Pondrelli et al. 2014 - Senckenberg

Int J Earth Sci (Geol Rundsch) DOI 10.1007/s00531-014-1069-7

ORIGINAL PAPER

Depositional evolution of a lower Paleozoic portion of the Southalpine domain: the Mt. Pizzul area (Carnic Alps, Italy) Monica Pondrelli · Carlo Corradini · Luca Simonetto · Maria Giovanna Corriga · Erika Kido · Angelo Mossoni · Claudia Spalletta · Thomas J. Suttner · Nicola Carta

Received: 13 December 2013 / Accepted: 29 July 2014 © Springer-Verlag Berlin Heidelberg 2014

Abstract The Carnic Alps represent the best exposed Paleozoic succession within the Alpine domain being fossiliferous, mostly non-metamorphic and largely complete. This study focuses on the area around Mt. Pizzul, because the bedrocks record well the basin dynamics and most of the units are conodont bearing. Our aims were to contribute to the procedure of formalization of the lithostratigraphic units and to understand the depositional and deformation history of the study area. The area has been mapped, tectonic overprint constrained and the successions described, measured and dated. As a second-order aim, we discuss these data to infer the relations with other parts of the Carnic basin and to recognize the global controls on sedimentation. The depositional evolution can be sketched Electronic supplementary material The online version of this article (doi:10.1007/s00531-014-1069-7) contains supplementary material, which is available to authorized users. M. Pondrelli (*) International Research School of Planetary Sciences, Università d’Annunzio, Viale Pindaro 42, 65127 Pescara, Italy e-mail: [email protected] C. Corradini · M. G. Corriga · A. Mossoni · N. Carta Dipartimento di Scienze Chimiche e Geologiche, Università di Cagliari, Via Trentino 51, 09127 Cagliari, Italy e-mail: [email protected] M. G. Corriga e-mail: [email protected] A. Mossoni e-mail: [email protected]

as follows: pre-Hirnantian ramp-type margin; Hirnantian glacioeustatic-related deposits and unconformity; pelagic deposition in a ramp-type margin (Prˇídolí–Eifelian); slope formation and differentiation in buildup, foreslope and pelagic environments (Eifelian–Frasnian); transgression and reef drowning (Frasnian–Visean); probable subaerial exposures likely during in uppermost Famennian and Visean times; and turbidite deposition (Visean). Global controls or deposits suggesting a global control are documented, including the Boda Event, the Hirnantian glaciation, the Middle Devonian reef growth, the Kacˇák Event, and the high-frequency sea-level fluctuations around the Devonian–Carboniferous boundary. The drowning of the buildups here and elsewhere in the Carnic Alps started during the Frasnian, unlike observed globally. This suggests that local tectonics lead to progressive deepening up to the transition to turbidite deposition. E. Kido · T. J. Suttner Institute for Earth Sciences (Geology and Paleontology), University of Graz, Heinrichstrasse 26, 8010 Graz, Austria e-mail: erika.kido@uni‑graz.at T. J. Suttner e-mail: thomas.suttner@uni‑graz.at C. Spalletta Dipartimento di Scienze Biologiche, Geologiche ed Ambientali, Alma Mater Studiorum Università di Bologna, Via Zamboni 67, 40126 Bologna, Italy e-mail: [email protected]

N. Carta e-mail: [email protected] L. Simonetto Museo Friulano di Storia Naturale, Via Marangoni 39‑41, 33100 Udine, Italy e-mail: [email protected]

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Keywords Carnic Alps · Paleozoic · Pre-Variscan succession · Stratigraphy · Variscan and Alpine deformation

Introduction The Carnic Alps, located across the boundary between northeastern Italy and south Austria (Fig. 1a), are one of the classic areas of study of the Paleozoic in Europe (Geyer 1894; Taramelli 1895; Frech 1887; Vinassa de Regny and

Fig.  1 a Location map (inset) and digital elevation model of the Carnic Alps (based on SRTM, resolution 3 arc/s). Main villages and roads are indicated. The location of b is emphasized by the white box.

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Int J Earth Sci (Geol Rundsch)

Gortani 1905; Gortani 1908, 1920; Vinassa de Regny 1910, 1915; von Gaertner 1931; Heritsch 1928, 1936; Selli 1963; Bandel 1972; Vai 1976; Schönlaub 1979, 1985; von Raumer et al. 2013). This area corresponds to the external non- to low-metamorphic portion of the Variscan substratum of the Southern Alps (Vai 1976; Spalletta et al. 1982; Brime et al. 2008), thus recording a complex Variscan and Alpine deformational history (Läufer 1996; Läufer et al. 1997; Venturini 1990; Brime et al. 2008). The Variscan succession is highly fossiliferous and ranges from the Upper

b Topographic map of the study area and surroundings. Contours are 25 m. The study area is located by the red box


Fig. 2 Sketch of the geology of the Carnic Alps (simplified after Brime et al. 2008). The red box represents the location of the study area

Ordovician to the lowermost part of the Upper Carboniferous (Bandel 1972; Schönlaub 1979, 1985; Selli 1963; Vai 1976; Spalletta et al. 1982) (Fig. 2). Because of these characteristics, the Carnic Alps represent the best exposed Variscan succession within the Alpine domain. As a consequence, they represent a key for the regional reconstructions in order to unravel the Variscan evolution of a sector that later underwent Alpine deposition and deformations which concealed part of the older stratigraphies. Deposition is recorded from the Sandbian–Katian with clastic shoreface sediments passing distally to fossiliferous offshore shales and silts (Schönlaub 1979; Vai 1971). In the proximal part of the basin, the sedimentation changed gradually to carbonatic during the Katian and up to the Hirnantian and led to the deposition of bioclastic parautochthonous limestones passing distally to deeper water limestone and then shales (Vai 1971; Schönlaub 1979, 1988; Dullo 1992; Berry 2003). Limestones—corresponding to the Boda Event (Fortey and Cocks 2005)—have been interpreted as cool-water carbonates (Dullo 1992). The following sea-level drop associated with the Hirnantian glaciation (Berry and Boucot 1973; Sheehan 1973; Schönlaub 1988; Berry 2003) led to the local deposition of calcareous sandstones and to the exposure of most of the basin, while shales continued to be deposited in the deepest parts of the basin (Schönlaub 1988; Schönlaub and Histon 1999; Hammarlund et al. 2012). As a consequence of non-deposition and erosion, a disconformity is recorded on top of Hirnantian limestones or sandstones (Schönlaub 1979, 1980, 1988; Schönlaub et al. 1992; Berry 2003). In

the more proximal parts of the basin, the missing strata can reach up to several conodont zones of Llandovery and early Ludlow age (Schönlaub and Histon 1999; Storch and Schonlaub 2012). From the Llandovery, transgressive limestones started to onlap on top of the disconformity while deposition continued in the distal parts of the basin with mixed carbonatic shale and graptolitic shale deposits (Jaeger 1975; Jaeger and Schönlaub 1977; Spalletta et al. 1982; Schönlaub 1988, 1997; Schönlaub et al. 1992; Storch and Schonlaub 2012). The whole Silurian succession has been subdivided into four facies associations reflecting different deposition depth and hydraulic conditions, from shallower to deeper: (1) shallow water bioclastic limestones passing to (2) pelagic limestones, (3) interlayered limestones and shales and finally (4) graptolite-bearing shales and cherts deposited in poorly oxygenated to anoxic water (Schönlaub 1992, 1997; Wenzel 1997). During the Prˇídolí, uniform limestone sedimentation suggests more uniform depositional conditions within the basin (Schönlaub 1997). The relative sea-level curve presented by Brett et al. (2009) matches the global sea-level changes as inferred by Johnson (2006). Similarly to other parts of the world, the presence of dark gray to black organic rich and laminated shaly intervals suggest changes from aerobic to anoxic bottom conditions (Brett et al. 2009). During the lower part of the Lochkovian, the basin physiography remained similar to the Silurian one—characterized by a ramp-type margin (Spalletta et al. 1982)—but, in analogy with other areas such as eastern North America (e.g., ver Straeten 2010, cum ref.), more differentiated than

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in the Prˇídolí, with inner ramp represented mostly by crinoidal bearing fine-grained rudstone and grainstone passing distally to finer-grained outer ramp grainstone and packstone and then pelagic limestone and shales (Kreutzer 1990). The increased differentiation of the basin is reflected in the thicknesses of the deposits, with differences between proximal and distal deposits that increased by a magnitude order than during Ordovician–Silurian times (Kreutzer 1990; Schönlaub 1980). During the Lochkovian, but more frequently during the Pragian, patch reefs started to grow on top of the crinoidal lithoclastic limestones (Vai 1963; Bandel 1969; Kreutzer 1990; Suttner 2007). The basin profile remained ramp type through most of the Emsian with depositional environments prograding and retrograding following sea-level fluctuations (Kreutzer 1992). In the upper part of the Emsian, the physiography of the basin has been hypothesized to have changed progressively into rimmed shelf (Hubmann et al. 2003). As a consequence, the transition between shallow water and basinal facies passed to a slope margin, as suggested by the emplacement of gravity-driven deposits (Bandel 1972, 1974; Kreutzer 1990; Hubmann et al. 2003). This change in the basin physiography—although implicit in the facies description—has only been poorly described (Pohler in Hubmann et al. 2003). Shallow water facies consists of reef, lagoon and related environments prograding and retrograding following relative sea-level fluctuation (Bandel 1969; Kreutzer 1990; Hubmann et al. 2003; Buggisch and Joachimski 2006; Ellwood et al. 2011). Forereef facies consists of debris flow, turbidites, concentrated and hyperconcentrated flows intercalated with pelagic deposition. These facies pass gradually to pelagic carbonate and then shale deposits moving distally (Bandel 1972; Kreutzer 1990). It has not yet been deeply investigated whether the reefs—and associated environments—built a single large structure, including buildups of the Mt. Coglians-Hohe Warte, Mt. Zermula and Mt. Cavallo (Fig. 2) or several smaller and not geographically associated bodies. The reef development reached its climax in Givetian times (Vai 1976; Kreutzer 1990; Hubmann et al. 2003) and then ended at the Frasnian–Famennian boundary (Spalletta et al. 1982; Kreutzer 1990), probably as a response of a combination of climatic-driven eustatic fluctuations (Hubmann et al. 2003; Perri et al. 2009; Joachimski and Buggisch 2002; Joachimski et al. 2002, 2009; Kaiser et al. 2006) and extensional syn-sedimentary tectonics (Spalletta et al. 1982; Spalletta and Vai 1984). Transgressive pelagic clymeniid and goniatitid-bearing limestone were deposited in most of the basin from the Frasnian until the lower Visean with the exception of the most proximal part of the basin where a few shallow water bodies and related gravity-driven deposits persisted until

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at least somewhere in the Famennian, and the deepest part of the basin where radiolarite deposition occurred (Spalletta et al. 1982; Schönlaub 1985; Schönlaub et al. 1992; Perri and Spalletta 1998a, b). Higher-frequency eustatic fluctuations are still recognizable and probably controlled the emplacement of black shales in correspondence of the Hangenberg Event (Schönlaub et al. 1992; Perri and Spalletta 2000, 2001). The depositional framework between the end of the Famennian and the Visean is complicated by local episodes of exposure of part of the basin (Schönlaub et al. 1991; Brigo et al. 2001) while sedimentation continued in other deeper parts (Spalletta 1983). Syn-sedimentary strike-slip tectonics have been proposed to control this evolution (Spalletta and Venturini 1988, 1995; Venturini and Spalletta 1998). The deposition of the Hochwipfel siliciclastic turbidites between the Visean and the Bashkirian marks the beginning of the Variscan tectogenesis (Spalletta and Venturini 1988). Vertically and laterally this unit grades into volcanic and volcanoclastic deposits (Spalletta et al. 1980). Mt. Pizzul is located in the central part of the Carnic Alps (Fig. 1a, b) in a strategic position, because of its geographical location and its sedimentary evolution from the Upper Ordovician to the Lower Carboniferous. The sedimentation was in fact continuous during most of its geological history; erosion and reworking—although present— are limited; most of the units are conodont bearing, with a locally rich fauna. As a consequence, although severely affected by Variscan as well as by Alpine tectonics, the succession is one of the most continuous of the Carnic Alps and it is almost entirely datable using conodonts. In spite of these characteristics, this area—and in general some of the units which crops out here—is largely ignored in the literature. The starting point of this study was the realization of a detailed geological map (mapped at 1:5,000 scale) in order to unravel the complex structural setting and to identify the different lithological units and their stratigraphic relations. All the intervals of transitions between these units have been sampled for conodonts in order to identify the timelines across this part of the basin. We inferred—where possible—the depositional environments in which the different units were deposited and their evolution through time. Finally, we discuss the significance of our data in the framework of the whole Carnic basin and in the supra-regional context with the aim to distinguish the controls related to local tectonics from the global—possibly eustatic—trends.

Geological setting The Mt. Pizzul area is located in the central Carnic Alps (Fig. 1a), close to the syn- to post-Variscan deposits of the

Fig. 3 Geological map of the study area (see SM1 for a large size version). The red/white lines represent the location of Figs. 7, 10 and 11


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Fig. 4 Geological cross-sections emphasizing the severe north–south shortening (see Fig. 3 for the location of the traces and the legend of the geological units). The top-right inset represents a 3D sketch

simplifying the structural architecture unravelling strike-slip deformations and E–W trending thrusts cutting the Mt. Pizzul fold. Parasitic folds are not shown

Pramollo basin (Fig. 2). According to conodont color alteration index (CAI) and Kübler index (KI) data, the succession in this zone has been assigned to low anchizone conditions (Brime et al. 2008). We realized a 1:5,000 geological map of the study area (Figs. 3, 4, SM1). The Variscan succession of the Carnic Alps is currently under revision in order to formally

define the lithostratigraphic units and unravel the confusion arisen by the fact that the Carnic Alps spread between two countries with different languages (and a long history of stratigraphic research), reflected in different names and distinction for the geological units at all the ranks. Accordingly, the subdivision used in the map is still based on informal units, but we anticipated the definitive name and

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Fig. 5 Images showing some of the main structural features of the study area. a North dipping E–W trending thrust leading the overturned Eifelian?–Frasnian shallow water facies of the Zuc della Guardia and Zuc di Malaseit on top of the succession of the Mt. Pizzul area where the Middle Devonian crops out in slope to pelagic facies. b North dipping E–W trending thrust leading the Eifelian?–Frasnian

shallow water facies cropping out in the Mt. Zermula on top of the succession of the Mt. Pizzul area where the Middle Devonian crops out in slope facies. c Panorama of the northern part of the study area showing the succession of the Mt. Pizzul folded as a result of the north–south oriented compression

description where possible. In order to allow comparison with previous work, a synonymy is included in the unit descriptions. We refer to Hubmann et al. (2013) for a complete list of synonymy. The succession is overturned with the exception of the easternmost flank of the Mt. Pizzul (Figs. 3, 4c). According to Venturini (1990), this area represents part of the overturned flank of a pluri-kilometer-scale recumbent southverging fold of Variscan age. Parasitic folds at different scales are abundant especially in the most ductile lithologies, as well as fault bend folds which can locally cut part of the succession (Venturini 1990). These Variscan structures were disrupted and/or locally enhanced by at least three Alpine compressional phases (Venturini 1990; Läufer 1996), which have been characterized using the stress tensor inversion method from fault striations (Gephart 1990; Gephart and Forsyth 1984; Pondrelli 1998).

Alpine structures in the study area consist mainly of E–W to ENE–WSW trending and 45° to 60° north dipping south-verging thrust/inverse faults which, by analogy with the rest of the Carnic Alps and the whole Southalpine domain, have been interpreted of Tortonian–Serravallian age (Venturini 1990; Läufer 1996; Pondrelli 1998) (Figs. 2, 3, 4, 5). These structures lead the shallow water Middle Devonian units to overthrust the Upper Ordovician—Lower Carboniferous succession, most probably enhancing the Variscan shortening (Figs. 3, 4, 5). These systems also obliterated a major portion of the most proximal Middle Devonian slope deposits (Vinz formation) which are preserved only in a single outcrop located at 1,909 m of elevation west of the Forca di Lanza (Figs. 3, 4a). The thrusts led to a roughly N–S oriented shortening (Figs. 3, 4). Shortening is greater in the central and eastern parts of the mapped area (Figs. 3, 4: sections BB’ and CC’, 5a) than in the western sector (Figs. 3, 4: section AA’A’’,

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5b), where the most proximal Middle Devonian slope deposits are preserved. Variscan thrusts and folds are disrupted by sinistral strike-slip NE–SW trending faults and compressional roughly N–S trending faults (Fig. 3). The role of these fault systems is constrained—rarely in the study area but frequently in other parts of the Carnic Alps—by slickensides and steps on faults mirrors. According to Pondrelli (1998), based on the stress tensor inversion method from fault striations performed on the whole Carnic Alps, the Tortonian– Serravalian phase might have occurred in a transpressional stress regime. Most of the N90°E to N120°E trending faults present in the Carnic Alps have been then reactivated as dextral strikeslip faults during the Plio-Pleistocene deformative phase (Venturini 1990; Läufer 1996; Pondrelli 1998), developed, according to Pondrelli (1998), in a strike-slip stress regime. In the study area, although located just south of the NW– SE trending dextral strike-slip Cason di Lanza fault (Fig. 2) (Venturini 1990), this phase is only rarely expressed by the dextral strike-slip reactivation of some minor E–W trending faults (e.g., slickensides found west of the Forca di Lanza area).

Description of the units The stratigraphic framework of the study area is sketched in Fig. 6. Deposition ranges from Upper Ordovician to Lower Carboniferous. Some of the most significant conodonts found in the succession are illustrated in Plate 1. The succession is carbonate dominated, except for most of the Upper Ordovician and the Lower Carboniferous. Deposition was quite uniform along the Mt. Pizzul area with the exception of the Eifelian–Frasnian where the basin differentiated in a proximal shallow water part passing distally to slope and then to pelagic deposits (Table 1).

Fig. 6 Stratigraphic space–time diagram of the study area

Valbertad formation The Valbertad formation crops out along the footwall of the E–W trending thrust spanning from the path between west of Forca di Lanza, Zuc della Guardia and Zuc di Malaseit and at the top of Mt. Pizzul (Fig. 3). This unit corresponds to the ‘Uggwa Shale’ (e.g., Schönlaub 1979, 1985; Schönlaub and Histon 2000) of the Austrian literature and to the ‘Uqua pelites and siltites, part of the Uqua formation,’ of the Italian literature (Vai 1971). The Valbertad formation consists of thin- to very thinbedded light olive gray pelites passing upward to arenaceous shales and then to very fine-grained graywacke. This coarsening upward trend is quite irregular, with some thin levels of very fine-grained graywacke rarely

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interbedded also in the lowermost part of the unit. The yet not described fauna consists of abundant brachiopods and bryozoans, trilobites and rare cystoids, crinoids and gastropods. In the upper part (roughly the last 3 m) of the unit, the pelites are increasingly interbedded with thin to very thin nodules of medium light gray mudstone and wackestone (Fig. 7). Nodules become more abundant, thicker and with more lateral continuity going upward emphasizing a gradual transition to the following Uqua formation (Fig. 7). The maximum thickness of this unit is about 7 m in the eastern flank of the Zuc di Malaseit, but the lowermost limit (here like everywhere in the Carnic Alps) is always tectonic (Fig. 7).

Findenig fm

La Valute fm

Moderate pink to moderate red nodular mud- to wackestone with interlayered thin to medium beige marly laminae; medium dark gray breccia with a medium-fine-grained grainstone matrix passing distally to grainstone

Very thin to medium Interlayered mudstone–wackebedded stone and shales of grayish black color Medium dark gray and brownish Thin bedded gray nodular mudstone to wackestone

Rauchkofel fm

Thin to very thin bedded (mudstone/wackestone); thick bedded (breccia/grainstone)

Thin to rarely medium bedded

Dark gray mudstone to wackestone

Alticola limestone

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10 Upper contact: Rauchkofel Fm. (conformable, sharp) 2 Upper contact: La Valute Fm. (conformable, sharp) Upper contact: Findenig 10 Fm. (conformable, transitional)

Sponge spicule, trilobites, articulated ostracod shells, orthocone nautiloids, crinoid stem plates, brachiopod shells, dacryoconarids, conodonts Dacryoconarids, frag- Upper contact (proximal): Vinz ments of trilobites, Fm. (conformable, ostracods, small sharp); Upper contact orthocone nautiloids, (distal): Hoher Trieb crinoid stem plates; Fm. (conformable, rare conodonts interfingered)

Crinoid loboliths, conodonts

Cephalopod, conodonts

5.5

Upper contact: Alticola lm. (disconformable)

Fossil debris (mainly brachiopods)

Thin to medium bedded; Hummocky, high-angle and low-angle crossstratification

Shales passing upward to dark gray calcareous sandstone; coarsening upward succession

Plöcken fm

Offshore part of a Katian (Walliser 1964; ramp-type margin Serpagli, 1967; Schönlaub 1988; Bagnoli et al. 1998; Ferretti and Schönlaub 2001) Hirnantian (Schönlaub, Offshore/offshore transition passing 1988; Ferretti and to shoreface and Schönlaub 2001; foreshore Storch and Schonlaub 2012) Prˇídolí Offshore part of a ramp-type margin 3.5

Conodonts

Thin to very thin beds

Medium light gray nodular mudstone and wackestone with light olive fine-grained graywacke interlayered

Uqua fm

Offshore part of a ramp-type margin Middle and upper Lochkovian (Corriga 2011; Corriga et al. 2011)

Upper Lochkovian (Cor- Offshore part of a ramp-type riga et al. 2011)—Eifmargin; gradual elian (sample ZMB emplacement of BF) a slope

Offshore part of a ramp-type margin

Lower Lockhovian

Offshore part of a ramp-type margin

Sandbian–Katian (Bagnoli et al. 1998)

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Lower contact: unknown; upper contact: Uqua Fm (conformable, transitional) Upper contact: Plocken Fm. (conformable, sharp)

Brachiopods, bryozoans, trilobites, rare cystoids, crinoids, gastropods

Thin to very thin bedded

Light olive gray pelites passing upward to arenaceous shales and then to very fine-grained graywacke

Valbertad fm

Depositional environment

Age

Maximum thickness (m)

Contacts

Fossils

Bedding and sedimentary structures

Lithology

Formation

Table 1 Summary of the characteristic features of the formations mapped in the study area


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Pal Grande fm (including Zollner fm.)

Hoher Trieb fm

Vinz fm


Fossils

Contacts

Thick bedded

Calcareous algae, fora- Tectonically sliced minifers, Amphiporae, stromatoporoids, rugose and tabulate corals, gastropods, ostracodes and brachiopods (Ferrari and Vai 1966; Corradini et al. 2012; Lamberty 2013) Medium to very thick bed- Disarticulated crinoid Tectonically sliced Medium gray clast-supported stems and corals, ded; planar lamination; and disorganized breccia in a fragmented shells, fine-grained grainstone matrix; normal grading; inverse conodonts grading packstone to locally laminated grainstone with interlayered rare fine-grained breccia; clast-supported breccia locally displaying an erosional base; crinoidal grainstone and/or fine-grained breccia passing upward to clast-supported breccia Upper contact: Pal Dacryoconarids, calThin to very thick bedBreccia layers of centimeterGrande Fm. (conformcispheres, foraminifded; planar lamination; scale thick clasts in a fineable, sharp) era, sponge spicules, normal grading; inverse grained grainstone matrix; algae, trilobites, grading medium gray grainstone bivalves, crinoids, to packstone locally passbrachiopods, palynoing upward to wackestone morphs, ostracods, and/or mudstone with thick reworked silicified laminae of silt/shale interbeds; corals, conodonts graywacke; black chert beds and nodules; black shales Upper contact: HochVery thin to thin (rarely Clymenids, bivalves, Light gray to moderate pink wipfel Fm. (disconmedium) bedded; slump; ostracods, radioand red nodular mudstone to formable) laria, brachiopods, wackestone; breccia with very trilobites, crinoids, angular to mostly sub-rounded conodonts carbonatic clasts centimeter to decimeter scale in diameter, separated by millimeter- to centimeter-scale thick olive green partings of silty shale; chert with limestone lenses

Lithology

Shallow water Amphipora bafflestone, algal laminites, floatstone with (undifferenmicritic matrix and fenestral tiated) loferite

Formation

Table 1 continued

Eifelian—lower Frasnian

Pelagic deposiMiddle Frasnian– tion (mudmiddle Famennian /wackestone (mudstone/wackestone and chert with unit); upper Famenlimestone lenses); nian (breccia); Tournaisian (chert with subaerial exposure (breccia) limestone lenses)

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20–25

Transition between forereef slope and basin floor

High-density gravity-driven flows in a forereef slope environment

Eifelian

25

Reef and back reef


Eifelian?–Givetian

Age

~250

Maximum thickness (m)


Nor visible in the study area

Contacts N/A

Maximum thickness (m) Middle Visean– Bashkirian (Venturini et al. 2009)

Age

The ages refer to the study area. Where data from the study area are not available, references from other Carnic Alps locations are reported. See text for details

Remains of plants Yellowish gray lithic sandstones Thin to medium (very rarely thick) bedded; and graywackes interbedded normal grading; planar with dark gray laminated lamination; ripple marks shales; breccia with centimeter-scale cherty clasts

Fossils

Hochwipfel fm


Lithology

Formation

Table 1 continued

Turbiditic deposits



Fig. 7 Stratigraphic log of the Upper Ordovician units measured west of Forca di Lanza (see Fig. 3 for location)

On the basis of the shelly fauna, the stratigraphic distribution of this unit has been determined as Sandbian–Katian in the nearby locality of Valbertad (Bagnoli et al. 1998). Even if there are no detailed studies on the brachiopod fauna (Schönlaub, pers. comm.) to constraint the

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Fig. 8 Images of some of the units mapped in the study area. a Nodular mudstone–wackestone of the Uqua formation. b Calcareous cross-bedded sandstones of the Plöcken formation. c Findenig forma-

tion. d Iron nodule in the Vinz formation. e Clast-supported breccia of the Vinz formation. f Crinoidal fine-grained breccia within the Vinz formation

correlations with supposedly coeval deep water Foliomena fauna occurring at Cellon mountain (Schönlaub et al. 2011), the Valbertad formation show sedimentological characters and fauna content consistent with deposition in an offshore setting, as suggested by Schönlaub (1969, 1988) and Vai (1971) in other parts of the Carnic Alps. The coarsening upward trend—even if crude—suggests a shallowing upward succession.

Uqua formation

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The Uqua formation crops out west of the Forca di Lanza and at the top of Mt. Pizzul (Fig. 3). It corresponds to the ‘Uggwa Limestone’ of the Austrian literature (e.g., Schönlaub 1979, 1985; Schönlaub and Histon 2000) and to the ‘Uqua limestone, part of the Uqua formation,’ of the Italian literature (Vai 1971).


Fig. 9 Images of some of the units mapped in the study area. a Packstone to wackestone interlayered with chert and black shales: Hoher Trieb formation. b Silicified coral within a breccia layer of the Hoher Trieb formation. c Traction structures within the Hoher Trieb forma-

tion. d Slump at the base of the Pal Grande formation. e Red mudstones within the Pal Grande formation. f Breccia-like deposits at the top of the Pal Grande formation

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Fig. 10 Schematic sketch of the Vinz Fm. measured west of Forca di Lanza (see Fig. 3 for location)

The Uqua formation consists of thin to very thin beds of medium light gray nodular mudstone and wackestone with very thin beds of light olive fine-grained graywacke interlayered (Figs. 7, 8a). Conodonts are the only abundant fossils. The transition to the following Plöcken formation is marked by a distinct change in color between light olive gray and medium dark gray shales (Fig. 7). The shales are very poorly exposed in the study area, so the nature of the contact is not clear. The maximum thickness of Uqua formation is about 3.5 m west of the Forca di Lanza. This unit has been assigned to the ordovicicus Zone on the basis of the conodont fauna described in other parts of the chain (e.g., Cellon, Uqua Valley, Valentintörl, Valbertad) (Walliser 1964; Serpagli 1967; Schönlaub 1988; Bagnoli et al. 1998; Ferretti and Schönlaub 2001). The Uqua formation at the Cellon locality has been interpreted as deposited in an offshore setting in a ramptype margin (Dullo 1992) which is consistent with the interpretation of the study area. The gradual transition from the Valbertad formation to the Uqua formation has been interpreted either as following warmer (Vai 1971) or cooler (Dullo 1992) conditions. Inferences about this issue were beyond the aim of this paper, but the crude coarsening upward trend registered in the uppermost part of the Valbertad formation which continues in correspondence

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of the siliciclastic laminae and layers within the Uqua formation, suggests the progressive approaching of the source area. Plöcken formation The presence of the Plöcken formation in the Pizzul area, precisely west of the Forca di Lanza, is here reported for the first time on the Italian side of the Carnic Alps (Fig. 3). We use here the term introduced by Schönlaub (1985, p. 38) which corresponds to the ‘Untere Schichten’ of the older Austrian literature (e.g., von Gaertner 1931; Schönlaub 1979, 1980: p. 18, Fig. 7a). This unit is not present in the stratigraphic scheme of the Italian literature, most probably because its presence on the Italian side of the Carnic Alps was not detected before. The Plöcken formation consists mainly of calcareous sandstone (Fig. 9b). At the base of the unit, a few centimeters of shales are present, passing upward to dark gray thin- to medium-bedded calcareous sandstone (Fig. 7). The succession is crudely thickening upward (Fig. 7). The calcareous sandstones are moderately sorted and medium- to fine-grained at the base passing upward to medium-, coarseand then very coarse-grained in correspondence of the topmost bed, thus defining a coarsening upward succession (Fig. 7). Disarticulated fossil debris (mainly brachiopods) is


Fig. 11 Stratigraphic log of the Middle to Upper Devonian units measured west of Forca di Lanza and east of Zuc di Malaseit (see Fig. 3 for locations)

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very common. Shells are dispersed in the sediment without the convex-down position orientation which might suggest that they are in situ or only slightly removed. The succession is relatively poorly exposed, so it is difficult to observe the sedimentary structures, but in the lowermost part hummocky cross-stratification appears to be present (Fig. 7). In the upper part of the succession, both high-angle and lowangle cross-stratifications are present, with only low-angle cross-stratification in the uppermost layer (Fig. 7). The total succession is about 5.5-m-thick west of Forca di Lanza (Fig. 7), while in the rest of the study area (with the possible exception of a block in the detritus in the eastern flank of Zuc di Malaseit) this unit does not crop out and the following Silurian deposits rest on top of the Valbertad formation. The contact of the Ordovician strata with the overlying Silurian succession is thus disconformable, marking the erosion of part of the Upper Ordovician succession and the erosion and/or non-deposition of the Lower Silurian one. The Plöcken formation has been assigned to the Hirnantian at the Cellon section (Schönlaub 1988) due to the presence of fossils belonging to the Hirnantia fauna which first occurs in the shales at the base of the succession and on the basis of a rich conodont fauna (Ferretti and Schönlaub 2001) and the index graptolite Metabolograptus persculptus (Storch and Schonlaub 2012). The unit is poorly exposed, thus preventing from a consistent analysis of the sedimentary structures. Still, the texture characteristics, the fossils content and possible reworking, the observable sedimentary structures and the overall sequence architecture are consistent with a progradational shoreline sequence (Fig. 7). Offshore/offshoretransition deposits appear to pass upward to shoreface and then foreshore deposits, presumably as a result of a sea-level drop associated with the Hirnantian glaciation. This interpretation is consistent with the one proposed by Schönlaub (1988). The disconformity on top, in correspondence of which locally the entire Plöcken formation, Uqua formation and part of the Valbertad formation are eroded, suggests that the sea-level drop and associated forced regression persisted and exposed the more proximal part of the Ordovician deposits. However, in the Cellon section, Schönlaub and Sheehan (2003) and Schönlaub and Lammerhuber (2009) hypothesized that the Plöcken Fm was deposited during the global sea rise following the regression, which is apparently in contradiction with the observations from the study area at Pizzul. Still, most probably the Plöcken formation is diachronous, and in the Pizzul area, we register only the regression prior to the following transgression observable at the Cellon section (Schönlaub and Sheehan 2003; Schönlaub and Lammerhuber 2009).

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Alticola limestone The Alticola limestone crops out north and west of the Forca di Lanza, close to the top of Mt. Pizzul and in a small tectonized strip in the east side of the Zuc di Malaseit (Fig. 3). The Alticola limestone represents a classic unit both in the Austrian and Italian literature (e.g., von Gaertner 1931; Vai 1976; Schönlaub 1979, 1985). In the study area, it consists of prevalently thin- to rarely medium-bedded medium dark gray cephalopod-bearing mudstone to wackestone. Beside cephalopods, the only abundant fossils observed in the study area are conodonts. The upward transition to the Rauchkofel formation is never exposed in the study area, but in nearby areas, e.g., Rio Malinfier West section (Corriga 2011; Corradini et al. 2012), it is very well exposed and marked by a distinct and sharp change in color. The maximum thickness measured in the study area is about 10 m at Mt. Pizzul. In this area, the Alticola limestone rests disconformably on top of Ordovician units, because the older Silurian carbonatic units of the Carnic Alps (Kok and Cardiola formations) are not present. On the basis of conodont data, the unit ranges from Ludlow to lowermost Lochkovian in age, spanning the Silurian–Devonian boundary in its uppermost part (Walliser 1964; Corriga and Corradini 2009; Corradini and Corriga 2010, 2012; Corradini et al. 2014). In the study area, the Alticola limestone appears to be limited to the Prˇídolí. In conformity with what is suggested for the whole Carnic Alps area (Wenzel 1997), we interpret the Alticola limestone within the study area as deposited in the offshore part of a ramp-type margin. Rauchkofel formation The presence of the Rauchkofel formation has been detected in this area only thanks to few loose blocks in the debris north of Forca di Lanza. The Rauchkofel formation corresponds to the ‘pelagic Rauchkofel limestone’ introduced by Schönlaub (1985) of the Austrian literature and to the ‘calcari lastroidi’ (e.g., Selli 1963; Spalletta et al. 1982) and ‘Calcari del Rauchkofel’ (Venturini et al. 2009) of the Italian literature. Because of the lack of exposure in the study area, the description is based on sections located nearby (e.g., Rio Malinfier West, Corradini et al. 2012). The Rauchkofel formation—in this part of the basin— consists of very thin- to medium-bedded interlayered mudstone–wackestone and shales of grayish black color (Corradini et al. 2012).


The overall thickness in this area is roughly estimated at two meters, although faulting of strata in or out of the section cannot be excluded. The upward transition to the La Valute formation is not exposed in the study area but in the nearby areas (e.g., La Valute) is marked by a sharp change in color. However, the boundary interval between the two formations, in and out of the study area, is often a preferential surface for thrust faults, and the gain or loss of section. Due to the poor exposure conditions, it is impossible to date the Rauchkofel formation in this area, with the exception of the occurrence of crinoid loboliths on a loose block north of Forca di Lanza. The presence of the crinoid loboliths in the Carnic Alps is considered to indicate a lowermost Lockhovian age, which has been confirmed by conodont data (hesperius Zone; Corradini et al. 2005). On the basis of the sections located in nearby areas, we interpret this part of the Rauchkofel formation as deposited in the offshore part of a ramp-type margin in conformity with what it is suggested in the whole Carnic Alps by Schönlaub and Histon (1999). We found in fact no evidence of gravity-driven deposition that might suggest the presence of a physiographic break. La Valute formation The La Valute formation crops out extensively over the whole study area (Fig. 3). This unit corresponds to the ‘Boden limestone’ of the Austrian literature (e.g., Schönlaub 1985; Kreutzer 1992) while in the Italian literature this unit was included within the ‘tentaculitid limestone’ (Selli 1963; Spalletta and Venturini 1990). The term ‘La Valute’ has been introduced by Corriga et al. (2011). It consists of thin-bedded medium dark gray and brownish gray nodular mudstone to wackestone. In the uppermost couple of meters of the unit, beds become very thin to thin with silty and marly intercalations and then gradually pass to the Findenig formation. The La Valute formation yields sponge spicule, trilobites, articulated ostracod shells, orthocone nautiloids, some strongly fragmented crinoid stem plates and brachiopod shells, dacryoconarids and conodonts (Corriga et al. 2011). The La Valute formation is roughly 10 meters thick in this area. On the basis of its conodont content, this unit can be assigned to the middle and upper Lochkovian (carlsi-pandora beta zones) (Carta 2011; Corriga 2011; Corriga et al. 2011). We interpret the La Valute formation as deposited in the offshore part of a ramp-type margin.

Findenig formation The Findenig formation crops out extensively across all of the study area (Fig. 3). This unit was named for the first time as ‘Findenig limestone’ by Schönlaub (1980) and corresponds to the ‘Roter Flaserkalk’ and ‘Knollenkalk’ of the older Austrian literature (e.g., Pölsler 1969; Bandel 1972, 1974; Schönlaub 1979) and to part of the ‘tentaculitid limestone’ of the Italian literature (e.g., Selli 1963; Vai 1963, 1980). It consists of thin- to very thin-bedded moderate pink to moderate red nodular mudstone to wackestone with interlayered thin to medium beige marly laminae (Fig. 8c). Near the top of the formation, a roughly 1.5-m-thick bed of medium dark gray breccia with a matrix made of mediumfine-grained grainstone is very distinctive. The layer is matrix supported and disorganized and represents the first breccia episode in this part of the Carnic basin (Fig. 11: breccia level I). This level passes distally to grainstone with diminished thickness of about 50 cm. The fauna is dominated by dacryoconarids, while fragments of trilobites, ostracods, small orthocone nautiloids and crinoid stem plates are less abundant. Conodonts— although present—are quite rare. This unit is overlain by the Vinz formation in the peak at 1,909 m of elevation, located just west of Forca di Lanza and by the Hoher Trieb formation in the rest of the study area (Figs. 3, 4). The upper boundary with the Vinz formation is sharp, while the upper transition to the Hoher Trieb formation is interfingered. For mapping purposes, the upper boundary of the Findenig formation is located just above the last pink layer. This unit is about 30 m thick (measured west of the Forca di Lanza), but due to its lithological character, it is often folded and thrusted, so that it is frequently subjected to either structural thickening or thinning. On the basis of the scarce conodont fauna, the Findenig formation has been shown to range from the upper Lochkovian (La Valute cave: Corriga et al. 2011) up to the Eifelian (sample ZMB BF). We interpret the mudstone/wackestone deposits of the Findenig formation as settled in offshore settings of a ramp-type margin. Unlike in relatively nearby areas such as Stua Ramaz, there is no evidence of the so-called allodapic layers (Bandel 1974; Vai 1980), which have been interpreted either as tempestites (Vai 1980) or turbidites (Bandel 1974). The first and only evidence of a gravity-driven process is the breccia to grainstone bed located a few meters below the upper formation boundary. The textural characteristics of this deposit suggest a deposition from a selective frictional flow, possibly an hyperconcentrated density flow passing distally to a grain flow (Mulder and Alexander

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Fig. 12 Correlation chart of the units measured in the study area

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Plate 1 Conodonts from the study area: a Palmatolepis rhomboidea Sannemann, 1955; sample PZW 5. b Palmatolepis minuta minuta Branson & Mehl, 1934; sample PZW 5. c Ancyrodelloides murphyi Valenzuela-Rios, 1994; sample CAD III. d Panderodus unicostatus (Branson & Mehl, 1933); sample CAD II. e Dapsilodus obliquicostatus (Branson & Mehl, 1933); sample CP 4. f Panderodus recurvatus (Rhodes, 1953); sample CAD II. g Belodella resima (Philip, 1965); sample CAD II. h Polygnathus nothoperbonus Mawson, 1987; sample PZS 2. i Polygnathus eiflius Bischoff & Ziegler, 1957; sample CAD P2. j Polygnathus ensensis Ziegler & Klapper,

1976; sample CAD P2. k Polygnathus timorensis Klapper, Philip & Jackson, 1970; sample CAD P1. l Polygnathus angusticostatus Wittekindt, 1966; sample FL 2. m Protognathodus praedelicatus Lane, Sandberg & Ziegler 1980; sample PZW H. n Gnathodus typicus Cooper, 1939; sample PZW H. o Klapperina ovalis (Ziegler & Klapper in Ziegler, Klapper & Lindström, 1964); sample FL 7. p Polygnathus pseudofoliatus Wittekindt, 1966; sample CAD P1. q Schmidtognathus wittekindti Ziegler, 1966; sample FL 6. r Polygnathus linguiformis linguiformis Hinde, 1876; sample FL 1. s Palmatolepis glabra glabra Ulrich & Bassler, 1926; sample PZW 5

2001). Such textures and processes suggest that the physiography of the basin might have started gradually—possibly locally—to develop a slope allowing gravity processes to be activated.

Shallow water undifferentiated units

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In the northern part of the study area (Mt. Zermula, Zuc della Guardia and Zuc di Malaseit), shallow water Middle


Devonian limestones crop out (Ferrari and Vai 1966). The detailed analysis of these units was beyond the goals of this work, so we simply grouped all the shallow water facies in a single unit without trying to distinguish the different formations (Fig. 3). According to Ferrari and Vai (1966), the facies association consists of thick-bedded Amphipora bafflestone, algal laminites, floatstone with micritic matrix and fenestral loferite. The succession is tectonically sliced at the base and at the top and repeated by E–W trending thrusts (Figs. 2, 3), so the reconstruction of its stratigraphy and facies evolution necessitates a detailed study. Calcareous algae, foraminifers, stromatoporoids, rugose and tabulate corals, gastropods, ostracodes and brachiopods (Ferrari and Vai 1966; Corradini et al. 2012; Lamberty 2013) are present in the mapped area, while stromatoporoids of the genus Amphipora are abundant more to the north close to the Cason di Lanza fault. Although we dissolved more than 30 limestones samples for microfossils (3–5 kg/ sample), no diagnostic conodonts have been obtained from the massive limestones so far. The overall thickness of this unit in the study area can be only roughly estimated, due to tectonic cuttings, to roughly 250 m. According to Ferrari and Vai (1966), no rocks older than the Givetian crop out, but on the basis of analogy with the units present in other shallow water areas, we suspect that possibly some Eifelian units can be present as well. The succession has been interpreted as deposited in reef and back-reef depositional environments by Ferrari and Vai (1966). Vinz formation The Vinz formation crops out only in a single location at the top of the hilltop at 1,909 m elevation west of Forca di Lanza along the path to Mt. Zermula (Fig. 3). The term ‘Vinz limestone’ was introduced by Kreutzer (1992) to distinguish the uppermost part of the ‘Kellerwand limestone’ (Schönlaub 1985). It corresponds to part of the ‘coral and crinoid limestone’ (e.g., Schönlaub 1980) of the Austrian literature and to part of the ‘yellow bedded limestone’ (Spalletta and Venturini 1990), the ‘calcareniti di transizione distali’ (Venturini et al. 2001) and the ‘calcareniti di Pal Grande’ and ‘calciruditi del Freikofel’ (Venturini et al. 2009) of the Italian literature. This unit characterizes the so-called transitional units (Kreutzer 1990, 1992) cropping out in the Passo di Monte Croce Carnico area (Fig. 2). Lithologies, facies and stratigraphic position appear to be consistent, so we maintained the subdivision here as well. Still, here, we observe no occurrence of the pelagic facies that in the Passo di Monte Croce Carnico area interfingers with the breccia beds. This

unit is not exactly the same as the Vinz formation found in the Passo di Monte Croce Carnico area but we consider that the absence of one facies is not a condition to introduce a new unit. The unit consists of various facies, all of them of medium gray color: thick-bedded clast-supported and disorganized breccia in a fine-grained grainstone matrix (Fig. 8e); medium-bedded packstone to locally laminated grainstone with interlayered rare thin- to medium-bedded fine-grained breccia; very thick-bedded clast-supported breccia locally displaying an erosional base; mostly crinoidal grainstone and/or fine-grained breccia (Fig. 8f) passing upward to clast-supported breccia (Fig. 10). Some clasts consist of mudstone; others are classified as grainstone or fine-grained breccia or yield fossils, mostly crinoids and corals. Silica replacement of corals is common. Ironrich nodules are quite abundant in a portion of the unit (Figs. 8d, 10). The facies description is partly hampered by the fact that the outcrop is very tectonized, which precludes measuring an accurate stratigraphic section (Fig. 10). Still, an overall, albeit irregular, thickening and coarsening upward trend seems to be recognizable (Fig. 10). The fauna is dominated by the presence of disarticulated crinoid stems and corals, and a few fragmented shells are observable in thin sections. Conodonts are present, although not abundant. The upper limit of the Vinz formation here is not observable, because of tectonic elision (Figs. 3, 10). The overall thickness of the unit in the study area is difficult to measure, again because of tectonic reasons, but we can roughly estimate a maximum thickness of about 25 m. Based on conodont data (sample FL2), the central part of the unit (Fig. 10) is dated to the late Eifelian (kockelianus Zone). Characteristics of Vinz formation facies in the study area are consistent with deposition by high-density gravity-driven flows in a forereef slope environment. Most of the breccia beds appear to have been deposited by hyperconcentrated density flows; however, normal grading in the top of some beds allows us to hypothesize deposition by concentrated density flows, according to Mulder and Alexander’s (2001) classification. Inverse grading at the base of the beds is common due to the development of traction carpets. Some finergrained deposits (i.e., laminated grainstone) might possibly represent turbidite deposition. No evidence of pelagic intercalations has been found in this unit. This may in part due to tectonic elisions, but the presence of mudstone, wackestone and iron-rich nodules within the clasts of the breccia, overlying the erosional base found in several breccia layers, allow us to hypothesize that some deeper or calmer water finegrained deposition occurred at times, which later was eroded by succeeding gravity-driven flows.

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Hoher Trieb formation The Hoher Trieb formation crops out extensively across all of the study area (Fig. 3). It represents the best exposed unit of the study area, with several easily reachable sections for the different parts of the unit and a section located west of Forca di Lanza where the unit is completely exposed (Fig. 11). Even if this unit in the study area is very similar to the ones presented in sections such as Hoher Trieb (Schönlaub 1969), Oberbuchach II (Schönlaub 1985) and Lodin/Findenigkofel (Pölsler 1969), differences in the relative thicknesses and distributions of the facies lead us to suggest this section as reference section in the framework of the procedure of formalization of the unit. The name Hohe Trieb limestone has been introduced by Schönlaub (1985) and Kreutzer (1992); since then, it is used in the Austrian literature (e.g., Schönlaub 1992; Hubmann et al. 2003). It corresponds to the ‘silicified coral bed and platy limestone’ of Schönlaub (1980) and to part of the ‘Monte Lodin formation’ of Selli (1963). We slightly change the name of the unit to Hoher Trieb formation, in compliance with the name of the type locality. This unit consists of several intercalated facies (Fig. 11). Four levels of medium gray meter-thick breccias punctuate the succession and have been found and mapped (Figs. 2, 11). The breccia layers consist of centimeter-scale thick clasts in a fine-grained grainstone matrix. Frequently, clasts consist of silicified corals (Fig. 9b). The texture is disorganized; matrix supported in the first two layers (breccia levels II and III) and clast supported in the last two (breccia levels IV and V). Breccia level IV in the uppermost part changes to well organized and display normal grading up to laminated grainstone (Fig. 11). The thickness of the same bed changes from place to place. Tectonic overprint complicates the framework, but—at least for the oldest two breccia episodes—it seems that the maximum thicknesses are reached east of the Zuc di Malaseit and the minimum west of Forca di Lanza. This trend suggests that the depositional processes are selective. The source area appears not to coincide with the one documented for the breccia level I inside the Findenig formation which thins in a different direction. Breccias are interlayered with medium gray thin- to thick-bedded grainstone to packstone locally passing upward to wackestone and/or mudstone with thick laminae of silt/shale interbeds. Sedimentary structures such as planar laminations, indicative of traction flows, suggest highvelocity flows (Fig. 9c). Bedding is quite regular, giving the unit the characteristic platy pattern which resulted in the name ‘platy limestone’ of the older Austrian literature (e.g., Schönlaub 1980). This pattern is emphasized by local interlayered medium gray thin- to medium-bedded wackestone beds. Rare intercalations of dusty yellow thin- to

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thick-bedded graywacke are also present. These beds often display lamination and normal grading passing upward from medium to fine-grained graywacke followed by shales (Fig. 11). Especially in the interval between breccia levels III and IV, but less commonly also in other intervals, limestone beds are interlayered with black thin to medium thick chert beds. Chert nodules are also found inside some of the limestone beds (Fig. 11). In the interval between breccia levels III and IV, some centimeter-thick black shale levels are also occurring (Fig. 9a). The fauna is dominated by the presence of dacryoconarids, but also calcispheres (including Bisphaera sp. and Radiosphaera sp.), foraminifera (including Baculella sp. and Parathurammina sp.), sponge spicules, algae (including dasycladacee), trilobites, bivalves, crinoids, brachiopods and ostracods have been detected in the wackestone to packstone beds, and reworked silicified corals within the breccia beds. Palynomorphs are also present. Conodonts are always present and very abundant in some levels. This unit passes upward to the Pal Grande formation with a sharp boundary (Figs. 3, 11), marked by a grain size decrease to mudstone–wackestone on top of the breccia level V. This sharp transition is often a preferential surface of decollementtype faulting, so it is rarely preserved. Still, west of Forca di Lanza, the boundary is preserved and is marked by a slump at the base of the Pal Grande formation (Fig. 9d). This unit is about 35–50 m thick depending on the location. This large variation is mostly likely due to either structural thickening and/or thinning, or to variations in depositional thicknesses of breccia units (Fig. 11). This unit shows intercalations between gravity-driven and pelagic deposits, the latter deposited in different environmental conditions. In particular, the overall textural and sedimentological characteristics of the breccia deposits— although not studied through a detailed microfacies analysis—are consistent with a deposition through gravity-driven hyperconcentrated to concentrated density flows (Mulder and Alexander 2001). The varying bed thickness is also consistent with selective processes. The grainstone to packstone beds are consistent with a formation by turbiditic currents, which would explain the platy shape, the common normal grading and the local evidence of traction. Instead, finer-grained limestone and black shales reflect deposition in deeper or calmer water through settling processes during the phases in which the re-deposition was not active or could not reach those parts of the basin. Black shales in particular suggest formation under dysoxic/anoxic conditions. Chert formation reflects diagenesis of either limestone or fine-grained siliciclastic deposits. We interpret the Hoher Trieb formation as formed in correspondence of the transition between forereef slope and basin floor.


Pal Grande formation The Pal Grande formation crops out extensively across all of the study area (Fig. 3). The term ‘Gross-Pal Clymenienkalk’ was mentioned by Frech (1894), but then this unit was named ‘Goniatiten– Flaserkalk’ in the Austrian literature (e.g., Bandel 1974; Schönlaub 1980) and ‘Climenid pelagic limestone’ (e.g., Spalletta and Venturini 1990),’clymeniid and goniatitidbearing pelagic limestone’ Venturini and Spalletta (1998), and ‘Calcari a goniatiti e climenie’ (Venturini et al. 2001) in the Italian literature, considering a unit spanning Upper Devonian to Lower Carboniferous in age. More recently, the term ‘Pal limestone’ was introduced by Schönlaub (1985) to represent the Upper Devonian part of the unit while ‘Kronhof limestone’ referred to the Carboniferous part. The Pal Grande formation is going to be formalized to indicate the whole Devonian–Carboniferous unit. In this area, the Pal Grande formation consists of light gray to moderate pink and red very thin- to thin (rarely medium)-bedded nodular mudstone to wackestone (Mossoni et al. 2013) (Fig. 9e). The nodular pattern is emphasized by thin laminae of silty shale. The sharp basal transition from the Hoher Trieb formation is often tectonically omitted because of the different lithologic characteristics at the boundary, but—when preserved—the first level of the Pal Grande formation is characterized by a convoluted and contorted bedding that we interpret as a slump deposit (Figs. 9d, 11). This kind of bedding is found only at that level, thus allowing— together with its contorted geometry—to exclude a tectonic origin. Moreover, the absence of current-associated deposits prevents an interpretation as convolutions. Softsediment deformation forms also due to sediment failure in response to density contrast and in the presence of fluids with high pore pressure. Still, we do not observe lithological differences (the contorted unit is made up entirely of mudstone/wackestone) able to create a density contrast nor fluid escape/load structures to support such a scenario. According to conodont fauna, the slump would have affected rocks of Frasnian age (between MN 3 and MN11 Zones). The moderate pink to moderate red limestone constitutes an intermediate unit in the succession while the lowermost and the uppermost parts of the Pal Grande formation are light gray. The very uppermost part of the Pal Grande formation—up to roughly a couple of meters but detected only west of the Mt. Pizzul top—consists of a breccia or a nodular limestone with very angular to mostly sub-rounded carbonatic clasts/nodules centimeter to decimeter scale in diameter, separated by millimeter- to centimeter-scale thick olive green partings of silty shale. Overall, the strata are very poorly sorted with no organization visible in the deposit (Fig. 9f). The carbonate clasts/nodules consist of

light gray mudstone/wackestone of Lower expansa Zone. On top of this facies, a 7–8-cm-thick level of dark gray to black thinly laminated chert has been found. In the following roughly 20 cm, cherts are interbedded with 5- to 10-cmthick dark gray mudstone/wackestone, that have been dated by means of conodonts to a short interval of Tournaisian age (anchoralis-latus Zone). This cherty limestone facies shows remarkable affinity with the radiolarites described by Spalletta (1983) in the Rio Chianaletta, and accordingly, it should be included in the Zollner formation and not in the Pal Grande formation. Still, because of its limited thickness, which prevents separate mapping of the units, and because of its single occurrence west of Mt. Pizzul, which discourages from mapping it as a horizon, we decided to include this unit within the Pal Grande formation. Within the Pal Grande formation, the fauna is characterized by the presence of clymenids and conodonts, but bivalves, ostracods, radiolaria, brachiopods, trilobites and crinoids are also observed in thin sections. On the basis of the conodont fauna, the mudstone/wackestone unit is shown to span the middle Frasnian (MN 11–12 Zones) to the middle Famennian (Upper marginifera Zone). In general, this unit passes upward with a disconformable contact to the Hochwipfel formation. The disconformity is emphasized by the sudden lithological change and by a hiatus that in the study area ranges from the Famennian (conodont data up to the Upper marginifera Zone) presumably up to the Visean, as in the whole Carnic area the start of the deposition of the Hochwipfel formation is thought to have occurred within the Visean (Venturini 1990). Still, the transition found west of Mt. Pizzul between the mudstone–wackestone (up to the Upper marginifera Zone), the breccia/nodular limestone (Lower expansa Zone) (Mossoni et al. 2013), the chert with limestone lenses (anchoralis-latus Zone) and the Hochwipfel formation suggests a more discontinuous history in correspondence of this boundary. The Pal Grande formation in the study area is about 20– 25 m thick. As for other units, the thickness can considerably change because of tectonic reduction. The Pal Grande formation reflects deposition under pelagic conditions (Mossoni et al. 2013). The reddish colored part of the Pal Grande formation in the study area ranges from the lower (Lower triangularis Zone) to the middle Famennian (Upper marginifera Zone). The red color might suggest well-oxygenated conditions in association with a particularly low sedimentation rate and diagenetic overprint (Bandel 1974). The breccia or nodular limestone at the top of the succession deserves some considerations. In general, clasts/ nodules do not appear to have been subjected to transport (Fig. 9f). They are not only very angular, but often

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their shape is complementary to the shape of the adjacent clasts like if they were deposited as mudstone/wackestone and then disrupted into breccia. This in turn limits the interpretations on the genetic origin of the deposit to an intraformational breccia or to a nodular pelagic limestone deposited in an extremely low sedimentation rate settling. The presence of some angular limits between the clasts/nodules as well as their different mutual dimensions (Fig. 9f) appears to be more consistent with a formation as intrabasinal breccia. The discussion on the processes responsible for the formation of such deposits will be deepened in the next chapters also by comparing data from other parts of the Carnic Alps (e.g., Spalletta et al. 1980, 1982; Spalletta 1983; Schönlaub et al. 1991; Spalletta and Venturini 1995). The following transition to cherts with interbedded limestone lenses and then to the Hochwipfel formation has been interpreted as evidence of the progressive deepening of the basin (Spalletta 1983). Hochwipfel formation The Hochwipfel formation crops out extensively across all of the study area (Fig. 3). The term Hochwipfel formation is—although informally—used in both Austrian and Italian literature after Heritsch (1928). It consists of very thin- to medium (very rarely thick)bedded yellowish gray lithic sandstones and graywackes interbedded with dark gray laminated shales. Locally, some levels of breccia with centimeter-scale cherty clasts have been found similar to what is observed in the rest of the Carnic Alps (e.g., Venturini et al. 2001, 2009). The sandstone/graywacke beds are frequently normally graded and sometimes show evidences of lamination and/or ripples. In the study area, fossil remains of plants which are frequently observed in other part of the Carnic Alps are quite rare, although locally present. The uppermost part of the Hochwipfel formation does not crop out in the study area, although the passage to the Dimon formation is visible just few kilometers south. According to Venturini et al. (2009), the Hochwipfel formation ranges from the middle Visean to the Bashkirian. The Hochwipfel formation as a whole has been interpreted as a flysch-type unit, partly made of turbiditic deposits (e.g., Spalletta et al. 1980; van Amerom and Schönlaub 1992).

Discussion The units of the lower Paleozoic of the Mt. Pizzul area have been mapped, measured, dated and described.

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The first aim of this work was to contribute to the procedure of formalization, even if the formalization itself will be finalized in a dedicated volume, of the lithostratigraphic units that are better exposed in the study area by evaluating their lithologic characters, stratigraphic relations and temporal distribution (Fig. 12); moreover, we propose the study area as reference area for the Hoher Trieb formation. This unit in fact changes remarkably from site to site (i.e., Schönlaub 1969, 1985; Pölsler 1969), not only in the overall thickness, but also for the processes affecting the lithoclastic bodies, the relative thickness and distribution of the different facies and the presence/absence of levels of chert/black shales intercalations in different intervals. As a consequence, several reference sections will be required to explain the unit complexity. As a higher-order aim of our study, the stratigraphic relations between the units cropping out in the area have been examined also in order to infer the depositional evolution of this sector of the Carnic Alps during the Upper Ordovician—Lower Carboniferous (Fig. 12) and its relation with the other parts of the Carnic basin. Moreover, we tried to investigate which response to global or supraregional controls can be recognized in the Carnic basin and which effects on its evolution depend instead on regional subsidence. The local basin was more or less uniform throughout the study area until the Eifelian, and then, it differentiated in shallow water, slope and slope–pelagic transition facies. This differentiation was possibly associated with a progressive change in the basin physiography from a ramp-type margin to a rimmed shelf margin or at least a margin with a slope separating the shallow water facies from the basin. The Upper Ordovician shows a transition from shale to sandy shale (Valbertad formation) and then nodular mudstone/wackestone with sandy shale to graywackes interlayered (Uqua formation) deposits. The Uqua formation (Figs. 7, 8a, 12) corresponds to the Boda Event (Berry and Boucot 1973; Berry 2003; Boucot et al. 2003). Fortey and Cocks (2005) interpreted the occurrence of limestone deposition interrupting older clastic sequences as evidence of an episode of global warming. This event—documented in many high latitude Gondwana settings—has been re-evaluated by Cherns and Wheeley (2007), who hypothesized that this transition might have been caused by a pre-Hirnantian icehouse interval. Sea-level fall would have led to cold water sinking at the south pole and to oceanic overturn through onset of thermohaline circulation, oxygenation of deep seafloors and local upwelling, with reduction in clastic input in response to expansion of terrestrial ice and shallowing that allowed cool-water carbonate facies to develop (Cherns and Wheeley 2007). This interpretation has been challenged by Armstrong et al. (2009) who dismissed the stratigraphic


reconstructions made by Cherns and Wheeley (2007). Still, the discussion about a pre-Hirnantian glacial interval is open. The possible presence of glacioeustasy or maybe some cooling during pre-Hirnantian times was, for example, recently shown to be consistent with isotopic data (Elrick et al. 2013). The debate is too complex and articulate to be solved from our study area and with our basic analyses. Nevertheless, some speculations can be offered, to add additional and possibly interesting perspectives to the discussion. The Uqua formation has been assigned to represent cool-water carbonates (Dullo 1992), but this does not automatically imply that this unit represents a cooler depositional environment than the older clastic shoreface to offshore deposits. Still, the crude coarsening upward trend within the Valbertad formation and within the clastic intervals of the Uqua formation (Fig. 7) suggests an increasing proximity of the source area. This trend is coupled with the increasing upward content—both in dimensions and amount—of limestone nodules and then layers (Fig. 7) which might suggest a shallowing upward succession. Although these elements alone are certainly not enough to recognize the warmer or cooler significance of these facies, evidences from the study area are consistent with a progressively reduced clastic input coupled with an increasing proximity of the source area. This might possibly be a consequence of the presence of ice sheets which caused sea-level drop and inhibited clastic deposition and would be consistent with isotopic data suggesting that orbitally controlled climatic changes acted in pre-Hirnantian times (Elrick et al. 2013), resulting in a protracted and dynamic late Ordovician greenhouse to icehouse transition. The effects of the Hirnantian glaciation are well exposed in the study area. We report for the first time the presence of the Plöcken formation in the Italian side of the Carnic Alps (Figs. 7, 8b, 12). This unit consists of a progradational shoreface sequence on top of the pelagic limestone of the Uqua formation, suggesting a dramatic drop of the sea level. In other parts of the Carnic Alps (i.e., Cellon section), the base of the Plöcken formation is marked by a level of diamictites (Schönlaub et al. 2011) which here is very poorly exposed. In most of the study area (but in most of the entire Carnic Alps as well), the sea-level drop led to progressive erosion also of the shoreface deposits of the Plöcken formation and the lowermost Uqua formation, thus suggesting that the Plöcken formation was deposited during an ongoing forced regression. The Hirnantia fauna that has been documented in the Carnic Alps (Schönlaub 1988; Berry 2003; Schönlaub et al. 2011) also strengthens this interpretation. The Carnic basin was not completely exposed during this episode, because continuous sedimentation has been reported in the most distal areas characterized by quartz sandstones/shale deposition (Schönlaub 1988) but the study

area was wholly subjected to subaerial erosion. The disconformity is documented on top of the Upper Ordovician carbonate succession throughout the entire Carnic Alps ranging up to the Wenlock and locally the Ludlow (e.g., Walliser 1964; Schönlaub 1980, 1988). Pelagic conditions were established from the following transgression starting probably during the late Hirnantian up to—in the study area—the Eifelian. Environmental conditions certainly changed through time as it is shown by changes in total thickness, shale content, fauna type and abundance, texture, layer thickness and color, but deposition remained below the storm wave base in a ramp-type margin. We found in fact no evidence of hummocky cross-stratification or graded beds, suggesting that reworking and re-deposition by storm action occurred in more proximal areas. Since a detailed layer by layer analysis was beyond the aim of this work, we cannot exclude the presence of some isolated bioclastic re-deposited layers. Still, we can exclude diffuse reworking such as the one recognized in some more proximal parts of the Rauchkofel formation in the area between Cellon and Pizzo Timau or in the Pragian Kellerwand formation (Kreutzer 1990). Also, we found no evidence of the so-called allodapic layers within the Findenig formation like the ones of Emsian age described in the relatively nearby area of Stua Ramaz (Vai 1980) or in more distant areas such as Oberbuchach (Schönlaub 1985). This again implies that the study area was—until the Eifelian—located in a pelagic part of the basin largely unaffected by storm and gravity-driven deposition in general. During the Eifelian, within the pelagic Findenig formation, the first evidence of gravity-driven flow occurs (sample BF, Figs. 11, 12). It consists of a meter-scale layers of matrix (grainstone) supported breccia passing distally to a decimetric-thick layer of grainstone. These characteristics suggest deposition by a selective process, possibly a grain flow. Later in the Eifelian, this part of the basin differentiated (Fig. 12) in shallow water facies including stromatoporoid and coral reefs (Zermula massif) passing distally to gravity-driven flows of fore reef slope environment (Vinz formation) and then to gravity-driven flows interlayered with pelagic deposition (Hoher Trieb formation) deposited in a slope to pelagic transition setting. Gravity-driven flows consist mostly of hyperconcentrated to concentrated density flow and turbidite flow deposits. Although it is difficult to precisely estimate the overall thickness of the shallow water facies in this part of the basin, due to tectonic reductions, a thickness of about 250 m can be estimated for the shallow water Eifelian?– Frasnian succession. This value is in agreement with the reconstructions made by Ferrari and Vai (1966). Laterally, this unit passes to the Vinz formation, whose thickness can also not be measured due to tectonic faulting (Fig. 10).

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More distally, this lateral facies transition correlates with the very uppermost part of the Findenig formation and the Hoher Trieb formation, reaching a maximum thickness of about 50 m. These data emphasize the dramatic physiographic change that occurred in a portion of the basin that—until the Emsian—was most probably uniform. The gradual appearance of gravity-driven deposits implies the progressive development of a slope-type setting. The gravity-driven deposits within the Findenig, Vinz and Hoher Trieb formations show reef-derived clasts, including corals and stromatoporoids (Figs. 10, 11). Accordingly, these units can give an indirect hint about the stratigraphic range of the reef evolution in the Mt. Zermula area. Breccia bed 1 (Findenig formation) represents the first evidence of the presence of a shallow water body, and it was attributed to the Eifelian (sample ZMB BF: Fig. 12). In the western Carnic Alps, in the Mt. Coglians/Hohe Warte area (Fig. 2), patch reefs are reported to develop from the Pragian (e.g., Vai 1963; Bandel 1969; Suttner 2007) and an about 1-km-thick shallow water succession of Pragian–Givetian age is present. The base of the Mt. Zermula shallow water facies is unknown because they have been faulted out, so we cannot exclude the presence of small mounds, patch reefs or in general a depositional high. Still, there are no reflections— i.e., reworked sediments—of such bodies within the basin in the study area; therefore, we assume that there were no major shallow water facies in the Mt. Zermula area before the Eifelian and that the margin was ramp-type with a transition to a relatively shallow basin. This is also consistent with data from the Emsian succession (Findenig formation) of the relatively nearby area of Stua Ramaz (Vai 1980), where the lithoclastic beds within the pelagic deposits have been interpreted as tempestites (Vai 1980). Tempestites would be consistent with a ramp-type margin. In the Findenigkofel (Pölsler 1969) and Oberbuchach II (Schönlaub 1985) sections, all the lithoclastic deposits (of Emsian to Eifelian age) within the pelagic mudstone/wackestone have been instead interpreted as re-deposition in a deepwater basin (Bandel 1974; Schönlaub 1985). The reef and in general shallow water facies have been never shown to prograde into more distal part of the basin, but were restricted to part of the study area. Sea-level fluctuations are reflected in intercalations between different forereef breccia and pelagic facies (Fig. 11). In the whole Carnic Alps, the Eifelian–Frasnian shallow water units appeared have occurred only in well-defined, limited areas (e.g., Kreutzer 1990, 1992); intercalations with the slope facies are not reported. This suggests that the transition between the shallow water and the forereef slope facies might have been possibly occurred through a non-depositional bypass margin. There are no data within the study area to quantify the basin depth after the emplacement of

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the slope, but certainly the basin became deeper, as suggested by Bandel (1974), although we do not have elements to quantify the actual depth. The reconstruction of the architecture of the slope bodies is beyond the aim of this work and necessitates much more detailed analyses on a much wider area. Still, some considerations can be proposed. Breccia beds from the Findenig and Hoher Trieb formations change their thickness as well as their lithology and/or texture distally becoming thinner and finer grained. The directions through which this happens are not consistent from bed to bed. Breccia level I (Findenig formation) thins from the Forca di Lanza toward the Zuc di Malaseit area, while breccia levels III and IV (Hoher Trieb formation) appear to reach their maximum thickness at the Zuc di Malaseit. The breccia levels can be followed along the whole study area and thus represent single depositional episodes. They might derive from a line source or multiple source points rather than a single point. Apart from these complexities, the Vinz formation deposits pass distally to the Hoher Trieb formation (Fig. 12). The Vinz formation consists of stacked gravitydriven deposits (mainly hyperconcentrated and concentrated flows) without evidence of pelagic deposition if not for iron-rich nodules reworked within some of the breccia beds. The Hoher Trieb formation consists of gravitydriven deposits (hyperconcentrated and concentrated flows alternating with fine-grained turbidites) interfingered with pelagic deposits (cherts and black shales). Cherts in particular are present along the whole lower part of the Hoher Trieb formation but are particularly extensive in the interval between breccia levels 3 and 4 (Fig. 11). Here, they are associated—in addition to turbiditic deposits—with some levels of black shales (Fig. 12). This interval spans the Eifelian–Givetian transition. A conodont fauna belonging to the kockelianus Zone has been documented just below a black shale bed (Fig. 12), equivalent to section Oberbuchach II (Schönlaub 1985). We hypothesize that these black shales might possibly correspond to the Kacˇák Event (e.g., Chlupac and Kukal 1986; House 1996, 2002; Kido and Suttner 2011; Kido et al. 2012). The pelagic intervals likely represent transgressive pulses strong enough to temporally inhibit shallow water carbonate production and consequent gravity-driven deposition. The areas where the Vinz formation was deposited were probably starved during transgressions, which might possibly explain the formation of iron-rich nodules later reworked by the younger re-sedimentation episode. The identification of the different transgressive pulses and of their order up to the higherfrequency episodes (Ellwood et al. 2011) necessitates more detailed analyses which are beyond the aims of this study. The reefs appear—from indirect evidences of the foreslope facies evolution—to increasingly expand—with higher-frequency order fluctuations within this broad


trend—from the Eifelian to the Givetian (Fig. 12). This is in agreement with what observed in the western Carnic Alps in the Mt. Coglians/Hohe Warte massif (e.g., Vai 1976; Kreutzer 1990; Pas et al. 2014), but also globally (Johnson et al. 1985; Copper and Scotese 2003). Apart from the different timing of the shallow water facies emplacement between the Mt. Coglians/Hohe Warte and Zermula area, reef evolution between Eifelian and Frasnian times appears to have been coarsely the same, which implies that although we cannot state whether these shallow water bodies were connected or distinguished as atoll-like reef bodies, they were subjected to the same controls (at least the high order controls). In comparison with our data, Kreutzer (1990) concluded for the central Carnic Alps a barrier-type reef model. In contrast, Tessensohn (1974) and Rantitsch (1992) preferred for the eastern continuation of the Carnic Alps, Karavanke Alps, isolated reef bodies. Moreover, the reef progressive extension until the Givetian and the presence of a global event such as the Kacˇák Event suggest that the study area was at least partially subjected to global or at least supra-regional controls (i.e., sea-level fluctuations). The transition between the Hoher Trieb and the Pal Grande formations is marked by a sharp change in lithology and by the presence of a slump close to the base of the Pal Grande formation (Fig. 11), which occurs between the Frasnian MN3 (sample FL6) and MN 12 (sample FL5) conodont zones. It records the sharp end of reef-fed gravity-driven deposition and the onset of pelagic conditions in this part of the basin. This most probably reflects the drowning of at least part of the reef, of which erosional products cannot reach this part of the basin anymore. Data from the study area are consistent with evidences from other parts of the Carnic Alps which suggest that reef drowning started earlier in the Carnic Alps than in other nearby basins (Vai 1976; Spalletta et al. 1980, 1982; Kreutzer 1990; Pas et al. 2014). Local extensional or transtensional tectonics has been claimed to control reef early drowning (Spalletta and Vai 1984). An extensional setting is consistent with the slump found close to the base of the Pal Grande formation that implies basin instability as well as with the sudden change in the depositional pattern. It would also explain the early drowning of the reefs in most of the Carnic Alps respect to the global reef extinction that occur at the Frasnian–Famennian boundary (e.g., Copper 2002). The Pal Grande formation in this part of the basin is quite monotonous with the exception of a distinct change of color from light gray to moderate pink and red and then again to light gray (Fig. 12). Reddish deposits (Fig. 9e) roughly range from the Lower triangularis to the Upper marginifera conodont zone. This facies has been related to well-oxygenated conditions in association with particularly low sedimentation rate and diagenetic overprint (Bandel

1974). We assume that these conditions might have been possibly associated with the transgressive trend documented from the triangularis to the marginifera conodont Superzones—with higher-order regressive fluctuations in this interval—in the whole Southern Europe (Buggisch and Joachimski 2006) and also globally (Johnson et al. 1985). This hypothesis is consistent with data from conodont biofacies from the whole Carnic Alps (Perri and Spalletta 2000). The transition to the following Hochwipfel formation in the study area is disconformable. It occurs either with a relatively large hiatus (middle Famennian Upper marginifera conodont zone to ?Visean) or with a passage between mudstone–wackestone (‘typical Pal Grande formation’: up to the Upper marginifera Zone), breccia/nodular limestone (Lower expansa Zone) (Mossoni et al. 2013), chert with limestone lenses (anchoralis-latus Zone) and finally the Hochwipfel formation. The breccia/nodular limestone (Figs. 9f, 12) is equivalent to deposits found in patches in the topmost part of the Pal Grande formation in other parts of the Carnic Alps such as Oisternig, Schönwipfel, Creta di Collinetta/Cellon and Pal Grande (Spalletta et al. 1980; Schönlaub et al. 1991). The geometry of such facies in the study area is not observable due to very poor exposure conditions, but in other parts of the Carnic Alps the lower boundary is very irregular, locally forming centimeter- to several meter-scale wide and several meter-scale long crevices, interpreted as neptunian dykes by Spalletta et al. (1980). In the study area, this facies has been constrained to the Lower expansa conodont zone (Mossoni et al. 2013) (Fig. 12), while in other parts of the Carnic Alps can be also Tournasian and Visean (Schönlaub et al. 1991; Perri and Spalletta 1998c). Such depositional geometry and facies—including textural characteristics like the presence of angular clasts and of clasts of very different diameter (sub-centimeter to decimeter wide)—suggest a formation as intraformational breccia rather than nodular pelagic limestone. Intraformational breccia might reflect formation related to soft-sediment deformation at a shallow burial depth (e.g., Chen et al. 2009), syn-tectonic deformation (e.g., Lehner 1991) or subsoil paleokarst (collapse breccia or rubble breccia resulting from weathering of bedrock and cementation in the vadose zone) (e.g., Edwars Clifton 1967; Szulczewski et al. 1996). Soft-sediment deformation can most probably be excluded in the study area because of the irregular geometry of the breccia beds, the limited presence of shales (necessary to provide injection of fluid materials) and the occurrence in a single stratigraphic interval. Spalletta (1983), Spalletta et al. (1980) and Spalletta and Venturini (1988) proposed a tectonically driven deepening of the whole Carnic basin with the deposition of olistolites and olistostromes at the base of the Hochwipfel formation turbidites. The breccia

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facies at the top of the Pal Grande formation might indeed reflect syn-sedimentary tectonic deformation, but this interpretation alone fails to explain the hiatus between the Pal Grande formation (middle Famennian) and the Hochwipfel formation (Visean) which is found in the study area where the breccia facies is not present. Instead, both the textural characteristics and the geological context are consistent with a formation of the breccia levels as paleokarst breccia as suggested in other areas by Schönlaub et al. (1991). The facies might correspond to either collapse breccia (Schönlaub et al. 1991) or rubble breccia resulting from weathering of bedrock and cementation in the vadose zone. Spalletta and Venturini (1995) taking into account the data presented by Schönlaub et al. (1991) proposed for the Carnic basin a scenario with a strike-slip geodynamic setting including gravity-driven tectonically induced resedimentation and the presence of tectonically rising blocks up to subaerial exposure. The outcrop is not well exposed (e.g., the basal contact is hidden), but the texture of the breccia (which suggest deposition and then fragmentation) and the absence of materials coming from younger parts of the succession appear to be more consistent with an interpretation as rubble breccia resulting from weathering of bedrock. The hypothesis of a subaerial exposure must of course be tested with more detailed analysis across the entire basin; this, however, might explain the patchy distribution of the unit (not only in the study area but in the whole Carnic Alps) and its stratigraphic distribution ranging, depending on the zones, between the Famennian and the Visean, in addition to the hiatus recorded on top of the shallow water facies where generally Frasnian rocks are disconformably superposed by the pelagic Tournaisian and/or Visean deposits of the Pal Grande formation. This occurs also in the Mt. Zermula area (Manzoni 1966), where shallow water facies of presumable Frasnian age are disconformably covered by Lower Carboniferous deposits of the Pal Grande formation. The existence of a subaerial exposure of Lower Carboniferous age has been proposed by Schönlaub et al. (1991) and Brigo et al. (2001) on the basis of the presence of siliceous crust-type mineralization between the Pal Grande and the Hochwipfel formations, in contrast to the genesis of an unconformity associated with deepening controlled by extensional or transtensional setting proposed by Spalletta (1983) and Spalletta et al. (1980). Still the transtensional setting—including coeval rising blocks, fault scarps and subsiding basins—might explain the existence of localized exposure (Spalletta and Venturini 1988, 1995; Venturini and Spalletta 1998), especially if found of different ages in different localities. We hypothesize the presence of different trends at a different scale. Progressive deepening starting from the Frasnian and reaching its climax in the Visean with the

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establishment of turbidite deposition—possibly related to extensional or transtensional tectonic—was superposed by higher-frequency fluctuations which caused sea-level drops and possibly two subaerial exposures in the study area. Sea-level drops—locally leading to exposure—have been reported worldwide before and after the Hangenberg Event, close to the Devonian/Carboniferous boundary (e.g., Kaiser et al. 2011; Myrow et al. 2013; De Vleeschouwer et al. 2013). De Vleeschouwer et al. (2013) associated these sealevel fluctuations with the presence of small continental ice sheets on Western Gondwana. We assume that a sea-level drop might have occurred in the study area leading to karstification—including breccia formation (collapse breccia or rubble breccia resulting from weathering of bedrock and cementation in the vadose zone)—and following erosion of the uppermost Devonian beds down to the Lower expansa Zone (breccia) or Upper marginifera Zone. Then, the area was drowned again with the deposition of interlayered chert and limestone during at least part of the Tournaisian. The transition to the following Hochwipfel formation is not exposed; therefore, we do not have data to assess whether there is evidence consistent with the subaerial exposure proposed by Schönlaub et al. (1991) and Brigo et al. (2001).

Conclusions The Mt. Pizzul area represents a strategic area to understand the geology of the Paleozoic of the Carnic Alps in the current stage of the knowledge. The sedimentation was mostly continuous between the Upper Ordovician and the Lower Carboniferous; erosion and reworking were limited; the environments record well the basin dynamics; most of the units are conodont bearing, although megafossils still await a detailed study; tectonic overprint—although present—can be constrained. We performed a detailed geological mapping and a stratigraphic/biostratigraphic survey in order primarily to (1) contribute to the procedure of formalization of the lithostratigraphic units exposed in the study area and (2) understand the depositional evolution and deformation history of the study area; as a secondorder aim, we tried (3) to infer the relations with other parts of the Carnic basin and (4) distinguish local and regional controls on sedimentation. In the framework of the formalization process, we report for the first time the occurrence of the Plöcken formation on the Italian side of the Carnic Alps and present a possible reference section for the Hoher Trieb formation. Moreover, we recognized the stratigraphic relations between the units by providing biostratigraphic constraints in close association with most of the formational boundaries. The area underwent a multistage deformational history, most of which are related to Variscan compression. This led


to the formation of a kilometer-scale recumbent fold and to the Tortonian–Serravallian Alpine transpressional phase that caused most of the Alpine shortening. The main characters of the depositional evolution of the study area can be sketched as follows: pre-Hirnantian deposition in a ramp-type margin; Hirnantian glacioeustaticrelated deposits and unconformity; Upper Silurian–Eifelian interval of offshore/pelagic deposition in a ramp-type margin; formation of a slope and associated differentiation of the basin in shallow water (including reefs), transitional (gravity-driven deposits) and pelagic parts (Eifelian–Frasnian); transgression of pelagic limestone and reef drowning from the end of the Frasnian, likely associated to local extensional or transtensional tectonics; possible existence of two subaerial exposures close to the Devonian– Carboniferous boundary and Visean, respectively, that might represent higher-frequency regressive fluctuations in a Frasnian–Visean context of progressive drowning and turbidite deposition during the Visean. The study area is located in the vicinity of the Zermula massif, where reef growth and related depositional environments have been reported but precise timelines are missing because of tectonic cuttings and difficulty to precisely date shallow water units. We used the indirect evidence of the stratigraphic distribution of the material derived from the shallow water units to infer the stratigraphic distribution of the shallow water units themselves. The basin started to receive material from shallow waters in the Eifelian (uppermost part of the Findenig formation). Accordingly, we infer that the Zermula massif likely started to develop during the Eifelian. The Mt. Coglians massif (the best known among the shallow water units in the Carnic Alps) instead shows the presence of patch reefs associated with development of extensive transitional facies beginning at least by the Pragian. Although there is no evidence to prove whether the shallow water bodies were connected or not, they started their development at different times, although the subsequent evolution, including the physiographic change from ramptype to rimmed shelf and up to the drowning beginning in the Frasnian, appears to be very similar and possibly coeval. The deposits of the study area reflect the presence of global or supra-regional controls such as the Boda Event, the Hirnantian glaciation, the Middle Devonian reef growth, and the Kacˇák Event, the high-frequency sea-level fluctuations close to the Devonian–Carboniferous boundary. The drowning of the Zermula buildup (and most probably of all the reef bodies in the Carnic Alps) instead appears to have started during the Frasnian, unlike what has been observed in other areas in the world where reefs continue to flourish throughout the Frasnian. This fact— coupled with some features consistent with the presence of syn-sedimentary tectonics such as a slump and a sharp formational boundary—suggests the presence of extensional

or transtensional tectonics beginning in the Frasnian, leading to progressive drowning of the basin up to the transition to turbidite deposition of the Hochwipfel formation. Acknowledgments We are indebted with all the members of the ‘Formal Lithostratigraphic Units in the Pre-Variscan Sequence of the Carnic Alps’ science group for fruitful, open and enthusiastic discussions. We warmly thank Hans Peter Schönlaub and Charles Ver Straeten for their thoughtful and stimulant reviews, which resulted in a greatly improved manuscript. MP research was founded by the Italian Ministry of Education, University and Research. EK and TJS are grateful to the Austrian Science Fund for financial support (Project: FWF P23775-B17). Researches by CC, MGC and AM were partly supported by grants RAS. This paper is a contribution to IGCP 591 (The Early to Middle Paleozoic revolution) and 596 (Climate change and biodiversity patterns in the Mid-Paleozoic). We wish to warmly thank Francesca, Silvio and all the wonderful people of the Lanza Hut for their friendship and hospitality.

References Armstrong HA, Baldini J, Challands TJ, Gröcke DR, Owen AW (2009) Response of the inter-tropical convergence zone to southern hemisphere cooling during upper ordovician glaciation. Paleogeogr Paleoclimatol Paleoecol 284(3):227–236 Bagnoli G, Ferretti A, Serpagli E, Vai GB (1998) Late Ordovician conodonts from the Valbertad Section (Carnic Alps). In: Perri MC, Spalletta C (eds) Southern Alps Field Trip Guidebook. In: Seventh international conodont symposium held in Europe, 1998. Giorn Geol, Special Issue, vol 60, pp 138–149 Bandel K (1969) Feinstratigraphische und biofazielle Untersuchungen unterdevonischer Kalke am Fuß der Seewarte (Wolayer See, zentrale Karnische Alpen). Jb Geol B-A 112:197–234 Bandel K (1972) Palökologie und Paläogeographie im Devon und Unterkarbon der zentralen Karnischen Alpen. Palaeontogr Ab A 141(1–4):1–117 Bandel K (1974) Deep-water limestones from the Devonian-Carboniferous of the Carnic Alps, Austria. In: Hsü KJ, Jenkyns HC (eds) Pelagic sediments: on land and under the Sea. IAS Spec Publ, vol 1. Wiley, New York, pp 93–116 Berry WBN (2003) Late Ordovician environmental changes in Carnic Alps and central Nevada: a comparative study. Bull Soc Geol Fr 174(3):211 Berry WBN, Boucot AJ (1973) Glacio-Eustatic control of late Ordovician-early silurian platform sedimentation and faunal changes. Geol Soc Am Bull 84(1):275–284 Boucot AJ, Jia-Yu R, Xu C, Scotese CR (2003) Pre-Hirnantian Ashgill climatically warm event in the Mediterranean region. Lethaia 36(2):119–131. doi:10.1080/00241160310001245 Brett CE, Ferretti A, Histon K, Schönlaub HP (2009) Silurian sequence stratigraphy of the Carnic Alps, Austria. Paleogeogr Paleoclimatol Paleoecol 279(1–2):1–28 Brigo L, Camana G, Rodeghiero F, Potenza R (2001) Carbonatehosted siliceous crust type mineralization of Carnic Alps (Italy–Austria). Ore Geol Rev 17(4):199–214. doi:10.1016/ s0169-1368(00)00014-7 Brime C, Perri MC, Pondrelli M, Spalletta C, Venturini C (2008) Polyphase metamorphism in the eastern Carnic Alps (N Italy–S Austria): clay minerals and conodont Colour Alteration Index evidence. Int J Earth Sci 97(6):1213–1229 Buggisch W, Joachimski MM (2006) Carbon isotope stratigraphy of the Devonian of Central and Southern Europe. Paleogeogr Paleoclimatol Paleoecol 240(1–2):68–88. doi:10.1016/j.palaeo.2006.03.046

13

Carta N (2011) Caratterizzazione Biostratigrafica delle Unità Litologiche affioranti nell’area di Monte Pizzul, Passo del Cason di Lanza (Alpi Carniche). Master Thesis, University of Cagliari, 130 pp Chen J, Chough SK, Chun SS, Han Z (2009) Limestone pseudoconglomerates in the Late Cambrian Gushan and Chaomidian Formations (Shandong Province, China): soft-sediment deformation induced by storm-wave loading. Sedimentology 56(4):1174–1195 Cherns L, Wheeley JR (2007) A pre-Hirnantian (Late Ordovician) interval of global cooling—the Boda event re-assessed. Paleogeogr Paleoclimatol Paleoecol 251(3–4):449–460. doi:10.1016/j.palaeo.2007.04.010 Chlupac I, Kukal Z (1986) Reflection of possible global Devonian events in the Barrandian area, CSSR. In: Global bio-events Proc 1st meet IGCP Project 216, pp 169–179 Copper P (2002) Reef development at the Frasnian/Famennian mass extinction boundary. Paleogeogr Paleoclimatol Paleoecol 181(1–3):27–65 Copper P, Scotese CR (2003) Megareefs in Middle Devonian supergreenhouse climates. Geol Soc Am Spec Pap 370:209–230 Corradini C, Corriga MG (2010) Silurian and lowermost Devonian conodonts from the Passo Volaia area (Carnic Alps, Italy). Boll Soc Paleont It 49(3):237–253 Corradini C, Corriga MG (2012) A Prˇídolí-Lochkovian conodont zonation in Sardinia and the Carnic Alps: implications for a global zonation scheme. Bull Geosci 87(4):635–650. doi:10.3140/bull.geosci.1340 Corradini C, Simonetto L, Serventi P, Rigo R, Calligaris C (2005) Loboliti (Crinoidea) del Devoniano basale di Monte Zermula (Alpi Carniche Italiane). Rend Soc Paleont It 2:29–36 Corradini C, Pondrelli M, Corriga MG, Simonetto L, Kido E, Suttner TJ, Spalletta C, Carta N, Geology and stratigraphy of the Cason di Lanza area (Mount Zermula, Carnic Alps, Italy) (2012). In: Kido E, Suttner TJ, Piller WE, Da Silva AC, Corradini C, Simonetto L (eds) IGCP 580—4th annual meeting, 24–30th June 2012, Graz, 2012. Ber Inst Erdw, Karl-Franzens-Universität Graz, 17, pp 83–103 Corradini C, Corriga MG, Männik P, Schönlaub H-P (2014) Revised conodont stratigraphy of the Cellon section (Silurian, Carnic Alps). Lethaia, 16 pp. doi:10.1111/let.12087 Corriga MG (2011) Biostratigrafia a conodonti attorno al limite Siluriano-Devoniano in alcune aree del Nord Gondwana. Ph.D., Università di Cagliari Corriga MG, Corradini C (2009) Upper Silurian and lower Devonian conodonts from the Monte Cocco II Section (Carnic Alps, Italy). Bull Geosci 84(1):155–168. doi:10.3140/bull.geosci.1112 Corriga MG, Suttner TJ, Kido E, Corradini C, Pondrelli M, Simonetto L (2011) The age of the La Valute limestone-Findenig limestone transition in the La Valute Section (Lower Devonian, Carnic Alps, Italy). Gortania Geol Paleont Paletn 32:5–12 De Vleeschouwer D, Rakocin´ski M, Racki G, Bond DPG, Sobien´ K, Claeys P (2013) The astronomical rhythm of Late-Devonian climate change (Kowala section, Holy Cross Mountains, Poland). Earth Planet Sci Lett 365:25–37 Dullo W (1992) Mikrofazies und Diagenese der oberordovizischen Cystoideen-Kalke(Wolayerkalk) und ihrer Schuttfazies (Uggwakalk) in den Karnischen Alpen. Jb Geol B-A 135(1):317–333 Edwars Clifton H (1967) Solution-collapse and cavity filling in the Windsor Group, Nova Scotia, Canada. Geol Soc Am Bull 78(7):819–832 Ellwood BB, Algeo TJ, El Hassani A, Tomkin JH, Rowe HD (2011) Defining the timing and duration of the Kacˇák Interval within the Eifelian/Givetian boundary GSSP, Mech Irdane, Morocco, using geochemical and magnetic susceptibility patterns. Paleogeogr Paleoclimatol Paleoecol 304(1–2):74–84

13

Int J Earth Sci (Geol Rundsch) Elrick M, Reardon D, Labor W, Martin J, Desrochers A, Pope M (2013) Orbital-scale climate change and glacioeustasy during the early Late Ordovician (pre-Hirnantian) determined from δ18O values in marine apatite. Geology 41(7):775–778. doi:10.1130/g34363.1 Ferrari A, Vai GB (1966) Ricerche stratigrafiche e paleoecologiche al M. Zermula. Giorn Geol 33:389–416 Ferretti A, Schönlaub HP (2001) New conodont faunas from the Late Ordovician of the Central Carnic Alps, Austria. Boll Soc Paleont It 40(1):3–15 Fortey RA, Cocks LRM (2005) Late Ordovician global warming—the Boda event. Geology 33(5):405–408. doi:10.1130/g21180.1 Frech F (1887) Über das Devon der Ostalpen, nebst Bemerkun- gen über das Silur und einen palaeontologischen Anhang. Zeit Deutsch Geol Gesellschaft 39:659–738 Frech F (1894) Die Karnischen Alpen. Ein Beitrag zur vergleichenden Gebirgs-Tektonik. Abh Natur Gesellsch Halle 1–514 Gephart JW (1990) FMSI: a fortran program for inverting fault/ slickenside and earthquake focal mechanism data to obtain the regional stress tensor. Comput Geosci 16(7):953–989 Gephart JW, Forsyth DW (1984) An improved method for determining the regional stress tensor using earthquake focal mechanism data: application to the San Fernando earthquake sequence. J Geophys Res 89(B11):9305–9320. doi:10.1029/JB089iB11p09305 Geyer G (1894) Zur Stratigraphie der paläozoischen Schichtserie in den Karnischen Alpen. Verh Geol R-A:102–119 Gortani M (1908) Nuove ricerche sul Nucleo centrale delle Alpi Carniche. Rend R Acc Lincei Classe di Scienze Fisiche Matematiche e Naturali 5(17):603–612 Gortani M (1920) I bacini della But, del Chiarsò e della Vinàdia in Carnia. Geologia, morfologia, idrografia. Carta della permeabilità delle acque. Pubbl Uff Idrogr R Mag Acque 104:1–71 Hammarlund EU, Dahl TW, Harper DAT, Bond DPG, Nielsen AT, Bjerrum CJ, Schovsbo NH, Schönlaub HP, Zalasiewicz JA, Canfield DE (2012) A sulfidic driver for the end-Ordovician mass extinction. Earth Planet Sci Lett 331:128–139 Heritsch F (1928) Tektonische Fragen im Karbon der Karnischen Alpen. Sitzber Akad Wiss Wien, math nat kl 137:303–338 Heritsch F (1936) Die Karnischen Alpen: monographie einer Gebirgsgruppe der Ostalpen mit variszischem und alpidischem Bau. Hrsg Geol Inst 1–205 House MR (1996) The middle Devonian Kacˇák event. Geosc S-W Engl 9(1):79–84 House MR (2002) Strength, timing, setting and cause of midPalaeozoic extinctions. Paleogeogr Paleoclimatol Paleoecol 181(1–3):5–25 Hubmann B, Pohler S, Schönlaub H, Messner F (2003) Paleozoic coral-sponge bearing successions in Austria. Ber Geol B-A 61:1–91 Hubmann B, Ebner F, Ferretti A, Histon K, Kido E, Krainer K, Neubauer F, Schönlaub HP, Suttner TJ (2013) The lithostratigraphic units of the Austrian Stratigraphic Chart 2004 (sedimentary successions): vol. I. Abh Geol B-A 66:137 Jaeger H (1975) Die Graptolithenführung im Silur/Devon des CellonProfils (Karnische Alpen). Carinthia II 165(185):111–126 Jaeger H, Schönlaub HP (1977) Das Ordoviz/Silur-Profil im Nölblinggraben (Karnische Alpen, Österreich). Verh Geol B-A 1977:349–359 Joachimski MM, Buggisch W (2002) Conodont apatite δ18O signatures indicate climatic cooling as a trigger of the Late Devonian mass extinction. Geology 30(8):711–714. doi:10.1130/0091-7613 (2002)0302.0.co;2 Joachimski MM, Pancost RD, Freeman KH, Ostertag-Henning C, Buggisch W (2002) Carbon isotope geochemistry of the Frasnian-Famennian transition. Paleogeogr Paleoclimatol Paleoecol 181(1–3):91–109. doi:10.1016/s0031-0182(01)00474-6

Int J Earth Sci (Geol Rundsch) Joachimski MM, Breisig S, Buggisch W, Talent JA, Mawson R, Gereke M, Morrow JR, Day J, Weddige K (2009) Devonian climate and reef evolution: insights from oxygen isotopes in apatite. Earth Planet Sci Lett 284(3–4):599–609. doi:10.1016/j.epsl.2009.05.028 Johnson ME (2006) Relationship of Silurian sea-level fluctuations to oceanic episodes and events. GFF 128(2):115–121. doi:10.1080/11035890601282115 Johnson JG, Klapper G, Sandberg CA (1985) Devonian eustatic fluctuations in Euramerica. Geol Soc Am Bull 96(5):567–587. doi:10.1130/0016-7606(1985)962.0.co;2 Kaiser SI, Steuber T, Becker RT, Joachimski MM (2006) Geochemical evidence for major environmental change at the DevonianCarboniferous boundary in the Carnic Alps and the Rhenish Massif. Paleogeogr Paleoclimatol Paleoecol 240(1–2):146–160. doi:10.1016/j.palaeo.2006.03.048 Kaiser SI, Becker RT, Steuber T, Aboussalam SZ (2011) Climate-controlled mass extinctions, facies, and sea-level changes around the Devonian–Carboniferous boundary in the eastern AntiAtlas (SE Morocco). Paleogeogr Paleoclimatol Paleoecol 310 (3–4):340–364. doi:10.1016/j.palaeo.2011.07.026 Kido E, Suttner T (2011) A new project has been launched: FWF P23775-B17 “Late Eifelian climate perturbations: Effects on tropical coral communities”. Jb Geol B-A 151(3–4):407–416 Kido E, Suttner TJ, Pondrelli M, Corradini C, Corriga MG, Simonetto L, Vodrážková S, Koptíková L (2012) Geochemical and geophysical records of the Middle Devonian section in the Carnic Alps. Ber Inst Erdwiss K-F Univ Graz 17:29–30 Kreutzer L (1990) Mikrofazies, Stratigraphie und Paläogeographie des Zentralkarnischen Hauptkammes zwischen Seewarte und Cellon. Jb Geol B-A 133(2):275–343 Kreutzer L (1992) Palinspastische Entzerrung und Neugliederung des Devons in den Zentralkarnischen Alpen aufgrund von neuen Untersuchungen. Jb Geol B-A 135(1):261–272 Lamberty P (2013) Sédimentologie du Dévonien dans les Alpes Carniques (entre l’Autriche et l’Italie). Université de Liège Läufer AL (1996) Variscan and Alpine tectonometamorphic evolution of the Carnic Alps (Southern Alps)-structural analysis, illite cristallinity, K–Ar and Ar–Ar geochronology. Tübinger Geow Arb:40 Läufer AL, Frisch W, Steinitz G, Loeschke J (1997) Exhumed faultbounded Alpine blocks along the Periadriatic lineament: the Eder unit (Carnic Alps, Austria). Int J Earth Sci 86(3):612–626 Lehner BL (1991) Neptunian dykes along a drowned carbonate platform margin: an indication for recurrent extensional tectonic activity? Terra Nova 3(6):593–602 Manzoni M (1966) Conodonti neodevonici ed eocarboniferi al Monte Zermula. Giorn Geol 33:461–488 Mossoni A, Corradini C, Pondrelli M (2013) Famennian (Late Devonian) conodonts from the Pizzul West section (Carnic Alps, Italy). Gortania Geologia Paleontologia Paletnologia 34:13–34 Mulder T, Alexander J (2001) The physical character of subaqueous sedimentary density flows and their deposits. Sedimentology 48(2):269–299. doi:10.1046/j.1365-3091.2001.00360.x Myrow PM, Hanson A, Phelps AS, Creveling JR, Strauss JV, Fike DA, Ripperdan RL (2013) Latest Devonian (Famennian) global events in western Laurentia: variations in the carbon isotopic record linked to diagenetic alteration below regionally extensive unconformities. Paleogeogr Paleoclimatol Paleoecol 386: 194–209 Pas D, Da Silva A-C, Suttner T, Kido E, Bultynck P, Pondrelli M, Corradini C, De Vleeschouwer D, Dojen C, Boulvain F (2014) Insight into the development of a carbonate platform through a multi-disciplinary approach: a case study from the Upper Devonian slope deposits of Mount Freikofel (Carnic Alps, Austria/Italy). Int J Earth Sci 103:519–553. doi:10.1007/ s00531-013-0969-2

Perri MC, Spalletta C (1998a) Updating of the conodont biostratigraphy in the Carnic Alps (Italy). In: Perri MC, Spalletta C (eds) Southern Alps field trip guidebook. Seventh international conodont symposium held in Europe, 1998. Giorn Geol, Sp Issue 60:116–119 Perri MC, Spalletta C (1998b) Conodont distribution at the Tournaisian/Visean boundary in the Carnic Alps (Southern Alps, Italy). In: Szaniawski H (ed) Proceedings of the sixth European conodont symposium (ECOS VI), Palaeont Pol 58:225–245 Perri MC, Spalletta C (1998c) Latest Devonian and Early Carboniferous conodonts from the Casera Collinetta di Sotto A section (Carnic Alps, Italy). In: Perri MC, Spalletta C (eds) Seventh international conodont symposium held in Europe, 1998. Giorn Geol Sp Issue 60:168–181 Perri MC, Spalletta C (2000) Late Devonian-early carboniferous transgressions and regressions in the Carnic Alps (Italy). Rec West Austr Mus Sup 58:305–319 Perri MC, Spalletta C (2001) Hangenberg Event al limite Devoniano/Carbonifero al Monte Zermula, Alpi Carniche, Italia. In: Perri MC (a cura di), Giornate di Paleontologia 2001, Giorn Geol, Ser 3a, 62, 2000, Suppl:31–40 Perri MC, Pohler S, Pondrelli M, Spalletta C, Schnellbacher P, Suttner TJ (2009) The Pragian–Frasnian ‘transitional’ facies of the Carnic Alps (Austria–Italy). In: Suttner TJ, Hubmann B, Piller WE (eds) Paleozoic Seas symposium, Ber Inst Erdw Karl-FranzensUniversität Graz 14:58–59 Pölsler P (1969) Stratigraphie und Tektonik im Nordabfall des Findenigkofels (Silur bis Karbon; Karnische Alpen, Österreich). Jb Geol B-A 112:355–398 Pondrelli M (1998) Evoluzione deformativa del nucleo ercinico carnico centro-orientale. Università di Bologna Rantitsch G (1992) Fazies und Diagenese devonischer Riffkalke des Seeberger Aufruches (Kärnten, Österreich). Jb Geol B-A 135:273–285 Schönlaub HP (1969) Das Paläozoikum zwischen Bischofalm und Hohem Trieb (Zentrale Karnische Alpen). Jb Geol B-A 112(2):265–320 Schönlaub HP (1979) Das Paläozoikum in Österreich. Abh Geol B- A 33:1–24 Schönlaub H (1980) Carnic Alps. Field Trip A. In: Schönlaub H (ed) Second European conodont symposium, ECOS II, vol 35. Abh Geol B- A, pp 5–57 Schönlaub HP (1985) Das Paläozoikum der Karnischen Alpen. Arbeitstag Geol B- A 1985:34–52 Schönlaub HP (1988) The Ordovician-Silurian boundary in the Carnic Alps of Austria. In: Cocks LRM, Rickards RB (eds) A global analysis of the Ordovician–Silurian Boundary. British Museum (Natural History), London, vol 43, pp 107–115 Schönlaub H (1992) Stratigraphy, biogeography and paleoclimatology of the Alpine Paleozoic and its implications for plate movements. Jb Geol B-A 135(1):381–418 Schönlaub HP (1997) The Silurian of Austria. In: Schönlaub HP (ed) IGCP Project 421 North Gondwanan Mid-Palaeozoic Biodynamics: Inaugural Meeting Vienna, Sept 17–21, 1997: Guidebook, Ber Geol B-A, vol 40, pp 20–41 Schönlaub HP, Histon K (1999) The Palaeozoic of the Southern Alps. In: Histon K (ed) V international symposium, Cephalopods— present and past, Carnic Alps Excursion Guidebook, 1999. Ber Geol B-A 47:6–30 Schönlaub HP, Histon K (2000) The palaeozoic evolution of the Southern Alps. Mitt Österr Geol Ges 92:15–34 Schönlaub HP, Lammerhuber A (2009) Geologie und Zeit an Beispielen aus den Karnischen Alpen. Edition Lammerhuber & Geologische Bundesanstalt, pp 1–48, Wien Schönlaub HP, Sheehan PM (2003) Die Krise des Lebens am Ende des Ordoviziums. In: Hansch W (ed) Katastrophen in der

13

Edgeschichte–Wendezeiten des Lebens, Museo 19, pp 82–98, Heilbronn Schönlaub HP, Klein P, Magaritz M, Rantitsch G, Scharbert S (1991) Lower carboniferous paleokarst in the Carnic Alps (Austria, Italy). Facies 25(1):91–117 Schönlaub HP, Attrep M, Boeckelmann K, Dreesen R, Feist R, Fenninger A, Hahn G, Klein P, Korn D, Kratz R (1992) The Devonian/Carboniferous boundary in the Carnic Alps (Austria)—a multidisciplinary approach. Jb Geol B-A 135(1):57–98 Schönlaub HP, Ferretti A, Gaggero L, Hammarlund EU, Harper DAT, Histon K, Priewalder H, Spötl C, Storch P (2011) The late Ordovician glacial event in the Carnic Alps (Austria). In: Gutiérrez-Marco JC, Rábano I, Garcia-Belido D (eds) Ordovician of the World. Cuad Mus Geomin, IGME, vol 14, pp 515–526 Selli R (1963) Schema geologico delle Alpi Carniche e Giulie occidentali. Giorn Geol 30(2):1–121 Serpagli E (1967) I conodonti dell’Ordoviciano superiore (Ashgilliano) delle Alpi Carniche. Boll Soc Paleont It 6:30–111 Sheehan PM (1973) The relation of Late Ordovician glaciation to the Ordovician-Silurian changeover in North American brachiopod faunas. Lethaia 6(2):147–154. doi:10.1111/j.1502-3931.1973. tb01188.x Spalletta C (1983) Concordanza stratigrafica tra carbonati, radiolariti e Flysch ercinico nelle alpi camiche (Devoniano-Silesiano). Mem Soc Geol It 24:11–21 Spalletta C, Vai GB (1984) Upper Devonian intraclast parabreccias interpreted as seismites. Mar Geol 55(1–2):133–144 Spalletta C, Venturini C (1988) Conglomeratic Sequences in the Hochwipfel Formation: a new palaeogeographic hypothesis on the Hercynian Flysch Stage of the Carnic Alps. Jb Geol B-A 131(4):637–647 Spalletta C, Venturini C (1990) Stratigraphic correlation form of the Palaeozoic sequence in the Carnic Alps. Rend Soc Geol It 12:417–421 Spalletta C, Venturini C (1995) Late Devonian-early Carboniferous synsedimentary tectonic evolution of the Palaeocarnic domain (Southern Alps, Italy). Giorn Geol 56(2):211–222 Spalletta C, Vai G, Venturini C (1980) Il flysch ercinico nella geologia dei Monti Paularo e Dimon (Alpi Carniche). Mem Soc Geol It 20(1979):143–265 Spalletta C, Vai GB, Venturini C (1982) La Catena Paleocarnica. Guida alia geologia del Sudalpino centro-orientale Guide Geol Reg Soc Geol It:281–292 Storch P, Schonlaub HP (2012) Ordovician-Silurian boundary graptolites of the Southern Alps, Austria. Bull Geosci 87(3):755–766 Suttner TJ (2007) Conodont stratigraphy, facies-related distribution patterns and stable isotopes (carbon and oxygen) of the uppermost Silurian to lower Devonian Seewarte section (Carnic Alps, Carinthia, Austria). Abh Geol B A 59:1–111 Szulczewski M, Belka Z, Skompski S (1996) The drowning of a carbonate platform: an example from the Devonian-Carboniferous of the southwestern Holy Cross Mountains, Poland. Sediment Geol 106 (1–2):21–49. doi:10.1016/0037-0738(95)00145-X

13

Int J Earth Sci (Geol Rundsch) Taramelli T (1895) Osservazioni sul Paleozoico delle Alpi Carniche. Boll Soc Geol It 14:277–280 Tessensohn F (1974) Zur Fazies paläozoischer Kalke in den Karawanken (Karawankenkalke II). Verh Geol B-A:89–130 Vai GB (1963) Ricerche geologiche nel gruppo del M. Coglians e nella zona di Volaia (Alpi Carniche). Giorn Geol 30:137–198 Vai GB (1971) Ordovicien des Alpes carniques. Mém BRGM Paris 73:437–450 Vai GB (1976) Stratigrafia e paleogeografia ercinica delle Alpi. Mem Soc Geol It 13(1):7–37 Vai GB (1980) Sedimentary environment of Devonian pelagic limestones in the Southern Alps. Lethaia 13(1):79–91 van Amerom AWJ, Schönlaub HP (1992) Pflanzenfossilien aus dem Karbon von Nötsch und der Hochwipfel-Formation der Karnischen Alpen (Österreich). Jb Geol B-A 135(1):195–216 Venturini C (1990) Field workshop on Carboniferous to Permian sequence of the Pramollo-Nassfeld Basin (Carnic Alps) (September 2–8, 1990). Guidebook: Udine 159 Venturini C, Spalletta C (1998) Remarks on the Palaeozoic stratigraphy and the Hercynian tectonics of the Palaeocarnic chain (Southern Alps). In: Perri MC, Spalletta C (eds) Southern Alps field trip guidebook. Seventh international conodont symposium held in Europe. Giorn Geol, Sp Issue, pp 69–88 Venturini C, Pondrelli M, Fontana C, Delzotto S, Discenza K (2001– 2002) Carta geologica delle Alpi Carniche (Geological map of the Carnic Alps), scala 1:25.000 (western and eastern sheets, 990 sq. km) Venturini C, Spalletta C, Vai G, Pondrelli M, Salvador G, Carulli G (2009) Note Illustrative della Carta Geologica d ‘Italia alla scala 1:50.000, Foglio 031 Ampezzo. ISPRA (ex-APAT-Servizio Geologico d’Italia):7–232 Ver Straeten CA (2010) Lessons from the foreland basin: Northern appalachian basin perspectives on the acadian orogeny, vol 206 Vinassa de Regny P (1910) Rilevamento geologico della tavoletta Paluzza. Boll R Com Geol It 41:173–200 Vinassa de Regny P (1915) Ordoviciano e Neosilurico nei gruppi del Germula e di Lodin. Boll R Com Geol It 44:295–308 Vinassa de Regny P, Gortani M (1905) Osservazioni geologiche sui dintorni di Paularo (Alpi Carniche). Boll Soc Geol It 24:1–15 von Gaertner HR (1931) Geologie der Zentralkarnischen Alpen. Denkschr Österr Akad Wissensch, Math-Naturw KL 102:113–199 von Raumer JF, Bussy F, Schaltegger U, Schulz B, Stampfli GM (2013) Pre-Mesozoic Alpine basements—their place in the European Paleozoic framework. Geol Soc Am Bull 125(1– 2):89–108. doi:10.1130/b30654.1 Walliser OH (1964) Conodonten des Silurs. Abh Hess L-Amt Bodenforsch 41:106 Wenzel BC (1997) Isotopenstratigraphische Untersuchungen an silurischen Abfolgen und deren paläozeanographische Interpretation. Erlanger Geol Abh 129:1–117