Introduction

Glaciations have a major impact on the evolution of life and affect the entire planet Earth (Hoffman 2009). Neoproterozoic glaciations played a major role as a bottleneck, and upon melting a major driver, for the origin of first large multicellular life, and the rise of the Ediacaran biota (Narbonne and Gehling 2003; Vickers-Rich and Fedonkin 2007). Cryogenian (c. 850–635 Ma) strata record the most extreme climate episodes in Earth’s history. The widespread occurrence of low-latitude glacial deposits globally together with the unusual geochemistry and sedimentology of cap carbonates (Hoffman et al. 1998; Bao et al. 2008) inspired the hypothesis of Snowball Earth (Kirschvink 1992, Kirschvink et al. 2000). Timing of Cryogenian glaciations is well established, represented by the Marinoan and Sturtian Snowball Earth events. The Marinoan Snowball Earth (Hoffman et al. 1998) is bracketed between c. 635 and 639 Ma (Hoffmann et al. 2004; Prave et al. 2016). Its predecessor, the Sturtian snowball, began around 717 Ma and ended at c. 660 Ma (Macdonald et al. 2010; Rooney et al. 2015). Cryogenian global glacial events were followed by a mid-Ediacaran Gaskiers glaciation aged between c. 581 and 579 Ma (Pu et al. 2016). The global nature of the Gaskiers ice age is still under debate. It is noteworthy that elements of the Ediacaran biota are first recorded about 5 Ma after the Gaskiers glaciomarine deposition of the Drook Formation, exposed on the Avalon Peninsula (Newfoundland) and hosting the Avalon assemblage, the oldest known community of large multicellular organisms on Earth (Narbonne et al. 2012, 2014).

Here, we introduce a new post-Gaskiers glacial event recorded in upper Ediacaran-aged strata occurring in the Cadomian orogen of Gondwanan Europe (Fig. 1). We present evidence for glaciomarine sedimentation in late Neoproterozoic evidenced in strata from the Saxo-Thuringian zone of the Central European Variscides (Germany) and from the Central Iberian zone of the Variscides in Spain. These sedimentary units are characterized by glaciomarine sedimentary features. Their maximum age is provided by U–Pb dates of youngest detrital zircon populations and of magmatic zircon from granitoid pebbles within the diamictites. U–Pb ages and Hf isotopes of detrital zircon were used as provenance indicators. We integrate all new data with the previous records and proposed the existence of a new late Ediacaran glacial episode.

Fig. 1
figure 1

Location of Variscan massifs in Central and Western Europe. AM Armorican massif, BM Bohemian massif, FCM French Central massif, IM Iberian massif, M Moravo-Silesian unit, S Sudetes. Variscan zones: CIZ Central Iberian, CZ Cantabrian, GTOM Galicia-Tras os Montes, MZ Moldanubian, OMZ Ossa-Morena, PL Pulo do Lobo, SPZ South Portuguese, SXZ Saxo-Thuringian, TBU Teplá-Barrandian unit, WALZ West asturian–Leonese. Ediacaran glaciomarine deposits in the Cadomian orogen: a Weesenstein and Clanzschwitz, b Orellana

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Geological setting

The Cadomian orogeny culminating at the end of the Precambrian and the earliest Cambrian is the principal pre-Variscan orogenic event in both the Bohemian Massif and Iberia (Fig. 1). Cadomian orogenic processes are characterized by a wide range of geotectonic events that took place in a marginal setting at the periphery of NW and N Africa from c. 750 to 540 Ma (e.g., Linnemann et al. 2000, 2014; Gutiérrez-Alonso et al. 2003; Drost et al. 2011). The type area for the Cadomian orogeny is the Armorican massif in north-western France (e.g., Graindor 1957). Rocks related to the Cadomian orogeny are commonly referred to by the collective term “Cadomian basement”.

Glacial deposits in the Saxo-Thuringian zone (Bohemian massif)

Our study focuses on rocks from the passive margin of the Cadomian arc–back-arc system (Linnemann 2007; Linnemann et al. 2014), best preserved in the Elbe zone and the North Saxon antiform in the southeastern part of the Saxo-Thuringian zone (NE Bohemian massif) (Fig. 2). In particular, we investigated the Ediacaran-aged Weesenstein and Clanzschwitz groups (Linnemann 2007). Furthermore, we included in our research the granodiorites of Laas and Dohna (earliest Cambrian). These plutonic rocks intruded the Ediacaran sedimentary units at c. 540 Ma after their deformation (Linnemann et al. 2010a).

Fig. 2
figure 2

Geological map of the Elbe zone and adjoining geological units showing the position of the Weesenstein, Clanzschwitz, and Rödern groups (Ediacaran) (modified from Pietzsch 1927; Linnemann 2007). 1 diamictites, quartzites, greywackes, and other siliciclastic metasedimentary rocks of the Weesenstein, Clanzschwitz, and Rödern groups (Ediacaran), 2 greywackes and mudstones (Ediacaran, Lausitz Greywacke complex), 3 Cambro-Ordovician quartzites and quartz schists, 4 Palaeozoic rocks not subdivided (Ordovician to Early Carboniferous), 5 Phyllitic Unit of the Elbe zone (Mühlbach-Nossen group, Cambro-Ordovician, to Devonian protoliths), 6 Early Variscan molasses of Hainichen (upper Viséan), 7 Rotliegend (Early Permian Variscan molasses), 8 Late Cretaceous, 9 granodiorites and anatexites (c. 540 Ma, Lausitz Granitoid complex), 10 Granodiorites of Dohna and Laas (c. 540 Ma, Elbe zone), 12 Großenhain orthogneiss, 13 Meißen massif (c. 335 Ma, not subdivided, monzonites, diorites, granites), 14 Upper Carboniferous volcanic and sub-volcanic rocks, 15 granulites of the Granulitgebirge (“Saxonian Granulites”, Variscan metamorphic core complex), 16 schist belt on the periphery of the Granulitgebirge, 17 flasergabbro, 18 greenschists (prasinites), 19 orthogneiss (c. 490 Ma) of the Frankenberg Zwischengebirge (Variscan nappe complex), 20 gneisses and other lithologies of the Erzgebirge nappe pile (high- and ultra-high-pressure rocks, not subdivided), 21 phyllites and other lithologies of the Erzgebirge nappe pile (low pressure rocks, not subdivided), 22 Variscan granitoids (c. 335–320 Ma), 23 major fault

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The Weesenstein group is exposed in the Elbe zone, which forms an NW–SE trending Variscan shear zone exhibiting a strong dextral strike-slip movement of at least 80–120 km sub-horizontal displacement (Linnemann 1994). Strike-slip movements took place in a time interval from c. 335 to 327 Ma (Hofmann et al. 2009). The Elbe zone divides the Variscan nappe pile of the Erzgebirge in the SW from the Lausitz Block in the NE (Fig. 2). While the latter consists mainly of Cadomian basement rocks (Lausitz Greywacke complex, and Lausitz Granodiorite complex, Linnemann et al. 2010a), the Erzgebirge is composed of a complex Variscan nappe pile of at least four separate nappes characterized by low pressure, and high and ultra-high pressure rocks (e.g. Sebastian 2013 and references therein). After reaching crustal levels at c. 340 Ma, and exhumation a short time later, a top-to-NW directed final emplacement was responsible for the origin of the dextral strike-slip movements along the Elbe zone (Linnemann 1994; Linnemann et al. 2010b and references therein). The Elbe zone is fault-bounded against the Erzgebirge nappe pile by the Mid Saxon fault, whereas the northern limit is the northern branch of the West Lausitz fault (Fig. 2). Variscan basement rocks of the Elbe zone show a protolith age range from the Ordovician to the Early Carboniferous and crop out in the Elbtalschiefergebirge (SE of Dresden) as well as in the Nossen-Wilsdruff Schiefergebirge (Fig. 2). Cadomian basement rocks form a narrow strip on the northeastern flank of the Elbe zone. In the Elbtalschiefergebirge (Figs. 2, 3), outcrops of the Weesenstein group and the Dohna granodiorite occur. Here, the Weesenstein group (Figs. 2, 3) is fault-bounded against Palaeozoic rocks along its southwestern limit by the southern branch of the West Lausitz fault. The strip of Cadomian rock units continues towards the northwest and is represented north of the Meißen massif by the Rödern group and the Großenhain orthogneiss (Fig. 2). The Cadomian basement of the Elbe zone continues further to the northwest into the North Saxon antiform. In that area, the Cadomian basement units of the Clanzschwitz group and the Laas granodiorite are present (Fig. 2). The complex pluton of the Meissen massif intruded into the strike-slip regime of the Elbe zone at 334 ± 3 Ma (Hofmann et al. 2009), emplaced into the southern branch of the West Lausitz fault, a distinct zone of weakness at that time (Linnemann 1994). Variscan strike-slip movements had ceased by 327 ± 3 Ma, the timing of intrusion of the Markersbach granite (Hofmann et al. 2009) that crosscuts all faults and cleavage planes in country rocks and thus provides a perfect time marker for the cessation of the Variscan strike-slip movement and the end of Variscan deformation in the region (Hofmann et al. 2009). At the beginning of the Permian, subsidence and felsic volcanism began in the Elbe zone resulting in formation of the Döhlen basin during the Rotliegend (Fig. 2) (Hofmann et al. 2009; Zieger 2015). This is the region in which the Weesenstein group and Dohna granodiorite outcrop were mapped in detail by Linnemann (1990, 1991). In the c. 2500 m-thick Weesenstein group, the older Seidewitz formation (c. 1500 m) occurs in the SW, while the younger Müglitz formation (c. 1000 m) is located in the NE. The section is rootless and fault-bounded at the base of the Seidewitz formation by the southern branch of the Westlausitz fault (Fig. 4). The Müglitz formation is limited by the intrusive contact with the Dohna granodiorite (Figs. 3, 4). The latter is dated at c. 538 Ma (Pb–Pb age, Gehmlich 2003). The general metamorphic overprint is characterized by a mid-to-high greenschist grade (in part biotite grade). Bedding is sub-vertical with two pronounced cleavages (Cadomian, Variscan, Linnemann 1990). The older Seidewitz formation represents a pre-glacial deposit, while the younger Müglitz formation contains glaciomarine sedimentary rocks (Fig. 5).

Fig. 3
figure 3

Geological map of the Elbtalschiefergebirge at Dresden (modified from Linnemann et al. 2008). 1 Weesenstein group containing the older Seidewitz formation including the Purpurberg Quartzite member and the younger Müglitz formation (Ediacaran), 2 Purpurberg quartzite member and its equivalents (Seidewitz formation, Weesenstein group, Ediacaran), 3 Quartz schists often under- and overlying the Purpurberg quartzite member (Seidewitz formation, Weesenstein group, Ediacaran), 4 Isolated pebbles and conglomerates of glaciomarine origin (Müglitz formation, Weesenstein group, Ediacaran), 5 Meta-basalts (Weesenstein group, Ediacaran), 6 Dohna granodiorite (Early Cambrian), 7 Cadomian metasediments of the Erzgebirge block overprinted under upper greenschist to amphibolite facies conditions during Variscan orogenic processes, 8 Ordovician, Silurian and Devonian metasediments and volcanic rocks of the Elbtalschiefergebirge overprinted under greenschist facies conditions during Variscan orogenic processes (so-called “Phyllitic unit” of the Elbe zone or Mühlbach-Nossen group), 9 Ordovician granite (Tourmaline granite, 485 ± 6 Ma, Linnemann et al. 2000), 10 Early Carboniferous metasediments, 11 Devonian volcano-sedimentary complex of the Elbtalschiefergebirge, 12 Devonian “diabase-limestone series” of the Elbtalschiefergebirge, 13 Early Carboniferous siliciclastic flysch deposits of the Elbtalschiefergebirge, 14 Variscan (Early Carboniferous) granitoids of the Meissen massif (c. 335 Ma) and the Markersbach granite (c. 320 Ma, Hofmann et al. 2009), 15 Permian and younger sedimentary cover, 16 major faults and dextral shear zones of the Elbtalschiefergebirge originated during Variscan orogenic processes, 17 Limit of Permian (Rotliegend) sedimentary rocks, 18 Limit of Cretaceous (Cenomanian–Turonian) sedimentary rocks, P-Top of the Purpurberg Mtn., type locality for the Purpurberg quartzite member

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Fig. 4
figure 4

Lithostratigraphic columns of the Weesenstein group (Elbe zone, Elbtalschiefergebirge, Saxo-Thuringian zone) (from Linnemann 1990, 1991). 1 conglomerate with pebbles of quartz and hornfels (diameter of pebbles up to 3 cm); 2 conglomerate with pebbles of quartz and rarely hornfels (diameter of pebbles up to 1 cm); 3 quartzite of the Purpurberg quartzite member (PBQM) and related quartzite horizons of the Seidewitz formation; 4 quartz schist; 5 chert; 6 glaciomarine Weesenstein diamictite of the Müglitz formation (“Weesenstein tillite”); 7 quartz wacke (immature greenish quartzite); 8 greywacke; 9 mudstone; 10 andalusite quartzfels; 11 metasediments strongly sheared and overprinted by contact metamorphism, often partial anatectic melting initialized by contact metamorphism is indicated; 12 meta-basalts; 13 Dohna granodiorite (538 ± 2 Ma, this study); 14 Hornblende diorite of the Meißen massif (334 ± 3 Ma, Hofmann et al. 2009); 15 sample location; PBQM Purpurberg quartzite member

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Fig. 5
figure 5

Outcrops and lithologies of the Weesenstein and Clanzschwitz groups. a Purpurberg quartzite member (Weesenstein group, Seidewitz formation) sub-vertically dipping and forming a ridge (cliffs of the Hallstein, Bahra valley, Elbtalschiefergebirge). b Purpurberg quartzite member, proximal thick-bedded facies, bedding is sub-vertical and shows ripple marks and cross bedding (Purpurberg near Friedrichswalde, Elbtalschiefergebirge, c cross bedding, b bedding plane), c Purpurberg quartzite member, distal thin-bedded facies, bedding is steeply dipping (outcrop in the “Röhrsdorfer Park” close to Röhrsdorf, Elbtalschiefergebirge). d Polished hand specimen of the conglomerate from the top-most part of the Purpurberg Quartzite member, white pebbles consist of quartzite, and black pebbles are composed of tourmaline hornfels (Purpurberg near Friedrichswalde, Elbtalschiefergebirge), e Isolated and tectonically stretched granitoid pebble in a matrix of greywacke and mudstone (Weesenstein group, Müglitz formation, “Weesenstein diamictite”, klippen at the railway station Weesenstein, Müglitz valley, Elbtalschiefergebirge), f Isolated and tectonically stretched granitoid pebbles in a matrix of greywacke and mudstone (Weesenstein group, Müglitz formation, “Weesenstein diamictite”, klippen at the footpath from Weesenstein to Köttewitz, Müglitz valley, Elbtalschiefergebirge), g Conglomerate of the Clanzschwitz group (member 3, North Saxon antiform)

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The Seidewitz formation is composed primarily of shelf deposits, such as quartz schist, quartzite, quartz wacke, greywacke, and mudstone (Fig. 5a–d). A prominent part of the sedimentary pile is the Purpurberg quartzite member (Fig. A), a sedimentary sequence of medium-to-thick-bedded quartzites (“Purpurberg quartzite”, Figs. 4, 5a, b) with quartz schist intercalations in the middle and conglomerate at the top (Fig. 5d). Primary sedimentary features, such as cross bedding and ripple marks (Fig. 5b), indicate that the lower part of the section is distributed in the SW, while the younger top in a transition to the overlying Müglitz formation occurs towards the NE (Fig. 4). The Purpurberg quartzite member represents the marine part of a fan delta deposit, which pinches out along strike towards the SE and to the NW (Fig. 4) (Linnemann 2007). The depositional centre is located at the Purpurberg and the Hallstein regions (Fig. 5a) along both flanks of the valley of the Bahra River (Fig. 4) (Linnemann 1990). The Purpurberg member is thin-bedded in the more distal parts of the fan delta and perfectly visible in the outcrops of the “Röhrsdorfer Park” (Figs. 4, 5c). The Seidewitz formation is limited in completeness at the base by the tectonic contact of the southern branch of the West Lausitz fault (Fig. 4). In that part of the formation, the geological situation becomes further complicated by intrusion of a Variscan hornblende granodiorite (Fig. 4), which is part of the Meissen massif (c. 335 Ma; Hofmann et al. 2009). Due to the contact-metamorphic overprint, most of the older part of the Seidewitz formation has been altered into an andalusite hornfels (Pietzsch 1917; Linnemann 1990). In its northwesternmost part, the Seidewitz formation was affected by partial melting and magmatization due to a contact-metamorphic overprint by the emplacement of the Meissen massif and strong shearing during the dextral strike-slip movements along the Elbe zone (Fig. 2) (Linnemann 1990). Within the Seidewitz formation, a number of amphibolites and meta-basalts occur (Fig. 4). In addition, one very badly exposed chert layer is known only from loose blocks found close to “Röhrsdorfer Park” (Pietzsch 1917) (Fig. 4).

The transition of the Seidewitz formation into the overlying (younger) Müglitz formation is gradual. The transitional interval is best exposed on the eastern flank of the Müglitz valley (Fig. 4) in the area northeast of the Weesenstein castle. Quartzite and quartz schist grade into greywacke, mudstone, and quartz wacke of the Müglitz formation. An additional, characteristic feature of the Weesenstein formation is occurrence of diamictites and conglomerates (Figs. 4, 5e–g). The diamictites crop out mainly along the flanks of the Müglitz Valley (Fig. 4). A few exposures are known in the valley of the Lockwitz River (Fig. 4) (Linnemann 1990). Isolated diamictite pebbles of granitoids, quartzite, and greywacke “swim” in a matrix of greywacke and mudstone (Fig. 5e, f), varying in size between 5 mm and 30 cm. Some of the pebbles are rounded and deformed due to Late Carboniferous tectonic overprint (Fig. 5e).

In some units, the diamictite is less deformed, particularly those cropping out around the railway station of Weesenstein. Here, diamictites which escaped penetrative deformation in the pressure shadows of more competent rock units, such as thick quartz wacke layers, can reach a thickness of more than 50 m. Glaciomarine features are documented by dropstones (Fig. 6a), dreikanters (Fig. 6b), flatiron-shaped pebbles (“Bügeleisen-Geschiebe”) (Fig. 6c, d), and facetted pebbles (Fig. 6d, e). Typical rainout deposits, formed by melting icebergs, also occur (Fig. 6f), with randomly oriented and predominantly quartzite and greywacke pebbles (47.2%). The second-largest group represented is volcanic rocks, such as rhyolite, felsite porphyry (andesite), and keratophyre (22.7%). Only 15.0% are igneous in origin including granodiorite, aplitic granite, and granite. The remaining 15.1% of the pebbles are vein quartz and quartz derived from pegmatites (Schmidt 1960). The dominance of sedimentary rocks among the loose pebbles in the diamictite is an additional argument for its glaciomarine origin, because glaciers erode mainly the shelf deposits in the source area before became floating ice.

Fig. 6
figure 6

Images from the glaciomarine deposits of the Weesenstein group (Müglitz formation, klippen at the railway station Weesenstein, Müglitz valley, Elbtalschiefergebirge). a Granitoid dropstone in a rainout deposit, b dreikanter, c bullet shaped clast (flat iron-shaped pebble) in a rainout deposit (upper part sharp top formed by ice abrasion, lower part stump end in the pressure shadow during ice abrasion) (“Bügeleisen-Geschiebe”), d facetted and striated pebble (flat iron-shaped pebble = “Bügeleisen-Geschiebe”), e sharp facetted pebble from a rainout deposit, f rainout sediment showing isolated pebbles with different orientations “swimming” a fine-grained matrix

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In addition to the observations of glaciomarine features in the macro-scale, we carried out studies on zircon surfaces recovered from the diamictites of the Müglitz formation and compared them with such surfaces derived from a sub-recent Quarternary lodgement till from Germany (Isle of Rügen, coastline north of Saßnitz, Fig. 7). Features indicative of glacial transport for detrital grains include high angularity, fresh faces, conchoidal breaks, and arc-step-like fractures reported by Immonen (2013). SEM backscatter images of investigated zircon grains show clear effects of ice abrasion in the Quarternary lodgement till (Fig. 7a–f), similar to such surfaces on the zircons recovered from the glaciomarine rainout deposit of the Weesenstein group (Fig. 7g–q, sample WST 1, Müglitz formation). Zircons from both samples exhibit scratches, flatiron shapes with relatively sharp top and stump end, and rip-up structures generated in the pressure shadows during ice abrasion. Edges of rip-ups had been sub-rounded subsequently during sedimentation processes and demonstrate that they are not formed during heavy mineral separation by the jaw crusher. Some zircon grains exhibit shortening by compression characterized by a release band running vertical to compressive strain (Fig. 7f, p). Cathodoluminescence images of zircons from the Weesenstein diamictite reflect a wide range of different tectonomagmatic settings from which such zircon grains were derived (Fig. 8). Most zircons show normal magmatic zoning, and this points to a felsic provenance. Complex zircons with core and rim are common. Other grains seem to have been derived from mafic melts or rocks which underwent ultra-high-pressure metamorphism.

Fig. 7
figure 7

SEM backscatter images of zircon grains affected by ice abrasion from a Quaternary glacial diamictite (a–f Germany, Isle of Rügen, coastline north of Saßnitz) and from a glaciomarine rainout deposit of the Weesenstein group (g–q sample WST 1, Müglitz formation). “p” indicates the assumed maximum pressure during ice abrasion and “s” the pressure shadow. Arrows show the resulting movement of the ice during abrasion processes. Images show examples of typical glacial features like high angularity, fresh faces, conchoidal breaks, and arc-step like fractures. a Scratches in a zircon grain. b Enlargement of a, c flat iron-shaped zircon grain. d zircon grain showing a more conic and a more stump end. e Scratches in the zircon grain shown in d. f Zircon grain illustrating shortening by compression, note the releasing band running perpendicular to the strain. g Zircon showing scratches (left) and rip-up structures (right), h enlargement of g. i Zircon showing rip-ups on the right side, which are sub-rounded by subsequent transport (see enlargement in j). j Enlargement of i, scratched zircon surface and sub-rounded edges of rip-ups, which rule out an origin during heavy mineral separation (e.g., by jaw crusher). k Zircon with relatively sharp top and stump end in the pressure shadow produced during ice abrasion., l zircon with rip-up showing sub-rounded edges. Similar to k, this grain hosts a sharp top and stump end in the pressure shadow produced during ice abrasion. m Zircon with rip-ups in the pressure shadow during ice abrasion. n Enlargement of M, zircon with rip-ups generated in the pressure shadow produced during ice abrasion; edges had been sub-rounded subsequently. o Zircon with scratches generated by ice abrasion. p Zircon grain showing shortening by compression, note the releasing band running perpendicular to the strain. q Flatiron-shaped zircon grain

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Fig. 8
figure 8

Cathodoluminescence (CL) images and U–Pb LA–ICP–MS ages of zircon grains from the Purpurberg Quartzite member (Seidewitz formation, Weesenstein group) (sample Pur 2). Complex zircon grains are very common (spots b3–b4; b6–b7; c15–c16). Zircon grains which were affected by high pressure metamorphism are frequent among the Eburnean-aged zircons (a52, c60). Very rare are zircon grains which originated in mafic melts (b35)

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Within the Elbe zone and the North Saxon antiform, rock units exist, which are related to the Weesenstein group. Towards the northeast, the Cadomian basement of the Elbe zone continues into the Radeburg area, bound by the northern and southern branches of the Westlausitz fault (Fig. 2). Here, Cadomian rocks are preserved in a narrow strip of meta-greywackes (Rödern group). Bedding and cleavage run sub-vertical. Locally layers of diamictite occur at Brotberg Mountain near Rödern (Schmidt 1960). The metasediments of the Rödern group exhibit a metamorphic overprint generated under greenschist facies conditions (biotite grade) and were intruded at c. 540 Ma by a protolith of the Großenhain orthogneiss (Fig. 4) (Schmidt 1960; Gehmlich 2003). During the Variscan strike-slip movements along the Elbe zone, the Cadomian basement strip was intruded by the c. 335 Ma old Meißen massif magma (Hofmann et al. 2009) (Fig. 2). The Rödern group is a strongly deformed equivalent of the Weesenstein group, while the Großenhain orthogneiss is a highgrade overprinted continuation of the Dohna granodiorite.

In the North Saxon antiform, about 30 km to the NW, there is a third area of importance with a Cadomian basement unit (Fig. 2). The metasedimentary Clanzschwitz group is subdivided into three un-named members (Fig. 9). The entire group is about 850 m thick and shows sub-vertical bedding and cleavage. Member 1 is composed of mudstones and quartz schists, and was intruded by the Laas granodiorite in the Early Cambrian. The strong contact metamorphism led to the origin of up to ~1 cm andalusite crystals. The second member of the Clanzschwitz group is composed of quartzite and quartz schist. Member 3 hosts diamictite (Clanzschwitz diamictite) and conglomerate layers (Fig. 5g). The Clanzschwitz group was intruded by the Laas granodiorite at c. 533 Ma (Pb–Pb age, Gehmlich 2003), and underwent a high greenschist facies and biotite grade regional metamorphism. After a gap of sedimentation, the Cadomian basement in the North Saxon antiform was transgressed by Lower Ordovician sediments, preserved as quartzites and mudstones of the Collmberg formation (Pietzsch 1962), cropping out south of Clanzschwitz (Fig. 2).

Fig. 9
figure 9

Geological map with sample locations and a lithostratigraphic column of the Clanzschwitz group (North Saxon antiform). 1 Mudstones and quartz schists affected by contact metamorphism of the Laas granodiorite with large andalusite crystals (up to 1 cm) (member 1 of the Clanzschwitz group), 2 quartzite and quartz schist (member 2), 3 diamictite (glaciomarine diamictite, “Clanzschwitz tillite”) (member 3), 4 quartzite and quartz schist (Ordovician, Collmberg formation), 5 Laas granodiotite, 6 Laas granodiorite with a gneissic texture, 7 fault, 8 roads, and 9 villages

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The three described Cadomian basement exposures host the same geological unit which forms a ca. 60 km-long strip along the northeastern flank of the Elbe zone and at the southeastern margin of the North Saxon antiform (Fig. 2). The sedimentary and igneous parts of all three Cadomian basement units exhibit striking similarities. Both the older parts of the Clanzschwitz (members 1 and 2) and Weesenstein groups (Seidewitz formation, Purpurberg quartzite member) consist of quartzite and quartz schists. In the younger parts, both groups are composed of diamictites, greywackes, and conglomerates, also similar to the Müglitz formation (Weesenstein group) and member 3 of the Clanzschwitz group. In addition, the package of greywacke and mudstone in the Rödern group hosts diamictite horizons, poorly exposed at the Brotberg area near the village of Rödern (Schmidt 1960). The occurrence of facetted pebbles and the c. 60 km-wide distribution of diamictites in the Weesenstein, Clanzschwitz, and Rödern areas at the same lithostratigraphic level supports a glaciomarine nature of the diamictites (Linnemann 2007). All three sedimentary units of Cadomian basement in the Elbe zone were intruded by granodiorites of a similar age (c. 533–540 Ma, Pb–Pb ages by Gehmlich 2003). All in all, the Cadomian basement units of the Elbe zone share striking similarities. Thus, we assume that these units formed a single pre-Variscan rock unit along the southern margin of the Lausitz block. The formerly coherent unit of Cadomian passive margin sequences, including diamictite levels, was later sliced into the described three sub-units at Weesenstein, Rödern, and Clanzschwitz due to dextral strike-slip movements along the Elbe zone (Linnemann 1994; Linnemann et al. 2010a, b) in the c. 335–327 Ma time interval (Hofmann et al. 2009).

Glacial deposits in the Central Iberian zone (Iberian massif)

Lotze (1945) proposed a division of the Iberian massif into six tectonostratigraphic zones, which still forms the basis of present subdivisions. Lotze’s Eastern Lusitanian–Alcudian Zone is a 100–150 km-wide band that stretches from Coimbra–Castelo Branco–Cáceres to Ciudad Real and contains the glaciomarine diamictites dealt with herein. It was later merged with the Galician–Castillian Zone into the Central Iberian Zone by Julivert et al. (1972). Neoproterozoic rocks occur in most tectonostratigraphic domains of the Variscan orogen in Iberia interpreted to have once belonged to peri-Gondwanan Cadomia (Quesada 1990, 1991). The various divisions are thought to represent an arc/back-arc system (Ossa-Morena zone) accreted to a passive margin of West Africa on which a retro-arc basin developed (Central Iberian, West Asturian–Leonese, and Cantabrian zones) as a consequence of such arc accretion in late Neoproterozoic-earliest Cambrian time (Quesada 1990, 1991, 1996, 2006). Deformation associated with this arc-continent event characterizes the Cadomian orogeny in this part of peri-Gondwana and variably affected the various parts involved (Gutiérrez-Alonso et al. 2003). Among the different tectonostratigraphic zones in Iberia, the Central Iberian zone is by far the one of the largest and thickest exposures of Cadomian basement rocks (commonly known as Schist and Greywacke Complex in this area).

Ovtracht and Tamain (1970) studied the Proterozoic rocks of the Schist and Graywacke complex in the core of the Alcudia anticline. They subdivided the Ediacaran-aged sedimentary rocks into Lower and Upper Alcudian groups (Fig. 10) and subsequently integrated into the so-called Alcudian supergroup (San José et al. 1990) (Fig. 11). These rocks have also been called the Extremenian Dome group (Álvarez-Nava Oñate et al. 1988), a name which unfortunately has no relation to the investigated area. Therefore, the terms Lower and Upper Alcudian group are preferred here. In the area around Orellana, Ediacaran and lower Palaeozoic strata were mapped and studied in detail by Pieren (2000). The exposed Lower Alcudian group is about 4000 m thick, but its base is nowhere exposed (Fig. 11). The group is characterized by a monotonous succession of turbidites composed of greywacke and mudstones (Fig. 11). According to García-Hidalgo et al. (1993) and Pieren (2000), the Lower Alcudian group from bottom to top is composed of (1) shales and greywackes (“La Coronada Shales formation”), (2) greywacke, microconglomerate, and shale (“Saint Mary of Zújar chapel greywackes and conglomerates formation”), and (3) shale, diamictite (pebbly mudstones), and greywacke (“Orellana matrix supported conglomerates formation”) (Fig. 11). Some greywacke turbidite beds are up to 6 m thick and have been used as marker horizons during geological mapping of the region. Fuenlabrada et al. (2016) analyzed major and trace elements, REE, and Sm–Nd isotopes of greywacke samples of the second unit. Their geochemical results point to an active margin setting as the most reasonable source for the deposition of the Ediacaran–Early Cambrian sequences of the Central Iberian zone. Trace element diagrams of the Ediacaran greywackes indicate a clear affinity to a continental island arc which matches with the paleogeographic/paleotectonic interpretations mentioned above (Quesada 1991, 1996, 2006; Pieren 2000).

Fig. 10
figure 10

Geological map of a part of the Alcudia antiform at the Orellana reservoir (Orellana region, Central Iberian zone, Iberia) (from García-Hidalgo et al. 1993; Pieren 2000; Clariana García et al. 2017). Locality 1: waterline of the sweetwater reservoir (images see Fig. 12), N 39°00′8.4″; W5°31′18.9″. Locality 2: roadcut near Orellana (images see Fig. 12), N 38°52′52.8″; W5°29′5.6″. Locality 3: roadcut near Orellana, sample location of OR-1 (images see Fig. 12), N 39°00′4.6″; W5°34′11.9″. Red dots are towns. Inset map CIZ Central Iberian zone, CZ Cantabrian zone, GTOM Galicia-Tras os Montes zone, MZ Moldanubian zone, OMZ Ossa-Morena zone, PL Pulo do Lobo zone, SPZ South Portuguese zone

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Fig. 11
figure 11

Idealized section of the Lower and Upper Alcudian groups (Alcudian supergroup) and overlying Ordovician strata from the Alcudia antiform near Orellana (Central Iberian zone, Spain). Left stratigraphic log of the Lower Alcudian group) (from García-Hidalgo et al. 1993; Pieren 2000). For abbreviations of inset, see Fig. 10

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The Lower Alcudian group was folded and locally cleaved during the Cadomian orogenic activity. Generally, a vertical and sub-vertical L1 lineation is present, which was generated by a (in the area) non-schistose gentle folding in the late Ediacaran (“intra-Alcudian” or Cadomian deformation). In addition, a Variscan schistose F1 exists. Most cobbles were oriented parallel to a 70ºN dipping L1 lineation. Some of them were not oriented or almost normal to the L1 lineation.

The Lower Alcudian group is unconformably overlain by the Upper Alcudian group (Fig. 11; Pieren 2000). This lithostratigraphic unit begins with conglomerates and coarse sandstones followed upsection by fine-grained sandstones and mudstones. Further upsection, shale and dolomite layers occur (Fig. 11). The basal unit of the Upper Alcudian group is considered to be latest Ediacaran. However, based on contained ichnofossils (Pieren 2000) and detrital zircons studied elsewhere in the region (Talavera et al. 2015), the upper part could be earliest Cambrian in age.

A second deformation affected both the Lower and Upper Alcudian groups. The Upper Alcudian group may well represent a Cadomian remnant or successor basin deformed during the latest stage of the Cadomian orogeny. After final Cadomian deformation, a time of uplift, intense weathering, and erosion may represent the local response to the generalized rift event recorded everywhere in Iberia during the Cambrian–Lower Ordovician (Quesada 1991; Sánchez-García et al. 2003, 2008, 2010). Subsequent subsidence, interpreted in response to opening of the Rheic Ocean as a culmination of rifting, allowed the area to be transgressed in the Lower Ordovician establishing a new passive margin of Gondwana that survived until Variscan collision in the Devonian (Quesada 1991, 2006). The Lower Ordovician-aged reddish sandstone and mudstone and the Armorican quartzite formation were the first to be deposited during this passive margin stage (Fig. 11; Pieren 2000).

Diamictites in the uppermost part the Lower Alcudian group (Figs. 11, 12) show distinct features indicative of a glaciomarine origin, including dropstones (Fig. 12b), rainout sediments (Fig. 12c), flat iron-shaped pebbles (“Bügeleisen-Geschiebe”) (Fig. 12d), and facetted pebbles (Fig. 12e). Pebble size varies between a few millimetres and 20 cm. Glaciomarine diamictite occur in c. 10–50 m-thick beds in the region of Orellana la Vieja, Orellana de la Sierra, and the Orellana reservoir (Embalse de Orellana; Pieren 2000).

Fig. 12
figure 12

Images from the glaciomarine deposits of the Lower Alcudian group (“Orellana diamictite”, Alcudia antiform, Central Iberian zone, Spain). a Diamictite containing pebbles of sedimentary rocks (quartzite, greywacke) (Locality 2, Fig. 10), b dropstone in a rainout deposit, c rainout deposit with millimetre-sized rock fragments, d faceted and flat iron-shaped pebble in a rainout deposit (flat faceted top and sharp left end formed by ice abrasion; stump right end formed in the pressure shadow during ice abrasion (“Bügeleisen-Geschiebe”) (Locality 3, Fig. 10), e Facetted pebble from a rainout deposit (flat iron-shaped pebble = “Bügeleisen-Geschiebe”) (Locality 3, Fig. 10), and f large flat iron-shaped pebble (=“Bügeleisen-Geschiebe”), a fine-grained diamictite matrix. (b, c, f images photographed in a roadcut at Casa el Canchal, N 38°52′52.8″; W5°29′5.6″)

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Samples for U–Pb dating and Hf isotope analyses

For our investigation to determine absolute age and provenance, we collected ten samples of clastic sedimentary rocks, six granodiorites, four granitoid pebbles, and one quartzite pebble. Pebbles were sourced from the glaciomarine diamictite layers in the Weesenstein and Clanzschwitz groups. One sample was taken from the Lower Alcudian group (Iberia) near Orellana. U–Pb LA–SF–ICP–MS* dating was carried out on zircons from all samples (*Laser ablation and inductively coupled plasma combined with a sector field mass spectrometer). Analysis of Hf isotopes by LA–MC–ICP–MS** was undertaken on zircon grains from Weesenstein and Clanzschwitz only (six sedimentary rock samples, three granitoid pebbles, one quartzite sample, and six granodiorite samples) (**Laser ablation and inductively coupled plasma combined with a multicollector mass spectrometer). The sample locations are marked in the columns and map of Figs. 4, 9 and 11. Exact location details are given in the headings of Table 1 (U–Pb ages) and Table 2 (Hf isotopes) (supplementary data).

The following sedimentary samples have been collected from the Müglitz formation of the Weesenstein group: WST1 (diamictite matrix, greywacke, Müglitz valley near Weesenstein, coordinates: N 50°56′09.77″; E 13°51′47.42″), GWQ1 (quartz wacke, Müglitz valley opposite to the castle in Weesenstein, coordinates: N 50°56′01.90″; E 13°51′34.80″), and W63d (conglomerate, Müglitz valley at the Fuchshübel, coordinates: N 50°56′23.21″; E 13°51′36.31″). From the Seidewitz formation (Purpurberg quartzite member) of the Weesenstein group, the quartzite sample P-Pur-1n-2-LA10 was collected (quartzite, top of the Purpurberg, coordinates: N 50°54′48.64″; E 13°54′23.03″). In addition, one quartzite pebble (WeG4) from the Müglitz formation was sampled (quartzite pebble, Müglitz valley at railway station Weesenstein, N 50°56′12.55″; E 13°51′49.67″). Furthermore, three granitoid pebbles from the diamictite in the Müglitz formation (Weesenstein group) were sampled: WeG2n, WeG3n, and WeG5 (granitoid pebbles, Müglitz valley at railway station Weesenstein, N 50°56′12.55″; E 13°51′49.67″). From the Dohna granodiorite, which is intrusive into the Weesenstein group, five samples were collected: Dohna1 (Müglitz valley, Köttewitzer Wehr, coordinates: N 50°56′29.12″; E 13°51′38.83″), Dohna2 (Müglitz valley, Köttewitzer Wehr, coordinates: N 50°56′29.28″; E 13°51′39.41″), Dohna 3 (Köttewitz, Müglitz valley on the foot path to Weesenstein, coordinates: N 50°56′33.25″; E 13°52′03.36″), Dohna 4 (roadcut near Dresden-Nickern and Goppeln, SE of the Gamighübel, temporary outcrop during road construction, coordinates: N 50°59′44.67″; E 13°46′28.07″), and DG1n (Bahra valley, coordinates: N 50°55′2.61″; E 13°54′41.60″.

From the Clanzschwitz group, two metasedimentary rock samples were sourced. The first sample (CL1) was taken from the matrix of the diamictite in member 3 (greywacke, Kleiner Steinberg near Clanzschwitz, coordinates: N 51°20′47.9″; E 13°09′41.5″. A second sample (CL2) came from the quartzite of member 2 (Großer Steinberg near Clanzschwitz, coordinates, N 51°21′1.17″; E 13°09′39.58″. A granitoid pebble (Cla 1n) from the diamictite (member 3) was sampled at the Schlangenberg near Wellerswalde (coordinates: N 51°20′44.01″; E 13°09′42.75″). From the Laas granodiorite that intruded the Clanzschwitz group, one sample was taken (Laas 1n, forest 1.5 km north of Clanzschwitz, N 51°21′34.34″; E 13°09′22.60″).

From the Lower Alcudian group, one sample of a diamictite matrix was taken from a road cut in the surroundings of Orellana (sample OR1, N 39°00′4.6″; W5°34′11.9″).

Hf isotope analysis by LA-MC- ICP-MS carried out on zircon grains from of five sedimentary rock samples (WST1, W63d, P-Pur-1b-2-LA10, CL1, and CL2), one quartzite pebble (WeG4), three granitoid pebbles (WeG2n, WeG5, Cla1n), and six granodiorite samples (Dohna 2, 3, 4, DG1n, Laas1).

Methods

Zircon concentrates were separated from 2 to 4 kg samples at the Senckenberg Naturhistorische Sammlungen Dresden (Museum für Mineralogie und Geologie) by crushing (jaw crusher), sieving, heavy mineral separation by heavy liquid (LST), and by making use of a magnetic separator. Final selection of the zircon grains for U–Pb dating was achieved by hand-picking under a binocular microscope. Zircon grains of all grain sizes and morphological types were selected, mounted in resin blocks, and polished to half their thickness. Concerning stratigraphic ages, the stratigraphic time scale of Ogg et al. (2016) has been used.

U–Th–Pb isotopes

Zircons (n = 1124) were analyzed for U, Th, and Pb isotopes by LA–SF ICP–MS techniques at the Museum für Mineralogie und Geologie (GeoPlasma Lab, Senckenberg Naturhistorische Sammlungen Dresden), using a Thermo-Scientific Element 2 XR sector field ICP-MS coupled to a New Wave UP-193 Excimer Laser System. A teardrop-shaped, low-volume laser cell was used to enable sequential sampling of heterogeneous grains (e.g., growth zones) during time resolved data acquisition. Each analysis consisted of approximately 15 s background acquisition followed by 30 s data acquisition, using a laser spot size of 25 and 35 µm, respectively. A common Pb correction based on the interference- and background-corrected 204Pb signal and a model Pb composition (Stacey and Kramers, 1975) was carried out if necessary. The necessity of the correction was judged on whether the corrected 207Pb/206Pb lies outside of the internal error of the measured ratios. Discordant analyses were interpreted with care. Raw data were corrected for background signal, common Pb, laser induced elemental fractionation, instrumental mass discrimination, and time-dependant elemental fractionation of Pb/Th and Pb/U using an Excel® spreadsheet program developed by Axel Gerdes (Institute of Geosciences, Johann Wolfgang Goethe-University Frankfurt, Frankfurt am Main, Germany). Reported uncertainties were propagated by quadratic addition of the external reproducibility obtained from the standard zircon GJ-1 (~0.6 and 0.5–1% for the 207Pb/206Pb and 206Pb/238U, respectively) during individual analytical sessions and the within-run precision of each analysis. According to the recommendation of Horstwood et al. (2016), a secondary zircon standard (Plesovice zircon) was analyzed. Sequences started with the analysis of five GJ1, one Plesovice, and ten unknowns followed by a repetition of a succession of three measurements of the GJ1 standard, one measurement of the Plesovice standard, and ten unknowns. U–Pb ages were in the recommended range of Sláma et al. (2008).

Concordia diagrams (2σ error ellipses) and Concordia ages (95% confidence level) were produced using Isoplot/Ex 2.49 (Ludwig 2001) and frequency and relative probability plots using AgeDisplay (Sircombe 2004). The 207Pb/206Pb age was taken for interpretation for all zircons >1.0 Ga, and the 206Pb/238U ages for younger grains. For further details on analytical protocol and data processing, see Gerdes and Zeh (2006). Th/U ratios are obtained from the LA–ICP–MS measurements of investigated zircon grains. U and Pb content and Th/U ratio were calculated relative to the GJ-1 zircon standard, and are accurate to approximately 10%. Analytical results of U–Th–Pb isotopes and calculated U–Pb ages are given in Table 1 (supplementary data).

Lu–Hf isotopes

Hafnium isotope measurements on zircons (n = 446) were made using a Thermo-Finnigan Neptune multicollector ICP-MS at Goethe Universität Frankfurt (GUF, Institut für Geowissenschaften, Mineralogie) coupled to RESOlution M50 193 nm ArF Excimer laser system (Australian Scientific Instruments) following the method described in Gerdes and Zeh (2006, 2009). Spots of 40 µm in diameter were drilled with a repetition rate of 5.5 Hz and an energy density of 6 J/cm2 during 50 s of data acquisition. The instrumental mass bias for Hf isotopes was corrected using an exponential law and a 179Hf/177Hf value of 0.7325. In case of Yb isotopes, the mass bias was corrected using the Hf mass bias of the individual integration step multiplied by a daily βHf/βYb offset factor (Gerdes and Zeh 2009). All data were adjusted relative to the JMC475 of 176Hf/177Hf ratio = 0.282160 and quoted uncertainties are quadratic additions of the within-run precision of each analysis and the reproducibility of the JMC475 (2SD = 0.0028%, n = 8). Accuracy and external reproducibility of the method were verified by repeated analyses of reference zircon GJ-1 and Plesovice. Results are in perfect agreement with previously published results (e.g., Gerdes and Zeh 2006; Sláma et al. 2008) and with the LA-MC-ICP MS long-term average of GJ-1 (0.282010 ± 0.000025; n > 800) and Plesovice (0.282483 ± 0.000025, n > 300) reference zircon at GUF.

The initial 176Hf/177Hf values are expressed as εHf(t), which is calculated using a decay constant value of 1.867 × 10−11 year−1, CHUR after Bouvier et al. (2008, 176Hf/177HfCHUR,today = 0.282785 and 176Lu/177Hf CHUR,today = 0.0336) and U–Pb and Pb–Pb ages obtained for the respective domains (Table 2, supplementary data). For the calculation of Hf two stage model ages (TDM) in billion years, the measured 176Lu/177Lu of each spot (first stage = age of zircon), a value of 0.0113 for the average continental crust, and a juvenile crust 176Lu/177LuNC = 0.0384 and 176Hf/177Hf NC = 0.283165 (average MORB; Chauvel et al. 2008) were used.

Cathodoluminescence (CL, supplementary data) and backscatter images were made using an EVO 50 scanning electron microscope (Zeiss) at Senckenberg Naturhistorische Sammlungen Dresden.

Results and discussion

Age constraints for glaciomarine sedimentation

U–Pb ages of detrital zircons are represented in Figs. 13 and 14. Because ash beds and fossils have not been found in the investigated sections, the maximum depositional age (MDA) had to be used for the stratigraphic age determination (Condon and Bowring 2011).

Fig. 13
figure 13

Combined binned frequency and probability density distribution plots of U–Pb LA–ICP–MS ages of detrital zircon grains from the Weesenstein group (Ediacaran of the Elbe zone) in the ranges of 400–1200 Ma (left) and 400 to 3800 Ma (right): W63d (conglomerate, Müglitz formation), WST1 (diamictite matrix, greywacke, Müglitz formation), GWQ1 (quartz wacke, Müglitz formation), and P-Pur1b-2-LA10 (quartzite, Purpurberg quartzite member, Seidewitz formation); MDA maximum depositional age

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Fig. 14
figure 14

Combined binned frequency and probability density distribution plots of U–Pb LA–ICP–MS ages of detrital zircon grains from the Weesenstein, Clanzschwitz, and Lower Alcudian groups (Ediacaran) in the ranges of 400–1200 Ma (left) and 400–3800 Ma (right): WeG4 (quartzite pebble from the Weesenstein diamictite, Ediacaran, Müglitz formation, Weesenstein group, Elbe zone), CL1 (diamictite matrix, Ediacaran, Clanzschwitz group, member 3, North Saxon antiform), CL2 (quartzite, Ediacaran, Clanzschwitz group, member 2, North Saxon antiform), and OR1 (greywacke, diamictite matrix, Ediacaran, Orellana diamictite, Lower Alcudian group, Central Iberian zone, Spain)

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The maximum depositional age for each sample was calculated from the youngest zircon population for each sandstone sample (Figs. 13, 14). For the Weesenstein diamictite (WST1, Müglitz formation, Weesenstein group), a maximum depositional age of 562 ± 5 Ma was obtained (Fig. 13). Underlying strata show similar maximum depositional ages of 562 ± 5 Ma for the Purpurberg quartzite member and 564 ± 5 Ma for quartz wacke sample GWQ1. Three granitoid pebbles from the diamictite horizons in the Müglitz formation exhibit ages of 576 ± 7, 572 ± 2, and 566 ± 4 Ma (Fig. 15), and support the maximum depositional age obtained from detrital zircon. The upper limit (minimum age) for the deposition of the Weesenstein diamictite is estimated based on the age of intrusion of the Dohna granodiorite (538 ± 2 Ma, Fig. 16).

Fig. 15
figure 15

U–Pb LA–ICP–MS zircon ages of three granitoid pebbles from the Müglitz formation of the Weesenstein group (Weesenstein diamictite) and from one granitoid pebble from member 3 of the Clanzschwitz group (Clanzschwitz diamictite)

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Fig. 16
figure 16

U-Pb ages from the granodiorites of Laas (top) and Dohna (bottom). For the Dohna granodiorite, U-Pb LA–ICP–MS zircon ages of five different samples were used (DG1n, Dohna 1–4). In the lower right diagram, a Concordia age of 538 ± 2 Ma is shown, which was calculated from the youngest zircon population of all samples. In addition, the diagram shows the large amount of inherited zircon grains occurring in the Dohna granodiorite

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The Clanzschwitz diamictite (CL1, member 3, Clanzschwitz group) exhibits a maximum depositional age of 570 ± 8 Ma supported by another age determination based on a zircon from the underlying member 2 (CL2), which is 565 ± 5 Ma (Fig. 14). The age of a zircon in the granitoid pebble taken from the diamictite (member 3, Cla1n) is 574 ± 5 Ma (Table 1, supplementary data). The intrusion age of the Laas granodiorite (537 ± 5, Fig. 16) offers the minimum age for the deposition of the Clanzschwitz diamictite.

The zircons from the Orellana diamictite (OR1) in the Lower Alcudian group of the Central Iberian zone yielded a maximum depositional age of 565 ± 4 Ma (Fig. 14).

In all three cases (Weesenstein, Clanzschwitz, and Orellana), the maximum depositional age for the diamictites is significantly younger than that for the Gaskiers glaciation in Newfoundland and in SE New England (Squantum). Thus, the three diamictites should represent a post-Gaskiers glacial event during Ediacaran time. Alternatively, a shifting of the icecenter from Avalonia (Gaskiers glaciation, c. 579–581 Ma) to West African peri-Gondwana (Weesenstein-Orellana glaciation, c. 565 Ma) over time would be possible. However, this solution seems to be more unlikely because of the existence of contemporaneous occurence of c. 565 Ma glacial deposits in Virginia (Hebert et al. 2010).

Provenance and crustal growth of the hinterland

U–Pb ages and Hf isotopes are classical tools for provenance studies and analysis of tectonomagmatic events, important for our understanding of patterns of crustal growth in the hinterland. Information of the pre-glacial hinterland can be obtained from the detrital zircon of quartzite pebble WeG4 taken from the Weesenstein diamictite, with a maximum age of deposition being 612 ± 7 Ma (Fig. 14). Because younger zircons aged at c. 565–600 Ma are lacking, it could not have been derived from the Purpurberg quartzite member, which underlies the diamictites of the Müglitz formation. Instead, the pebble must be derived from much older strata and, therefore, representative of the sedimentary cover of the paleolandscape in the hinterland at c. 600–610 Ma. The detrital zircon populations derived from the pebble exhibit a dominant peak at 630 Ma and a number of secondary peaks of Neoproterozoic age at c. 710 and 790 Ma. No Tonian and Mesoproterozoic zircons were found. Palaeoproterozoic zircon dates range between c. 1769 ± 66 and 2471 ± 13 Ma, while the Archaean zircon population occupies the range of 2525 ± 59 to 3163 ± 26 Ma (Fig. 14). Such a pattern of the zircon population is typical for peri-Gondwanan deposits originating in the periphery of the West African craton (Linnemann et al. 2014). A pronounced peak ranging from 1942 ± 18 to 2168 ± Ma is coeval with the so-called Eburnean orogeny. Neoproterozoic zircons came from one of the Pan-African orogens. The latter ones formed during continent–continent collision of the amalgamation of the Gondwana supercontinent at c. 610–630 Ma (Black et al. 1994). All in all, the zircon population found in the quartzite pebble from the Weesenstein tillite (Müglitz formation) represents the composition of rocks in a pre-glacial hinterland, which consisted of eroded parts of the Pan-African orogen and the West African craton. In this region, no Mesoproterozoic zircon grains were found.

All samples of clastic sediments from the Weesenstein, Clanzschwitz, and Orellana areas show a very similar pattern of zircon populations (Figs. 13, 14) typical for Cadomia and its West African hinterland (Linnemann et al. 2014). The U–Pb ages of the detrital zircons from this study can be grouped in three major groups: (1) the Neoproterozoic zircons (c. 565–950 Ma), the Palaeoproterozoic zircons (c. 1800–2200 Ma), and the Archaean zircons (c. 2500–3600 Ma). In addition, some samples of siliciclastic sedimentary rocks yielded a few Mesoproterozoic zircon grains (c. 1000–1200 and c. 1550 Ma). The Palaeoproterozoic and Archaean pattern is typical of a West African Craton provenance (e.g., Linnemann et al. 2000, 2014; Abati et al. 2010, 2012). Neoproterozoic zircon populations are diagnostic of provenance from the Cadomian orogen and from the Pan-African basement of the Trans-Saharan Belt or a related area (Linnemann 2007, Linnemann et al. 2011; Drost et al. 2011). The scarce Mesoproterozoic zircons might well have been derived from the cratonic neighbours of the West African Craton, which were at c. 570 Ma the Amazonian or Sub-Saharan cratons (e.g., Gärtner et al. 2013).

The pre-glacial deposits of the Purpurberg quartzite member (Seidewitz formation, Weesenstein group) and of member 2 of the Clanzschwitz group show a much stronger sedimentary input from the cratonic hinterland indicated by the strong Palaeoproterozoic and Archaean zircon populations (Figs. 13, 14).

Detrital zircon populations in the range of c. 565–580 Ma in our samples of sedimentary rocks and granitoid pebbles aged at c. 566–576 Ma indicate that the main peak of the Cadomian orogeny was over and the Cadomian orogen was exposed during the time of glacial processes.

Similarities of the zircon populations of glacigenic and related sedimentary rocks from Bohemia and Iberia document no differences in the composition of the West African hinterland and a coeval plate tectonics evolution. The relative large distance of about 1800 km between the glacial diamictites in Bohemia and Iberia constrains a relative wide distribution of the glaciated area within the Cadomian orogen.

Neoproterozoic zircons of glacial deposits and related rocks from the Saxo-Thuringian zone show a wide variability of the epsilon Hf (εHf) values, in the range of 13 and −41 (Fig. 17). About 15% have positive (εHf). Such an initial epsilon Hf notation indicates that these zircons crystallized either from magmas directly derived from a depleted mantle source (=juvenile magmas) contaminated with different amounts of older crust or by partial melting of a relatively juvenile crust of a Cadomian or Pan-African magmatic arc affinity, respectively (Fig. 17). Zircons close to the depleted mantle array are very scarce and about 660–690 Ma old (Fig. 17), which is in line with the early magmatic arc setting in the Pan-African orogens. Most zircon with ages of c. 540–700 Ma has lower εHf (up to −41), suggesting that juvenile magmas were contaminated by older crust or recycling of the juvenile island arcs together with some older crust (Fig. 17). About 85% of Neoproterozoic zircon show negative εHf. Zircons aged between ca. 540 and 700 Ma with negative εHf, in the range of −7 to −21, and model ages between c. 1.7 and 2.2 Ga can be explained by almost pure reworking of older Eburnean crust (Fig. 17). Surprisingly, eleven Neoproterozoic zircon grains show more negative εHf between −20 and −41. Such zircons lie on a linear array with most Eburnean and Archaean zircons, and may have been formed by recycling of c. 2500–3800 Ma old Archaean crust (Fig. 17). The relatively abundant group with negative and moderate positive εHf (+8 to −8) is only compatible with mixing of the newly formed Cadomian crust at c. 540–700 Ma and of Eburnean and Archaean crust (see mixing field in Fig. 17). Zircons within the c. 537–540 Ma age group of old granodiorites of Laas and Dohna show a spread of the εHfi in the range of +2 to −17 and fall in the same mixing field. Thus, Cadomian granitoid magmatism seems to be derived from an evolved Cadomian crust (Fig. 17).

Fig. 17
figure 17

Hf isotope evolution diagrams summarizing the data of zircons from the studied metasedimentary rocks of the Weesenstein and Clanzschwitz groups. The probability plot of the zircon age populations is represented in grey. See text for discussion. Top Hf isotope evolution diagram of analyzed zircon grains in the age range of ca. 500–4000 Ma. Below Hf isotope evolution diagram of analyzed zircon grains in the age range of 500–1200 Ma. For details and references of depleted mantle evolution, see Gerdes and Zeh (2006) and Dhuime et al. (2011). Data were analyzed using the decay constant of 1.867 × 10−11 (Scherer et al. 2001) and the CHUR parameters of Bouvier et al. (2008)

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From Mesoproterozoic zircon grains, only four analyses could be performed. Three of them point to the recycling of a Palaeoproterozoic crust, and one could be derived from the recycling of a true Mesoproterozoic crust (Fig. 17). Palaeoproterozoic zircons (c. 1.7–2.4 Ga), in most cases, exhibit εHf of +8 to −10 and model ages of about 2.0 and 2.7 Ga (Fig. 17). Such grains document a recycling of Palaeoproterozoic and Neoarchean crust. A number of zircon grains have positive εHf notations. Some of them fall on the depleted mantle curve. These zircons show a strong affinity to a juvenile Eburnean crust. Other Palaeoproterozoic zircon grains are characterized by εHf of −10 to −26 (Fig. 17). These zircons demonstrate a recycling of a ca. 3.0–3.8 Ga old Archean crust. Two Palaeoproterozoic zircons from the Purpurberg quartzite member are exceptional. These two zircons are 2056 ± 34 and 2103 ± 40 Ma old, and show a strongly negative εHf of −34 and −32 (Fig. 17). Calculated TDM model ages are 4.27 and 4.21 Ga, respectively. Both grains lie on a linear array parallel to the 176Lu/177Hf evolution of a continental crust extracted from the depleted mantle during the Hadean that mirrors a Jack Hills Trend (Fig. 17), named after the location of the oldest zircon on Earth (Wilde et al. 2001). U–Pb ages combined with Hf isotope data demonstrate that a c. 4.2–4.3 Ga old crust must have survived in the framework of the West African Craton at least until c. 2.05–2.1 Ga ago.

Most Neo- to Palaeoarchaean zircon grains dated at 2.5 to 3.5 Ga exhibit εHf values between +4 and −18 with expecting model ages of 2.7–3.7 Ga (Fig. 17). Neo- to Palaeoarchaean crustal evolution points to the recycling of late Neo- to Eoarchean crust of the West African Craton. TDM model ages indicate three different Archean crustal domains, which would have been recycled. These units are characterized by model ages of c. 3.0, 3.2–3.3, and 3.5–3.7 Ga, respectively. Most known West African Archean ages of granitoid magmatism fall between c. 2.7 and 3.0 Ga (e.g. Rocci et al. 1991). Scarce occurrence of 3.42–3.52 Ga old felsic rocks in the Amsaga area (Mauritania) has been reported by Potrel et al. (1996, 1998). It is very obvious, that both, Liberian (c.3.0 Ga) and Leonian (c. 3.4 Ga) crust, were recycled during Archaean time.

Conclusions

Sedimentary features, aspects of basin development, plate tectonics framework, U–Pb ages, and Hf isotope data together allow reconstruction of Ediacaran glacial processes, palaeogeography, timing of orogenic processes, and crustal growth of the piece of the Cadomian orogen preserved in the Elbe zone and the North Saxon antiform of the Saxo-Thuringian zone (NE Bohemian Massif) and from the Alcudian group of the Central Iberian zone in Iberia. Sedimentary and other features of the deposits further allow tracing a glacial event within sedimentary rocks related to the Cadomian orogen (Weesenstein and Clanzschwitz groups, and Lower Alcudian group). Indicative for glaciomarine deposition are dropstones, flatiron-shaped pebbles (“Bügeleisen-Geschiebe”), facetted pebbles, dreikanters, rainout sediments, and zircon grains affected by ice abrasion. Pre-glacial and glaciomarine sedimentation took place in a time span between c. 565 and 540 Ma. Age constraints completely rule out a correlation of this glacial event recorded in the Cadomian metasedimentary rocks of the Saxo-Thuringian and Central Iberian zones with the Gaskiers glaciation, which occurred around 579–581 Ma (Bowring et al. 2002; Thompson et al. 2014; Pu et al. 2016). Instead, a correlation with a glacial event recorded in Northwest Africa (Morocco) (Vernhet et al. 2012) and in Saudi Arabia (Vickers-Rich and Fedonkin 2007) is much more realistic and provides more information about the timing of this young Neoproterozoic glacial event. At Bou Azzer (Moroccan Anti-Atlas), Vernhet et al. (2012) describe glaciomarine structures (striated pavement on rhyolites) within the Ouarzazate group. The age of the latter is constrained by published U–Pb ages of the Ouarzazate group and can be placed in a time window between c. 570 and 562 Ma (Karaoui et al. 2015). In Saudi Arabia, another glaciomarine diamictite has been detected in Ediacaran sedimentary deposits of the Dhaiqa formation. An ash bed just below the diamictite provided a U–Pb age of 560 ± 4 Ma (Vickers-Rich and Fedonkin 2007). Furthermore, an Ediacaran glacial deposit (Moelv diamictite) crops out in Baltica (Bingen et al. 2005). However, available age data are not sufficient to constrain an age younger than the Gaskiers glaciation (Bingen et al. 2005). All in all, Ediacaran diamictites of the Saxo-Thuringian zone are bracketed in a time range of 565–540 Ma. We correlate glacial deposits of Bohemia and Iberia with the closer occurrences of glaciomarine diamictites in the Anti-Atlas (Bou Azzer) and in Saudi Arabia (Dhaiqa formation), which suggest a sedimentation age closer to 565 Ma for the Weesenstein, Clanzschwitz, and Orellana diamictites. Because of the clear post-Gaskiers age of all these glacial deposits, the data allow to define a new Ediacaran glacial event younger than the Gaskiers glaciation. Because of the main occurrences, we propose to name the glacial event at c. 565 as Weesenstein–Orellana glaciation (Fig. 18).

Fig. 18
figure 18

Timing of the Weesenstein–Orellana glaciation and its relation to the δ13C curve (Halverson et al. 2005) (left); palinspastic map showing palaeopositions of Ediacaran post-Gaskiers glacial deposits in peri-Gondwana, North Africa, and Arabia (right). Cadomia composite includes the different terranes of the Cadomian orogen distributed now in Western and Central Europe (peri-Gondwanan West Africa)

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Existence of this Cadomian glaciation could be the reason for the negative excursion of the δ13C curve (Shuram–Wonoka anomaly) (Fig. 18; Halverson et al. 2005). The origin of the Shuram–Wonoka anomaly is under discussion. Le Guerroue et al. (2006) concluded that the anomaly was not related to glaciation, since it lasted 50 Myr, and that it included the Gaskiers glaciation (i.e., it began before the Gaskiers). They also suggested that it did not end until 550 Ma, which suggests that it is not caused by glaciation. Halverson et al. (2005) suggest duration of the Shuram–Wonoka anomaly between c. 590 and c. 560 Ma. A shorter time interval (581 ± 6–567 ± 6 Ma) is inferred by an age model of Retallack et al. (2014). All in all, precise timing and the basic cause for the Shuram–Wonoka anomaly are not well known, but Ediacaran glaciations certainly contributed to that negative anomaly, because glacial events require a response in the ocean water chemistry.

Glacigenic sedimentary and related rocks act as an archive of evolution of the hinterland. The Hf isotope record of detrital and magmatic zircons from investigated strata gives some clues about the nature of the major geologic events forming the Cadomian orogen and its West African hinterland. The Cadomian crustal evolution is dominated by the recycling of continental crust as suggested by the dominance of zircons with negative εHfi values. Juvenile arc magmas became contaminated by the recycling of Eburnean and Archaean crust during long Cadomian magmatic arc activity. A minority of the zircons consists of those derived from the depleted mantle array. Mixture with continental crust is always present. The required geotectonic setting is a continental magmatic arc developed during the Neoproterozoic on stretched Archaean and Palaeoproterozoic (Eburnean) crust.

For the West African hinterland, important episodes of crustal growth and/or recycling can be defined: (1) a 2.2–4.3 Ga old basement became recycled in most cases during Eburnean (c. 1.8–2.2 Ga) orogenic processes; (2) two Eburnean zircons found in the Purpurberg quartzite member and aged at 2056 ± 34 and 2103 ± 40 Ma imply recycling of pre-existing Hadean crust and show TDM model ages of 4.21 and 4.27 Ga, respectively (Jack Hills trend). Hadean crust existed until c. 2.0 Ga; (3) juvenile zircons are scarce; (4) Archean magmas recycled a c. 2.7–3.8 Ga old crust; and (5) zircons with ages of 3.2–3.3 and 3.4–3.5 Ga show a juvenile trend.