002, and cited as STD, 2002 in references) was prepared by the Deutsche Stratigraphische Kommission (www.stra- tigraphie.de). At that time an updating of the global time scale was necessary in order to integrate isotopic age determinations that had become available over the previous decade since the publication of the previous Phanerozoic time scales (Harland et al., 1990; Odin, 1994; Gradstein and Ogg, 1996; Young and Laurie, 1996; IUGS, 2000), and the separate time scales for the Devonian (Tucker et al., 1998), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001) periods. The STD 2002 became the model for the Devonian – Carboniferous – Permian Correlation Chart 2003 (herein abbreviated to DCP 2003, and cited as DCP, 2003 in references), and in consequence, both charts have identical time scales. Version 2 and Version 4g of the DCP 2003 were presented as posters, respectively at the 15th International Congress on Carboniferous and Permian 2003 in Utrecht, and at the 32nd International Geological Congress 2004 in Florence. By August 2004, about 40 authors had contributed to more than 45 supraregional and regional columns, and to over 50 columns with marine and terrestrial biozonations from six continents. Work on several of these columns is still in progress. Both the DCP 2003 and the Explanations 2008 on the Devonian – Carboniferous – Permian Correlation Chart 2003 should be available in 2008. As a major part of the DCP 2003 its Global Stratigraphic Scale (GSS) , the numerical ages of the stages, and (part of) Regional Stratigraphic Scales ( RSS ; supraregional composite sections) of Central and West Europe, East Europe, Tethys, South China (eastern Tethys), and North America are presented here for the first time. It is most important to emphasize there are significant differences between homonymous global and supraregional stratigraphic units which must be taken into account, when time-related geochemical data and events of the DFG-Project 1054 are compared on a global scale. For the East European Platform, the DCP 2003 presents the Resolutions (1990a,b) ( “ Stratigraphic Guide ” of the Soviet Union 1990) for the Carboniferous and Permian periods, and additional terms where necessary. The data used in the DCP 2003 are published or will be published in near future. Arrows indicate questionable ages and positions of stratigraphic boundaries. Frequently used abbreviations are: The numerical calibration of the Global Stratigraphic Scale (GSS) as used in the DCP 2003 is identical to that of the STD 2002. The DCP 2003 time scale is composed of time scales for three periods: Devonian (Weddige et al., 2005), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001). These time scales are numerically calibrated to portray geological sequences and events in a linear time frame. Every endeavour has been made, as far as data allows, to remove unnecessary, artificial compression and expansion of time intervals, biozonations and depositional events. The numerical ages of the stages are mostly rounded to the nearest 0.5 Ma. The accepted and proposed Global Stratotype Section and Points (GSSP) and their index fossils are shown for the DCP 2003. For stages yet to be defined by a GSSP, traditional index fossils are used, as far as possible, in order to establish their boundaries for global correlation. An exception is the base of the Serpukhovian Stage, which is traditionally defined by foraminifers and ammonoids, but when defined by conodonts, lies stratigraphically slightly below the traditional boundary (i.e. the use of the FAD of Lochriea ziegleri leads to the referral of the latest Late “ Visean ” to the Serpukhovian). In the DCP 2003 time scale, as in most time scales, the Middle Devonian is considerably shorter than the Early and the Late Devonian epochs. The Emsian and Famennian are the longest of the Devonian stages. The DCP 2003 age of ca. 358 Ma for the Devonian – Carboniferous boundary (DCB) is a compromise between the ca. 354 Ma age, based on U – Pb SHRIMP dates (Claoué-Long et al., 1993), and the ca. 362 Ma age, based on U – Pb ID-TIMS dates (Tucker et al., 1998). Recently, an age of 360.7 ± 0.7 Ma has been derived for the DCB from U – Pb ID-TIMS dating of two metabentonites (360.2 ± 0.7 Ma for bed 70, Early Siphonodella duplicata Zone, and 360.5 ± 0.8 Ma for bed 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) (Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table 2). The 320 Ma age for the Mississippian – Pennsylvanian boundary is based on the 40 Ar/ 39 Ar date of 324.6 Ma for the Jaklovec Member in Upper Silesia (Lippolt et al., 1984), close to the Pendleian – Arnsbergian boundary (intra-Namurian A) (Menning et al., 2000: Fig. 6). The 40 Ar/ 39 Ar age of 319.5 Ma for the Poruba Member (Lippolt et al., 1984), within the Middle Arnsbergian (Menning et al., 2000: Fig. 6), is indicative of a younger age for this boundary, and a ca. 318 Ma is suggested (this work). The latter age is more consistent with the U – Pb ID-TIMS ages from the Visean of Trapp (pers. com.). The Mississippian/Early Carboniferous (Tournaisian – Serpukhovian) is considerably longer than the Pennsylvanian/Late Carboniferous (Bashkirian – Gzhelian). For a long time the duration of the Pennsylvanian was overestimated, mainly because of the tremendous thicknesses of coal-bearing deposits in the United States and Europe. In the DCP 2003 time scale, the time relationship of Mississippian to Pennsylvanian is 38 my to 24 my. When the boundary ages of ca. 361 Ma (Devonian – Carboniferous), 319 Ma (Mississippian – Pennsylvanian) and 299 Ma (Carboniferous – Permian) are used, the Mississippian (ca. 42 my) is about twice as long as the Pennsylvanian (ca. 20 my). The most significant shift of a numerical age of the Permian is at the top of the Kungurian Stage (top Cisuralian Epoch). The ages suggested for this boundary have swung from 255 – 256 Ma (Harland et al., 1982, 1990; Gradstein and Ogg, 1996) to 274 – 270 Ma (Menning, 1989, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004), despite the upward expansion of the Kungurian of the GSS at the expense of of the Ufimian Stage (cf. 5.1, 5.2.3). The age of the Permian – Triassic boundary (PTB) has been changed from 245 Ma (Harland et al., 1982) to 251 Ma (Menning, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004) based on the 40 Ar/ 39 Ar dating of the Grenzbitumen-Zone of 233 ± 9 Ma (Hellmann and Lippolt, 1981) and the U – Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB (Claoué-Long et al., 1991). Using the U – Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The STD 2002 (Menning and Deutsche Stratigraphische Kommission, 2002)/DCP 2003 time scales differ only slightly from the Geologic Time Scale 2004 (herein abbreviated as GTS 2004, and cited as GTS 2004; Gradstein et al., 2004). However, they all are in marked contrast to those of Haq and Van Eysinga (1987), Geologic Time Scale 1989 (herein abbreviated to GTS 1989, and cited as GTS, 1989 in references), (Harland et al., 1990), Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the more reliable, e.g. for the Sakmarian (6 my in DCP 2003, or 10.2 my in GTS 2004); for the Emsian (15 my in DCP 2003, 9.5 my in GTS 2004, or 17 my in Kaufmann, 2006). Since the publication of the GTS 1989 (Harland et al., 1990) time scale, the most significant changes in age for the Devonian – Permian time interval are for: the Silurian – Devonian boundary from 408 Ma to 416 – 418 Ma, the Dinantian – Silesian (Visean – Namurian) boundary from 332.9 Ma to 326.5 – 325 Ma, the Carboniferous – Permian boundary from 290 Ma to 296 – 299 Ma, the Early – Late/Middle Permian boundary from 256 – 258 Ma to 270 – 274 Ma, and for the Permian – Triassic boundary from 245 Ma to 251 – 252.5 Ma (cf. Fig. 1, Table 2; Menning, 1989). In the DCP 2003 questionable ages and positions of stratigraphic boundaries are marked by arrows, rather than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, however these is not well documented (Fig. 1). For example, comparing ages for the global Emsian – Eifelian boundary (FAD of Eucostapolygnathus partitus ) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma (Kaufmann, 2006) the error bars (± 2 σ ) do overlap hardly (Tables 2 and 3). The Devonian System was introduced by Sedgwick and Murchison (1839). There was, however, no ...

Context 23

... the beginning of Project 1054 of the Deutsche Forschungsgemeinschaft (DFG) “ The evolution of the Late Palaeozoic in the light of sedimentary geochemistry ” German geochemists were using time scales with significant age differences. Therefore, it was decided to use a uniform time scale and a stratigraphic correlation chart with sections from several continents. This was done to ensure variations in isotopic profiles of carbon, oxygen, sulfur, nitrogen, boron, and osmium would be directly comparable, and hence prove to be more reliable indicators of isochronous global events. During 2001 – 2002 the “ Stratigraphische Tabelle von Deutschland 2002 ” (herein abbreviated to STD 2002, and cited as STD, 2002 in references) was prepared by the Deutsche Stratigraphische Kommission (www.stra- tigraphie.de). At that time an updating of the global time scale was necessary in order to integrate isotopic age determinations that had become available over the previous decade since the publication of the previous Phanerozoic time scales (Harland et al., 1990; Odin, 1994; Gradstein and Ogg, 1996; Young and Laurie, 1996; IUGS, 2000), and the separate time scales for the Devonian (Tucker et al., 1998), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001) periods. The STD 2002 became the model for the Devonian – Carboniferous – Permian Correlation Chart 2003 (herein abbreviated to DCP 2003, and cited as DCP, 2003 in references), and in consequence, both charts have identical time scales. Version 2 and Version 4g of the DCP 2003 were presented as posters, respectively at the 15th International Congress on Carboniferous and Permian 2003 in Utrecht, and at the 32nd International Geological Congress 2004 in Florence. By August 2004, about 40 authors had contributed to more than 45 supraregional and regional columns, and to over 50 columns with marine and terrestrial biozonations from six continents. Work on several of these columns is still in progress. Both the DCP 2003 and the Explanations 2008 on the Devonian – Carboniferous – Permian Correlation Chart 2003 should be available in 2008. As a major part of the DCP 2003 its Global Stratigraphic Scale (GSS) , the numerical ages of the stages, and (part of) Regional Stratigraphic Scales ( RSS ; supraregional composite sections) of Central and West Europe, East Europe, Tethys, South China (eastern Tethys), and North America are presented here for the first time. It is most important to emphasize there are significant differences between homonymous global and supraregional stratigraphic units which must be taken into account, when time-related geochemical data and events of the DFG-Project 1054 are compared on a global scale. For the East European Platform, the DCP 2003 presents the Resolutions (1990a,b) ( “ Stratigraphic Guide ” of the Soviet Union 1990) for the Carboniferous and Permian periods, and additional terms where necessary. The data used in the DCP 2003 are published or will be published in near future. Arrows indicate questionable ages and positions of stratigraphic boundaries. Frequently used abbreviations are: The numerical calibration of the Global Stratigraphic Scale (GSS) as used in the DCP 2003 is identical to that of the STD 2002. The DCP 2003 time scale is composed of time scales for three periods: Devonian (Weddige et al., 2005), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001). These time scales are numerically calibrated to portray geological sequences and events in a linear time frame. Every endeavour has been made, as far as data allows, to remove unnecessary, artificial compression and expansion of time intervals, biozonations and depositional events. The numerical ages of the stages are mostly rounded to the nearest 0.5 Ma. The accepted and proposed Global Stratotype Section and Points (GSSP) and their index fossils are shown for the DCP 2003. For stages yet to be defined by a GSSP, traditional index fossils are used, as far as possible, in order to establish their boundaries for global correlation. An exception is the base of the Serpukhovian Stage, which is traditionally defined by foraminifers and ammonoids, but when defined by conodonts, lies stratigraphically slightly below the traditional boundary (i.e. the use of the FAD of Lochriea ziegleri leads to the referral of the latest Late “ Visean ” to the Serpukhovian). In the DCP 2003 time scale, as in most time scales, the Middle Devonian is considerably shorter than the Early and the Late Devonian epochs. The Emsian and Famennian are the longest of the Devonian stages. The DCP 2003 age of ca. 358 Ma for the Devonian – Carboniferous boundary (DCB) is a compromise between the ca. 354 Ma age, based on U – Pb SHRIMP dates (Claoué-Long et al., 1993), and the ca. 362 Ma age, based on U – Pb ID-TIMS dates (Tucker et al., 1998). Recently, an age of 360.7 ± 0.7 Ma has been derived for the DCB from U – Pb ID-TIMS dating of two metabentonites (360.2 ± 0.7 Ma for bed 70, Early Siphonodella duplicata Zone, and 360.5 ± 0.8 Ma for bed 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) (Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table 2). The 320 Ma age for the Mississippian – Pennsylvanian boundary is based on the 40 Ar/ 39 Ar date of 324.6 Ma for the Jaklovec Member in Upper Silesia (Lippolt et al., 1984), close to the Pendleian – Arnsbergian boundary (intra-Namurian A) (Menning et al., 2000: Fig. 6). The 40 Ar/ 39 Ar age of 319.5 Ma for the Poruba Member (Lippolt et al., 1984), within the Middle Arnsbergian (Menning et al., 2000: Fig. 6), is indicative of a younger age for this boundary, and a ca. 318 Ma is suggested (this work). The latter age is more consistent with the U – Pb ID-TIMS ages from the Visean of Trapp (pers. com.). The Mississippian/Early Carboniferous (Tournaisian – Serpukhovian) is considerably longer than the Pennsylvanian/Late Carboniferous (Bashkirian – Gzhelian). For a long time the duration of the Pennsylvanian was overestimated, mainly because of the tremendous thicknesses of coal-bearing deposits in the United States and Europe. In the DCP 2003 time scale, the time relationship of Mississippian to Pennsylvanian is 38 my to 24 my. When the boundary ages of ca. 361 Ma (Devonian – Carboniferous), 319 Ma (Mississippian – Pennsylvanian) and 299 Ma (Carboniferous – Permian) are used, the Mississippian (ca. 42 my) is about twice as long as the Pennsylvanian (ca. 20 my). The most significant shift of a numerical age of the Permian is at the top of the Kungurian Stage (top Cisuralian Epoch). The ages suggested for this boundary have swung from 255 – 256 Ma (Harland et al., 1982, 1990; Gradstein and Ogg, 1996) to 274 – 270 Ma (Menning, 1989, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004), despite the upward expansion of the Kungurian of the GSS at the expense of of the Ufimian Stage (cf. 5.1, 5.2.3). The age of the Permian – Triassic boundary (PTB) has been changed from 245 Ma (Harland et al., 1982) to 251 Ma (Menning, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004) based on the 40 Ar/ 39 Ar dating of the Grenzbitumen-Zone of 233 ± 9 Ma (Hellmann and Lippolt, 1981) and the U – Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB (Claoué-Long et al., 1991). Using the U – Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The STD 2002 (Menning and Deutsche Stratigraphische Kommission, 2002)/DCP 2003 time scales differ only slightly from the Geologic Time Scale 2004 (herein abbreviated as GTS 2004, and cited as GTS 2004; Gradstein et al., 2004). However, they all are in marked contrast to those of Haq and Van Eysinga (1987), Geologic Time Scale 1989 (herein abbreviated to GTS 1989, and cited as GTS, 1989 in references), (Harland et al., 1990), Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the more reliable, e.g. for the Sakmarian (6 my in DCP 2003, or 10.2 my in GTS 2004); for the Emsian (15 my in DCP 2003, 9.5 my in GTS 2004, or 17 my in Kaufmann, 2006). Since the publication of the GTS 1989 (Harland et al., 1990) time scale, the most significant changes in age for the Devonian – Permian time interval are for: the Silurian – Devonian boundary from 408 Ma to 416 – 418 Ma, the Dinantian – Silesian (Visean – Namurian) boundary from 332.9 Ma to 326.5 – 325 Ma, the Carboniferous – Permian boundary from 290 Ma to 296 – 299 Ma, the Early – Late/Middle Permian boundary from 256 – 258 Ma to 270 – 274 Ma, and for the Permian – Triassic boundary from 245 Ma to 251 – 252.5 Ma (cf. Fig. 1, Table 2; Menning, 1989). In the DCP 2003 questionable ages and positions of stratigraphic boundaries are marked by arrows, rather than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, however these is not well documented (Fig. 1). For example, comparing ages for the global Emsian – Eifelian boundary (FAD of Eucostapolygnathus partitus ) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma (Kaufmann, 2006) the error bars (± 2 σ ) do overlap hardly (Tables 2 and 3). The Devonian System was introduced by Sedgwick and Murchison (1839). There was, however, no ...

Context 24

... the beginning of Project 1054 of the Deutsche Forschungsgemeinschaft (DFG) “ The evolution of the Late Palaeozoic in the light of sedimentary geochemistry ” German geochemists were using time scales with significant age differences. Therefore, it was decided to use a uniform time scale and a stratigraphic correlation chart with sections from several continents. This was done to ensure variations in isotopic profiles of carbon, oxygen, sulfur, nitrogen, boron, and osmium would be directly comparable, and hence prove to be more reliable indicators of isochronous global events. During 2001 – 2002 the “ Stratigraphische Tabelle von Deutschland 2002 ” (herein abbreviated to STD 2002, and cited as STD, 2002 in references) was prepared by the Deutsche Stratigraphische Kommission (www.stra- tigraphie.de). At that time an updating of the global time scale was necessary in order to integrate isotopic age determinations that had become available over the previous decade since the publication of the previous Phanerozoic time scales (Harland et al., 1990; Odin, 1994; Gradstein and Ogg, 1996; Young and Laurie, 1996; IUGS, 2000), and the separate time scales for the Devonian (Tucker et al., 1998), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001) periods. The STD 2002 became the model for the Devonian – Carboniferous – Permian Correlation Chart 2003 (herein abbreviated to DCP 2003, and cited as DCP, 2003 in references), and in consequence, both charts have identical time scales. Version 2 and Version 4g of the DCP 2003 were presented as posters, respectively at the 15th International Congress on Carboniferous and Permian 2003 in Utrecht, and at the 32nd International Geological Congress 2004 in Florence. By August 2004, about 40 authors had contributed to more than 45 supraregional and regional columns, and to over 50 columns with marine and terrestrial biozonations from six continents. Work on several of these columns is still in progress. Both the DCP 2003 and the Explanations 2008 on the Devonian – Carboniferous – Permian Correlation Chart 2003 should be available in 2008. As a major part of the DCP 2003 its Global Stratigraphic Scale (GSS) , the numerical ages of the stages, and (part of) Regional Stratigraphic Scales ( RSS ; supraregional composite sections) of Central and West Europe, East Europe, Tethys, South China (eastern Tethys), and North America are presented here for the first time. It is most important to emphasize there are significant differences between homonymous global and supraregional stratigraphic units which must be taken into account, when time-related geochemical data and events of the DFG-Project 1054 are compared on a global scale. For the East European Platform, the DCP 2003 presents the Resolutions (1990a,b) ( “ Stratigraphic Guide ” of the Soviet Union 1990) for the Carboniferous and Permian periods, and additional terms where necessary. The data used in the DCP 2003 are published or will be published in near future. Arrows indicate questionable ages and positions of stratigraphic boundaries. Frequently used abbreviations are: The numerical calibration of the Global Stratigraphic Scale (GSS) as used in the DCP 2003 is identical to that of the STD 2002. The DCP 2003 time scale is composed of time scales for three periods: Devonian (Weddige et al., 2005), Carboniferous (Menning et al., 2000), and Permian (Menning, 2001). These time scales are numerically calibrated to portray geological sequences and events in a linear time frame. Every endeavour has been made, as far as data allows, to remove unnecessary, artificial compression and expansion of time intervals, biozonations and depositional events. The numerical ages of the stages are mostly rounded to the nearest 0.5 Ma. The accepted and proposed Global Stratotype Section and Points (GSSP) and their index fossils are shown for the DCP 2003. For stages yet to be defined by a GSSP, traditional index fossils are used, as far as possible, in order to establish their boundaries for global correlation. An exception is the base of the Serpukhovian Stage, which is traditionally defined by foraminifers and ammonoids, but when defined by conodonts, lies stratigraphically slightly below the traditional boundary (i.e. the use of the FAD of Lochriea ziegleri leads to the referral of the latest Late “ Visean ” to the Serpukhovian). In the DCP 2003 time scale, as in most time scales, the Middle Devonian is considerably shorter than the Early and the Late Devonian epochs. The Emsian and Famennian are the longest of the Devonian stages. The DCP 2003 age of ca. 358 Ma for the Devonian – Carboniferous boundary (DCB) is a compromise between the ca. 354 Ma age, based on U – Pb SHRIMP dates (Claoué-Long et al., 1993), and the ca. 362 Ma age, based on U – Pb ID-TIMS dates (Tucker et al., 1998). Recently, an age of 360.7 ± 0.7 Ma has been derived for the DCB from U – Pb ID-TIMS dating of two metabentonites (360.2 ± 0.7 Ma for bed 70, Early Siphonodella duplicata Zone, and 360.5 ± 0.8 Ma for bed 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) (Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table 2). The 320 Ma age for the Mississippian – Pennsylvanian boundary is based on the 40 Ar/ 39 Ar date of 324.6 Ma for the Jaklovec Member in Upper Silesia (Lippolt et al., 1984), close to the Pendleian – Arnsbergian boundary (intra-Namurian A) (Menning et al., 2000: Fig. 6). The 40 Ar/ 39 Ar age of 319.5 Ma for the Poruba Member (Lippolt et al., 1984), within the Middle Arnsbergian (Menning et al., 2000: Fig. 6), is indicative of a younger age for this boundary, and a ca. 318 Ma is suggested (this work). The latter age is more consistent with the U – Pb ID-TIMS ages from the Visean of Trapp (pers. com.). The Mississippian/Early Carboniferous (Tournaisian – Serpukhovian) is considerably longer than the Pennsylvanian/Late Carboniferous (Bashkirian – Gzhelian). For a long time the duration of the Pennsylvanian was overestimated, mainly because of the tremendous thicknesses of coal-bearing deposits in the United States and Europe. In the DCP 2003 time scale, the time relationship of Mississippian to Pennsylvanian is 38 my to 24 my. When the boundary ages of ca. 361 Ma (Devonian – Carboniferous), 319 Ma (Mississippian – Pennsylvanian) and 299 Ma (Carboniferous – Permian) are used, the Mississippian (ca. 42 my) is about twice as long as the Pennsylvanian (ca. 20 my). The most significant shift of a numerical age of the Permian is at the top of the Kungurian Stage (top Cisuralian Epoch). The ages suggested for this boundary have swung from 255 – 256 Ma (Harland et al., 1982, 1990; Gradstein and Ogg, 1996) to 274 – 270 Ma (Menning, 1989, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004), despite the upward expansion of the Kungurian of the GSS at the expense of of the Ufimian Stage (cf. 5.1, 5.2.3). The age of the Permian – Triassic boundary (PTB) has been changed from 245 Ma (Harland et al., 1982) to 251 Ma (Menning, 1995a,b; Young and Laurie, 1996; Gradstein et al., 2004) based on the 40 Ar/ 39 Ar dating of the Grenzbitumen-Zone of 233 ± 9 Ma (Hellmann and Lippolt, 1981) and the U – Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB (Claoué-Long et al., 1991). Using the U – Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan (Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB (Fig. 1; Menning et al., 2005). The STD 2002 (Menning and Deutsche Stratigraphische Kommission, 2002)/DCP 2003 time scales differ only slightly from the Geologic Time Scale 2004 (herein abbreviated as GTS 2004, and cited as GTS 2004; Gradstein et al., 2004). However, they all are in marked contrast to those of Haq and Van Eysinga (1987), Geologic Time Scale 1989 (herein abbreviated to GTS 1989, and cited as GTS, 1989 in references), (Harland et al., 1990), Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the more reliable, e.g. for the Sakmarian (6 my in DCP 2003, or 10.2 my in GTS 2004); for the Emsian (15 my in DCP 2003, 9.5 my in GTS 2004, or 17 my in Kaufmann, 2006). Since the publication of the GTS 1989 (Harland et al., 1990) time scale, the most significant changes in age for the Devonian – Permian time interval are for: the Silurian – Devonian boundary from 408 Ma to 416 – 418 Ma, the Dinantian – Silesian (Visean – Namurian) boundary from 332.9 Ma to 326.5 – 325 Ma, the Carboniferous – Permian boundary from 290 Ma to 296 – 299 Ma, the Early – Late/Middle Permian boundary from 256 – 258 Ma to 270 – 274 Ma, and for the Permian – Triassic boundary from 245 Ma to 251 – 252.5 Ma (cf. Fig. 1, Table 2; Menning, 1989). In the DCP 2003 questionable ages and positions of stratigraphic boundaries are marked by arrows, rather than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, however these is not well documented (Fig. 1). For example, comparing ages for the global Emsian – Eifelian boundary (FAD of Eucostapolygnathus partitus ) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma (Kaufmann, 2006) the error bars (± 2 σ ) do overlap hardly (Tables 2 and 3). The Devonian System was introduced by Sedgwick and Murchison (1839). There was, however, no ...

Context 25

... 79, late part of the Siphonodella sulcata Zone) taken from the Hasselbach auxiliary global stratotype section in the Rhenish Slate Mts. (Rheinisches Schiefergebirge) ( Trapp et al., 2004). Using the probable FAD of S. sulcata in bed 84 of this section, and its thickness, the age of the DCB is estimated at ca. 361.4 ± 0.7 Ma (Menning, this work; Fig. 1, Table ...

Context 26

... of 233 ± 9 Ma ( Hellmann and Lippolt, 1981) and the U-Pb SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan which is only slightly older than the PTB ( Claoué-Long et al., 1991). Using the U-Pb ID-TIMS age of 252.6 ± 0.2 Ma for bed 25 in Meishan ( Mundil et al., 2004), an age of 252.5 Ma has been derived for the PTB ( Fig. 1; Menning et al., ...

Context 27

... abbreviated to GTS 1989, and cited as GTS, 1989 in references), , Odin (1994), and Gradstein and Ogg (1996). Comparing the DCP 2003 and the GTS 2004, the maximum age differences of stage boundaries are 4.6 Ma in the Permian, 1.9 Ma in the Carboniferous and 7.6 Ma in the Devonian. It is difficult to assess which of the stage durations (Table 2, Fig. 1) are the Table 2; Menning, ...

Context 28

... than by error bars. The ages of the DCP 2003 are mostly rounded to the nearest 0.5 Ma in order to avoid estimates of questionable accuracy, whereas ages of 0.1 Ma in the GTS 2004 and their small error bars of ± 0.4 Ma to ± 2.8 Ma for the ages of the Devonian to Permian stages suggest an improved accuracy, howev- er these is not well documented (Fig. 1). For example, comparing ages for the global Emsian-Eifelian boundary (FAD of Eucostapolygnathus partitus) of 397.5 ± 2.7 Ma (GTS 2004) and 391.9 ± 3.4 Ma ( Kaufmann, 2006) the error bars (± 2σ) do overlap hardly (Tables 2 and ...

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... Loch- kovian; Kaufmann, 2006). Using the same ages as Tucker et al. (1998), and part of their error bars, an age of 417.5 Ma has been derived for the DCP 2003. Kaufmann (2006 favours an age of 418.1 ± 3.0 Ma which may have been influenced at least from the Early Emsian Bundenbach ID-TIMS age of 407.7 ± 0.7 Ma of Kaufmann et al. (2005) (cf. Fig. 1). House and Gradstein (2004: Fig. 14.5;GTS 2004) estimated an age of 416 Ma using the Ludfordian Upper Whitcliffe ID- TIMS age of 420.2 ± 3.9 Ma ( Tucker et al., 1998) as tie point, giving the other ages a lower weight. Compston (2000) suggested a SHRIMP age of 409.9 ± 1.1 Ma for the base of the Devonian, after he reprocessed original ...

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... the other ages a lower weight. Compston (2000) suggested a SHRIMP age of 409.9 ± 1.1 Ma for the base of the Devonian, after he reprocessed original data, allowing for the heterogeneity in the zircon SL13 standard. His entire Devonian time scale, except for the upper limit, is in strong contrast to the other scales based mainly on ID-TIMS ages (Fig. ...

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... appears to be flawed by the exclusion of some very significant ID-TIMS ages in its calibration procedure. Thus, the age of the upper boundary of the Emsian Stage at 397.5 ± 2.7 Ma may be too old and its duration of 9.5 Ma is probably too short. Most probably, the Emsian Stage is longer than the combined ages of the Lochkovian and Pragian stages (Fig. ...

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... the Frasnian Stage, not least because of its division into a greater number of conodont and ammonoid zones. However, this time-relation has yet to be confirmed by a reliable isotopic age determination. Only in the time scale of Compston (2000), using SHRIMP and selected/ reinterpreted MSID ages, is the Frasnian longer than the Famennian Stage (Fig. 1). Since the work of Tucker et al. (1998) the Famennian has a duration of about 15 my (Table 2) and, therefore is one of the longest Phanerozoic ...

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... numerical age for the base of the Tournaisian Stage (base of the Carboniferous Period) is between 354 Ma (Young and Laurie, 1996) and 362 Ma (Tucker et al., 1998) (Fig. 1). The 354 Ma age was derived from an Early Carboniferous SHRIMP age , which is now known to be too young because of the inherent problem in the SL13 zircon standard used for the determination. Nevertheless, the 354 Ma age was applied as a tie point in the Phanerozoic time scales of Young and Laurie (1996) and Gradstein and Ogg (1996), ...

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... , which is now known to be too young because of the inherent problem in the SL13 zircon standard used for the determination. Nevertheless, the 354 Ma age was applied as a tie point in the Phanerozoic time scales of Young and Laurie (1996) and Gradstein and Ogg (1996), and in the Carboniferous time scales of Jones (1995) and Menning et al. (2000) (Fig. 1), not least, because the dated sample was collected from the S. sulcata conodont Zone in the well known Hasselbach section in the Rhenish Slate ...

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... renormalized by Claoué-Long et al., 1995) are both from the same tuff in bed 79. Using the ages of 360.5 ± 0.8 Ma and 360.2 ± 0.7 Ma of Trapp et al. (2004) and the thicknesses of the Hasselbach section instead of the thicknesses of the Lali section (South China, Trapp et al., 2004), the age of the DCB is estimated at 361.4 Ma (Menning, this work, Fig. 1). A Re-Os age of 361.3 ± 2.4 Ma of a black shale from the Exshaw Formation, Canada (Selby and Creaser, 2005) is very close to ...

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... numerical age for the base of the Bashkirian Stage is between 323 Ma and ca. 318 Ma (Fig. 1) excluding the age of 314 Ma which was derived from the U-Pb SHRIMP age of 314.45 ± 3.3 Ma from the Harewood borehole, middle Arnsbergian, middle Namurian A ( Riley et al., 1993;cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of ...

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... al., 1993;cf. Menning et al., 2000). In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member ( Lippolt et al., 1984). This is allocated to approximately the Pendleian-Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. ...

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... In the DCP 2003/STD 2002 an age of ca. 320 Ma is used, which has been derived by Menning et al. (2000) using the 40 Ar/ 39 Ar age of 324.6 ± 8.0 Ma from the Jaklovec Member ( Lippolt et al., 1984). This is allocated to approximately the Pendleian-Arnsbergian boundary (Weyer, in Menning et al., 2000: Fig. 6). Fig. 15.5) in the GTS 2004 (Table 2, Fig. 1), using both the above 40 Ar/ 39 Ar ages of Lippolt et al. (1984) and a SHRIMP age (using the SL 13 zircon standard; Riley et al., 1993) which was rejected by Menning et al. ...

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... of 314 to 311 Ma are estimated for the base of the Moscovian Stage (Fig. ...

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... of 306.5 ± 1.0 Ma to 303 Ma are estimated for the base of the Kasimovian Stage (Fig. 1). They are based mainly on 40 Ar/ 39 Ar ages from coal tonsteins of Westphalian and Stephanian age of Central Europe ( Lippolt et al., 1984;Burger et al., 1997). In the DCP 2003/STD 2002 an age of 305 Ma is used, which is an optimal estimate using the above-mentioned dates. The age of 306.5 ± 1.0 Ma ( Davydov et al. in the GTS, 2004) is ...

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... of ca. 295 to ca. 304 Ma are estimated for the base of the Gzhelian Stage (Table 2, Fig. 1). The main reason for this relatively wide age variation is the poor biostratigraphic correlation between the southern Ural Mts. ...

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... (Early Epoch and Late Epoch). In East Europe the Early Permian consists of the Asselian, Sakmarian, Artinskian, and Kungurian stages and the Late Permian, starting at the base of the Solikamskian Horizon, includes the Ufimian, Kazanian, and Tatarian stages (Resolutions, 1965). These seven (supra)regional stages were also used as global stages (Fig. 1), although the base of the Ufimian was not well defined in large areas of the East European Platform ( Grunt, ...

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... ages of 290 to 299 ± 1.0 Ma are given for the base of the Asselian Stage (Fig. 1). An age of 296 Ma (Menning, 1989;ff.) is used in the DCP 2003/ STD 2002, which is derived from ages from Central Europe ( Lippolt and Hess, 1983;Lippolt et al., 1984;. This age range encompasses those previously suggested viz., ca. 292 Ma ( Chuvashov et al., 1996, SHRIMP;cf. Menning et al., 2000: Fig. 7), 298 Ma (Jones, 1995, Central ...

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... my of the DCP 2003 or the 8.8 ± 0.7 my of the GTS 2004 (Wardlaw et al.)], and for the Kungurian Stage [the ca. 7 my of the DCP 2003 or the 5.0 ± 0.7 my of the GTS 2004 (Wardlaw et al.) (Table 2, Fig. 1). ...

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... Late Permian of East Europe consisted for 40 years of the three regional stages Ufimian, Kazanian, and Tatarian (Resolutions, 1965), but these names were also used for stages of the GSS (Fig. 1). Their stratotype and reference sections are described in detail in a monograph by Esaulova et al. (1998). The introduction of the GSSP concept resulted in the controversial subdivision of the Late Permian into the Middle and Late Permian, the substitution of the East European stages by North American and Chinese ones including new ...

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... magnetostratigraphic evidence permits a corre- lation of the marine parts of the Kazanian and Urzhumian successions of East Europe with parts of the entirely terrestrial Rotliegend of Central and West Europe ( Fig. 4; Menning, 2001), and a correlation of the marine early Zechstein of Central and West Europe with the entirely terrestrial late Tatarian (Vyatkian) deposits of East Europe. Consequently, marine fossils cannot be used to demonstrate that the successions in Central and West Europe are approximately isochronous with those of East Europe. Thus, marine ...

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... which is based on a) the 265.3 ± 0.2 Ma ID-TIMS age (Bowring et al., 1998) for a bentonite bed just below the base of the Capitanian Stage in its boundary stratotype at the Nipple Hill near the Guadalupe Mts., West Texas, b) the numerous ID- TIMS ages from the sections Meishan and Shangsi around the PTB ( Bowring et al., 1998;cf. Menning, 2001: Fig. 1) as well as the SHRIMP age of 251.2 ± 3.4 Ma for bed 25 in the GSSP of Meishan (ClaouéLong et al., 1991) which is only slightly older than the PTB in Bed 27c ( Yin et al., 2001), and c) geological time indications (cf. Section 2.1). The latter suggest that the age of 253.4 ± 0.2 Ma from just above the base of the Changhsingian Stage ( ...

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... latter suggest that the age of 253.4 ± 0.2 Ma from just above the base of the Changhsingian Stage ( Bowring et al., 1998) may be little too young. In the DCP 2003 are allocated durations of ∼ 4 my to the Changhsingian, of ∼ 5.5 Ma to the Wuchiapingian, and of ∼ 4.5 my to the Capitanian stages (Figs. 1 and 4). ...

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... using the age of 252.6 ± 0.2 Ma for bed 25 in Meishan ( Mundil et al., 2004), an age of ∼ 252.5 Ma has been derived for the PTB ( Fig. 1; Menning et al., 2005). The allocation of ∼ 0.4 my to each of the 25 Lopingian parasequences (17.3 in the Wuchiapingian and 7.7 in the Changhsingian) of Chen et al. (1998: Fig. 9) results in durations of ∼ 10 my for the Lopingian Epoch, of ∼ 6.9 my for the Wuchiapingian Stage and of ∼ 3.1 my for the Changhsingian Stage ( Menning et ...

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... 1; Menning et al., 2005). The allocation of ∼ 0.4 my to each of the 25 Lopingian parasequences (17.3 in the Wuchiapingian and 7.7 in the Changhsingian) of Chen et al. (1998: Fig. 9) results in durations of ∼ 10 my for the Lopingian Epoch, of ∼ 6.9 my for the Wuchiapingian Stage and of ∼ 3.1 my for the Changhsingian Stage ( Menning et al., 2005; Fig. 1). Thus, the age for the Capitanian-Wuchiapingian boundary is ∼ 261 Ma according to an age of ∼251 Ma for the PTB (STD 2002;GTS 2004) or ∼ 262.5 Ma according to an age of ∼ 252.5 Ma for the PTB ( Menning et al., 2005). The corresponding ages for the base of the Changhsingian Stage are 254.1 Ma and 255.6 Ma respectively, whereas in the ...

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... are rarely developed in the intra-Variscan basins, a complete succession is represented in the foreland basin between Central England and East Poland. There is no evidence of a break in time of ≥ 0.1 my in the migrating basin centre in North Germany during the 266 Ma to 229 Ma interval (STD 2002;Menning et al., 2005), but Wardlaw et al. (2004: Fig. 16.2) show a gap between the top of the Zechstein at 259.8 Ma and the base of the Buntsandstein at 251.0 ...