Two-phase late Paleozoic magmatism (~ 313–312 and ~ 299–298 Ma) in the Lusatian Block and its relation to large scale NW striking fault zones: evidence from zircon U–Pb CA–ID–TIMS geochronology, bulk rock- and zircon chemistry

Abstract

Late Paleozoic (Variscan) magmatism is widespread in Central Europe. The Lusatian Block is located in the NE Bohemian Massif and it is part of the Saxothuringian Zone of the Variscan orogen. It is bordered by two major NW-trending shear zones, the Intra-Sudetic Fault Zone towards NE and the Elbe Fault Zone towards SW. The scarce Variscan igneous rocks of the Lusatian Block are situated close to these faults. We investigated 19 samples from Variscan plutonic and volcanic rocks of the Lusatian Block, considering all petrological varieties (biotite-bearing granites from the Koenigshain and Stolpen plutons, amphibole-bearing granites from three boreholes, several volcanic dykes, and two volcanites from the intramontane Weissig basin). We applied whole-rock geochemistry (18 samples) and zircon evaporation dating (19 samples). From the evaporation data, we selected six representative samples for additional zircon SHRIMP and CA–ID–TIMS dating. For the Koenigshain pluton, possible protoliths were identified using whole-rock Nd-isotopes, and zircon Hf- and O-isotopes. The new age data allow a subdivision of Variscan igneous rocks in the Lusatian Block into two distinct magmatic episodes. The spatial relation of the two age groups to either the Elbe Fault Zone (298–299 Ma) or the Intra-Sudetic Fault Zone (312–313 Ma) together with reports on the fault-bound character of the dated intrusions suggests an interpretation as two major post-collisional faulting episodes. This assumption of two distinct magmatic periods is confirmed by a compilation of recently published zircon U–Pb CA–ID–TIMS data on further Variscan igneous rocks from the Saxothuringian Zone. New geochemical data allow us to exclude a dominant sedimentary protolith for the Koenigshain pluton as supposed by previous investigations. This conclusion is mainly based on new O- and Hf-isotope data on zircon and the scarcity of inherited zircons. Instead, acid or intermediate igneous rocks are supposed as the main source for these I-type granitoids from the Koenigshain pluton.

Introduction

In the Palaeozoic, the closure of the Rheic Ocean and the amalgamation of the supercontinent Pangaea caused the Variscan orogeny in North America, Europe and Asia. In the Bohemian Massif, located in the easternmost part of the European Variscides, the orogen is subdivided into a series of orogenic zones by subordinate subduction zones (Fig. 1, e.g. McCann 2008). Subduction processes at ca. 340 Ma marked the peak metamorphism and the end of plate convergence in the Saxothuringian Zone of Central Europe (Fig. 1, e.g., Schmädicke et al. 1995; Kröner and Willner 1998; Tichomirowa et al. 2005). During the post-collisional period, transpressional and transtensional tectonics resulted in significant strike-slip displacement along continental scale NW-trending strike-slip faults, that cut and offset the original EW-trending Variscan zones of the Bohemian Massif (e.g., the Saxothuringian and the Tepla-Barrandean Zones, Fig. 1) in the easternmost part of the European Variscides (e.g. Mazur et al. 2020). Recent tectonic models emphasize a repeated reactivation of these structures under different tectonic stress fields (Edel et al. 2018).

Fig. 1

Simplified Variscan tectonic map of Central Europe after Mazur et al. (2020). RS Rhenohercynian Suture, STS Saxothuringian Suture, LB Lusatian Block, EFZ Elbe Fault Zone, ISF Intra-Sudetic Fault Zone, OFZ Odra Fault Zone, DFZ Dolsk Fault Zone, ADF Alpine Deformation Front

Late Paleozoic late- to post-Variscan igneous rocks (combined here under the term Variscan) of the Bohemian Massif are in many cases bound to strike-slip faults (Oberc-Dziedzic et al. 2015) and their radiometric dating may give information about the temporal evolution of the tectonic activity on these structures. Most of the Variscan igneous rocks have been dated repeatedly by different geochronological methods (summarized by, e.g., Förster and Romer 2010; von Seckendorff 2012) and extensively analysed geochemically (e.g., Hammer 1996; Förster et al. 1999; Hammer et al. 1999; Słaby and Martin 2008; Tichomirowa et al. 2019b). Recently, first high-precision zircon U–Pb CA–ID–TIMS data of Variscan granitic rocks allow to reliably differentiate single magmatic pulses of the post-Variscan period indicating distinct short (1–2 Myr) magmatic periods (Kryza et al. 2014a; Tichomirowa et al. 2019a).

The Lusatian Block, the Jizera mountains and the Karkonosze pluton are part of the Saxothuringian Zone of the NE Bohemian Massif (Fig. 1) and constitute a NW-striking terrane that is bound by two major NW striking fault zones, the Elbe Fault Zone (EFZ) and the Intra-Sudetic Fault Zone (ISF; Figs. 1, 2). Towards the north, the ISF splits into the Inner Lusatian and the Main Lusatian Fault (Fig. 2a). The northernmost constituent of the Lusatian-Jizera-Karkonosze Block is the Lusatian Block, which is bound towards NW by the Cambrian sediments of the Torgau-Doberlug syncline (Geyer et al. 2014). Towards SE, The Lusatian Block is confined by the ESE striking normal faults of the Eger Graben (the prolongation of the Krušné Hory Fault towards E, Fig. 2) against the Early Paleozoic igneous rocks of the Jizera mountains, which are intruded by the granitic Variscan Karkonosze pluton (Fig. 2, Kozdrój et al. 2001).

Fig. 2

Geological maps (based on Kozdrój et al. (2001): (a) geological map of the Lusatian (-Jizera-Karkonosze) Block with major structures. b sample locations of the Stolpen pluton, volcanic dykes, and the volcano-sedimentary Weissig basin, c sample locations of amphibole-bearing granitoids and of the biotite-bearing granites of the Koenigshain pluton (Cenozoic sediments and volcanic rocks removed). Abbreviations in a-d are: LTF Lusatian Thrust Fault, SKF Stolpen-Klotzsche Fault, ISF Intra-Sudetic Fault, ILF Inner Lusatian Fault, MLF Main Lusatian Fault, SBF Sudetic Boundary Fault, KHF Krušné Hory Fault, DB Doehlen basin

In the Lusatian Block, the late- to post-Variscan granitoids are subordinate and Variscan volcanic rocks are scarce (Fig. 2). Variscan magmatic activity manifested near the NW striking border faults of the Lusatian Block as granitic plutons, volcanic rocks, and mainly fault-parallel volcanic dykes. These rocks are geochemically and petrologically highly diverse (e.g., Hammer et al. 1999). A relationship between strike-slip-faulting and magmatic activity seems plausible. To understand the tectono-magmatic evolution of the Lusatian Block and its NW striking boundary faults, it is necessary to define precise ages of the fault-bound granitoids. Sources of these rocks need to be understood to interpret the interrelation of faulting and magmatism.

In this study, we present new whole-rock geochemistry and isotope data (Nd- for whole rocks, Hf- and O-isotopes for zircon) of Variscan granitic and rhyolitic-dacitic rocks of the Lusatian Block to infer on sources of these rocks. To determine ages for the igneous rocks, we first applied zircon evaporation dating (Kober et al. 1987) on 19 different Variscan igneous rocks of the Lusatian Block. Based on these data, we chose one sample for zircon U–Pb SHRIMP and six representative samples for high-precision CA-ID-TIMS dating.

Geological setting

Geological setting of the Lusatian Block

The basement of the Lusatian Block is characterized by monotonous flysch-like Precambrian greywacke and pelite sequences (Kröner et al. 1994; Linnemann et al. 2010) that are mainly located in the northern part (Fig. 2a). These sediments were consolidated during the Cadomian orogeny (ca. 570–540 Ma, e.g., Linnemann et al. 2000; Kroner et al. 2007; Linnemann et al. 2010) and intruded by granodioritic plutons at ca. 540–530 Ma (Fig. 2a; Linnemann et al. 2000; Linnemann 2007; Tichomirowa et al. 2001; Tichomirowa 2002) that are mainly located in the southern part of the Lusatian Block (Fig. 2a). The high-grade Variscan metamorphic overprint that characterizes the basement of the adjacent Erzgebirge (Fig. 2a) only affected the rocks of the Lusatian Block along major tectonic lineaments.

During various magmatic events, the Cadomian basement of the Lusatian Block was penetrated by pre-Variscan alkaline basaltic and gabbroic dyke swarms (ca. 400 Ma; Kramer 1977; Kindermann et al. 2003; Abdelfadil et al. 2013; Fig. 2), late Variscan calc-alkaline lamprophyres (330–340 Ma, Abdelfadil et al. 2013), post-Variscan compositionally diverse felsic igneous rocks (290–340 Ma, e.g., Hammer et al. 1999; Förster et al. 2012; Kryza et al. 2014a), and, subordinately, Mesozoic lamprophyric dykes (ca. 230 Ma, Kramer et al. 1977).

Roughly simultaneously with the post-Variscan magmatism, beginning post-orogenic denudation and contemporaneous deformation of the Variscan mountains led to the formation of intramontane Late Carboniferous/Early Permian molasse basins in the Saxothuringian Zone. Since the basins were uplifted and largely eroded together with the orogen, most of the present intramontane basins represent erosional remnants (Schneider and Romer 2010). Several such basins surround the Lusatian Block, namely the North Sudetic basin, the Intrasudetic basin, the Karkonosze piedmont basin, and the Doehlen basin (Fig. 2a). The Lusatian Block itself is free of molasse sediments, except for the 2.7 km long and 1 km wide Weissig basin, which is located in the prolongation of the Lusatian Thrust Fault of the EFZ (Fig. 2b). The basin fill of the Weissig basin comprises ca. 350 m of interlayered sediments and volcanic rocks. Altogether, ca. 150 m of the basin fill consists of crystal-rich intermediate lava and crystal tuff (Reichel 2012). The stratigraphic sequence of the Weissig basin is subdivided into two formations, namely the Hutberg Formation that predominates in the northwestern part of the basin and the Napoleonstein Formation that predominates in the southeastern part (Reichel 2012). Further dismembered blocks of Permian volcanic and sedimentary rocks (Radebeul, Rossendorf, Rosinendörfchen, Vlči hora, Hodkovice nad Mohelkou; Reichel 2012; Huhle and Lange 2010; Fig. 2a) appear along the Lusatian Thrust Fault.

Variscan magmatism in the Lusatian Block

Variscan plutons are encountered in the entire Bohemian Massif (Fig. 1, Cháb et al. 2007). In the Lusatian-Jizera-Karkonosze Block, the largest of these plutons is the ca. 70 km long and > 20 km wide Karkonosze Pluton, while smaller plutons occur in the northern part (Koenigshain and Stolpen plutons; Fig. 2a–c).

The Karkonosze composite pluton was studied in detail by several authors. Based on whole-rock geochemistry and petrologic evidence, the various lithologies of the Karkonosze composite pluton have been interpreted as the product of mixing of a peraluminous crustal magma with a mafic magma (Słaby and Martin 2008). Epsilon Nd data vary from -7 to -1 and are specific for single rock types of the pluton. These data were interpreted as mixing between two endmembers: (i) an enriched mantle source and (ii) a granitic crustal source (Słaby and Martin 2005). The concept of mixing is supported by the existence of mafic magmatic enclaves and syn-plutonic dykes (Barbarin 2005). The existence of chemically zoned feldspar-phenocrysts and chemical modelling indicates that fractional crystallization also took place during the evolution of the Karkonosze granite (Słaby and Götze 2004). Several attempts have been made to determine the intrusion age and sequence of the various rock types of the Karkonosze pluton using different methods and minerals (e.g., Duthou et al. 1991; Kröner et al. 1994; Kusiak et al. 2009). They thereby observed large scatter in intrusion ages (290–340 Ma) contradicts recent high-precision zircon U–Pb CA-ID-TIMS data that display ages between 312.5 ± 0.3 and 312.3 ± 0.3 Ma for the two main granite facies (Kryza et al. 2014b).

Smaller plutons as well as other subsurface granites were studied in less detail. Petrologically, these rocks were subdivided into amphibole- and biotite-bearing intrusions (Fig. 2b, c, Hammer 1999). The monzogranite of Wiesa, the granodiorite (tonalite) of Kleinschweidnitz (Eidam et al. 1995; Hammer 1996), and a granodiorite encountered in a borehole northeast of Arnsdorf (borehole “Ober Prauske”, Hammer 1996) represent amphibole-bearing granitoids. The most prominent of the biotite (± muscovite-)-bearing intrusions is the massif of Koenigshain/Arnsdorf (Fig. 2c), which is composed of three textural varieties: equigranular leucogranite, porphyritic granite, and fine-grained monzogranite (Eidam and Götze 1991; Hecht et al. 1999). The biotite-bearing rocks also include the monzogranites of Stolpen (Hammer 1996). The slightly NW elongated shape and the spatial proximity of the Stolpen pluton to the Stolpen-Klotzsche Fault (Fig. 2b, parallel to the EFZ) support the assumption that the magma intruded along this fault (Kozdrój et al. 2001; Lisowiec et al. 2014). The Koenigshain pluton is bound to the Inner Lusatian Fault (Fig. 2c) and provides petrologic evidence for a relationship between faulting activity and magma intrusion (Hammer 1996; Thomas and Davidson 2016).

Magma sources of the granitoid rocks have been proposed based on whole-rock geochemistry and Sr-, Nd-, O-, and H-isotopic data (Hammer 1996). According to these data, the amphibole- and biotite-bearing granites had different sources. Hammer (1996) interpreted the amphibole-bearing granites as melts formed from mafic amphibolite-facies rocks of the lower crust. Distinctly, the biotite-bearing granites are suggested to be water-undersaturated partial melts from metatonalitic or metapelitic protoliths (Hammer 1996). The composition of altered domains and the mineralogy of secondary accessories indicate a local influence of post-magmatic fluids to the Stolpen biotite-bearing granite. Enrichment in some HFSE and LILE (Th, U, Y, Hf, Nb, and Ta) and a high halide content (fluorite) of altered domains indicate a mantle origin of these fluids (Lisowiec et al. 2013). Different minerals and methods have been used to determine the intrusion age of the granitoid rocks of both the Koenigshain pluton and the Stolpen pluton (Fig. 3; Kröner et al. 1994; Hammer et al. 1999; Thomas et al. 2009; Förster, et al. 2012; Lisowiec et al. 2014). The age data span a range from 331 to 290 Ma and are partly contradictory. Consequently, the intrusion sequence of the granitoid rocks still has to be resolved (Fig. 3).

Fig. 3

Published age data from late- to post-Variscan granites of the German part of the Lusatian Block. Data are from (1) Hammer et al. (1999), (2) Förster et al. (2012), (3) Lisowiec et al. (2014), (4) Kröner et al. (1994), (5) Reichel et al. (2012), (6) Barthel et al. (2010)

Variscan volcanic dykes of rhyolitic to rhyodacitic composition accompany and partly cross-cut the granite plutons of Stolpen and, to a lesser amount, of Koenigshain (Kozdrój et al. 2001; Fig. 2b, c). Near the Weissig basin and the Stolpen pluton, almost all dykes strike NW, roughly parallel to the EFZ. In addition, some dykes strike NE to ENE (Fig. 2b). Subordinate NE to ENE striking dykes prevail near the Koenigshain pluton (Fig. 2c). These dykes, the lavas, and tuffs from the Weissig basin have not yet been dated by radiometric methods. For the Weissig basin, age estimates based on fossil insect wings indicate an Asselian age for the sediments interlayered with volcanic rocks (Schneider and Werneburg 2012; Reichel 2012).

Materials and methods

Samples

This study aimed to acquire intrusion ages for Variscan igneous rocks of the northern Lusatian Block (Fig. 2b, c). Four samples were taken from the Koenigshain pluton (BGK1, BGK2, BGK3, BGK4) and three samples from the Stolpen pluton (BGSt11, BGSt13, BGSt14; Fig. 2b, c) to cover different textural varieties of the two biotite-bearing granitic plutons. From the amphibole-bearing granitoids, three samples from different drill holes were chosen (AG59, AG61, AG65; Fig. 2c). Exemplarily for the widespread rhyolitic and rhyodacitic dykes of the Lusatian Block, we sampled eight dykes around the Stolpen pluton (VR03, VR09, VR10, VR11, VR15, VR20, VR28, VR34; Fig. 2b). Most of the dyke samples represent loose blocks of rocks found at surface. From the Weissig basin, we selected a porphyritic tuff of the Napoleonstein Formation (VRW65) and a porphyritic lava from the Hutberg Formation (VRW67). Table 1 summarizes the sample numbers, localities, coordinates, and rock types of studied samples.

Table 1 Sample locations and description

Bulk-rock geochemistry

18 out of 19 whole-rock samples were analysed for major and trace element contents at Activation Laboratories (Actlabs Canada; “4 Litho” research analytical protocol) by Fusion-ICP and Fusion-MS, respectively. Samples were fused with lithium metaborate/tetraborate and afterward diluted and analysed by Perkin Elmer Sciex ELAN 6000, 6100, or 9000 ICP/MS. Three blanks and five controls (three before the sample group and two after) were analysed per group of samples. Duplicates were fused and analysed every 15 samples. The instrument was recalibrated every 40 samples. Reproducibility was better than 1% for major elements and better than 5% for trace elements based on analyses of certified standards.

Zircon separation and dating

Zircons have been extracted from all 19 samples by the usual procedure (crushing, Wilfley table, Frantz magnetic separator, heavy liquids, final handpicking). The selected zircons were characterized using secondary electron (SE) images to evaluate their morphology. Cathodoluminescence (CL) images of the same grains visualize their internal structure. Energy-dispersive X-ray spectroscopy (EDX) was used to determine the elemental composition of mineral inclusions.

Zircon U–Pb CA–ID–TIMS dating is a highly precise and accurate, but also time-consuming, laborious, and expensive method. A preceding sample selection, based on zircon characteristics like morphology and internal structure and a more rapid dating approach like zircon U–Pb LA–ICP–MS or the evaporation method, can be used to select the best suited samples for additional CA–ID–TIMS and SHRIMP dating. Here, we used the zircon evaporation method because this is a routine method in our laboratory.

Zircon dating by evaporation (Kober method)

All 19 samples were analysed with single zircon evaporation at TU Bergakademie Freiberg following the methods outlined by Kober (1987). Filament assemblies were mounted on a Finnigan MAT262 mass spectrometer. Before zircon evaporation, the second (ionization) filament was heated to 1800 °C to strip the filament from possible lead-bearing phases (additional outgassing). The evaporation filament was then heated to 1450 °C to remove common lead hosted in less stable phases of the zircon grain. Evaporation was performed at 1600 °C after cooling the ionization filament. This was done in one step to obtain high signal intensities for measurement. Data acquisition was performed by peak jumping using a secondary electron multiplier equipped with an ion counter with mass sequence 207–206–204–206–207 (counting time in seconds 4–4–8–4–4, respectively). A mean value was calculated from the two ²⁰⁷Pb/²⁰⁶Pb and ²⁰⁴Pb/²⁰⁶Pb ratios of each scan to minimize intensity changes during measurements. Ten blocks (composed of ten scans) were recorded corresponding to 90 scans per measurement. Since ²⁰⁴Pb/²⁰⁶Pb ratios bear a large uncertainty due to the low intensity of ²⁰⁴Pb, a trend line was defined through subsequently measured ²⁰⁴Pb/²⁰⁶Pb ratios.

Temora 2 was dated as a secondary standard as suggested by Horstwood et al. (2016; analytical data are given in Supplement 4). The mean evaporation age of secondary standard Temora 2 is with 419.3 ± 3.0 Ma ca. 1% older, but still within error identical to published values (416.8 ± 0.3 Ma, Black et al. 2003, 417.5 ± 0.1 Ma and 417.3 ± 0.1 Ma, von Quadt et al. 2016). According to these data and the primary standard 91500, the precision and accuracy of our zircon evaporation data is ca. 2% (Supplement 4).

Zircon U–Pb SHRIMP dating

Zircons from one sample (BGK1) were additionally analysed by the SHRIMP II technique (Sensitive High mass Resolution Ion MicroProbe) at the Centre of Isotopic Research (VSEGEI, St. Petersburg, Russia). Each analysis consisted of 5 scans through the mass range. The spot diameter was about 18 μm, and the primary beam intensity was about 4 nA. The data have been reduced in a manner similar to those used by Williams (1998, and references therein), using the SQUID Excel Macro of Ludwig (2000). The zircon standard Temora2 was used for reference of the U/Pb ratio and concentrations (Black et al. 2003). Primary standard data cannot test the accuracy of a dating method, because these standard data are used to calibrate the unknowns and match by definition the accepted values. Nevertheless, these data are valuable, because they illustrate the precision of the method. Therefore, we present primary standard values for SHRIMP (Temora2, Supplement 4, Tichomirowa et al. 2019a). Accordingly, the precision of our zircon U–Pb SHRIMP dating is at about 1–2% (see also Schaltegger 2015).

To trace the accuracy of any dating method, secondary reference materials should be dated together with the unknowns (Horstwood et al. 2016). Our SHRIMP data were acquired in 2011. This is the reason why a secondary standard was not analysed by the SHRIMP method.

Corrections for common lead (Pb_c) were made using measured ²⁰⁴Pb and by applying the Pb evolution model of Stacey and Kramers (1975). Uncertainties given for individual analyses (ratios and ages) are at the 1σ level, for calculated Concordia ages at the 2σ level.

Zircon U–Pb CA–ID–TIMS dating

For this labour-intensive method, we chose 1–2 representative samples from each of the biotite-bearing granites of Koenigshain (BGK1, BGK4) and Stolpen (BGSt11, BGSt14), the amphibole-bearing granitoids (AG61), and the volcanic rocks of the Weissig basin (VRW67).

Selected zircon grains (ca. 30—50 per sample) were annealed for 72–96 h at 900 °C, and subsequently chemically abraded for 12 h at 210 °C with concentrated HF and HNO₃ in a pressurized Parr dissolution vessel. This procedure dissolves crystal domains with strong radiation damage which are suspected to have experienced post-crystallization lead loss (Mattinson 2005). The removal of Pb loss after application of this lab protocol has been proven for the Plešovice reference zircon but does not need to be appropriate equally for zircons from other samples (Widmann et al. 2019). Afterwards, the acid together with dissolved zircon material was completely pipetted out and 3.5N HNO₃ was added to the remaining zircons grains and fragments and left for 30 min at 50 °C to remove surface lead. Several cleaning cycles with water combined with repeated ultrasonic treatment were conducted before single zircon fragments were selected for further processing. Single zircon grains/fragments were washed with 3.5N HNO₃ and transferred into cleaned microcapsules with a small drop of this fluid and four drops of concentrated HF. Samples were spiked with a ²⁰⁵Pb- ²³³U-²³⁵U- tracer solution (ET535 at TU Bergakademie Freiberg, Condon et al. 2015). For final dissolution, the microcapsules were placed in pressurized Parr dissolution vessels and heated to 200 °C for 48 h, followed by drying at 130 °C and then re-dissolution in 6N HCl for 24 h at 200 °C to transfer them into chlorides. After repeated drying, the samples were dissolved in ten drops of 3.1N HNO₃ and transferred into micro-columns for column chemistry. U and Pb were separated from the rest of the sample by anion exchange chromatography using HCl and H₂O. The U and Pb containing fraction was loaded on pre-degassed rhenium filaments with a drop of silica gel (Gerstenberger and Haase 1997) and measured with a Finnigan Mass Spectrometer MAT262 using a secondary electron multiplier (SEM). Alternatively, the samples were measured on an IsotopX Phoenix Mass Spectrometer using Daly ion counter and/or Faraday collectors (ATONA). The comparability of the results of both mass spectrometers was proven by repeated measurement of zircon standards 91500 (Wiedenbeck et al. 1995) and Temora (Black et al. 2004). The published ages of Temora2 are 416.8 ± 0.3 Ma (Black et al. 2003), 417.5 ± 0.1 Ma for SEM-measurements and 417.3 ± 0.1 Ma for measurements on Faraday cups (von Quadt et al., 2016). Our date of 417.3 ± 0.6 Ma (Supplement 4) perfectly matches these values. Few analyses of Temora 2 show older ages (418–431 Ma) that can be interpreted as the presence of slightly older zircon grains in the Temora diorite and were not used for the mean age calculation. Additionally, the accuracy of zircon U–Pb CA–ID–TIMS ages was monitored by dating the standard 91500. This standard was determined to be 1062.4 ± 0.4 Ma (Wiedenbeck et al. 1995) or 1063.6 ± 0.3 Ma (Schoene et al. 2006). Our weighted mean ²⁰⁶Pb/²³⁸U-age of 1064.6 ± 1.3 Ma (Supplement 4) is within 0.1% of the accepted values. Based on the results of standard dating, we presume the here presented CA–ID–TIMS ages to be accurate on the 0.1% level.

In contrast to Pb–Pb evaporation and U–Pb SHRIMP dating, the CA–ID–TIMS method is not calibrated by an external zircon sample (e.g. a calibrated zircon reference) but by a mixed U–Pb tracer solution. Currently, the use of precisely and accurately calibrated EARTHTIME tracers reduced the inter-laboratory bias to 0.1% (Condon et al. 2015). Consequently, the CA–ID–TIMS method can reliably yield accurate U–Pb dates (e.g., Schaltegger et al. 2015) if this well-calibrated tracer is used.

Bulk rock Sm- and Nd-isotopes

About 200 mg of powder from whole‐rock was dissolved in 50% HF‐12N HNO₃, then attacked with 8N HNO₃ and finally with 6N HCl. Samarium and neodymium were separated by ion-exchange resins. The isotope ratios were measured on a Finnigan MAT262 spectrometer and the quoted errors are given at the 2σ level. Concentrations of Sm and Nd were obtained by isotope dilution. The ¹⁴³Nd/¹⁴⁴Nd ratios were normalized to ¹⁴⁶Nd/¹⁴⁴Nd = 0.7219 (DePaolo 1981). Sm‐Nd model ages were calculated using the depleted mantle model (TDM, Liew and Hofmann 1988). The mean value for ¹⁴³Nd/¹⁴⁴Nd of the standard JNdi was 0.512098 ± 0.000010 (n = 8).

Zircon geochemistry: Hf- and O-isotope ratios

For sample BGK1, in addition to in situ U–Pb zircon ages analysed by SHRIMP, the Hf- and O-isotope composition was determined from the same spots on these zircon grains. Hafnium isotopes were measured on a Thermo-Finnigan Neptune multi-collector ICP–MS coupled to a Resonetics 193 nm ArF excimer laser (CompexPro 102, Coherent) system at Goethe-University Frankfurt (GUF) (Gerdes and Zeh 2006). Spots of 40 to 60 μm in diameter were ablated with a repetition rate of 5.5 Hz and an energy density of 5 J/cm² during 55 s of data acquisition. All data were adjusted relative to the JMC475 standard (¹⁷⁶Hf/¹⁷⁷Hf = 0.282160) and quoted uncertainties are quadratic additions of the within-run precision of each analysis and the reproducibility of JMC475 (2 SD = 0.0028%, n = 6). We verified the accuracy and external reproducibility by repeated analysis of the reference zircons, Temora and GJ-1. They yielded ¹⁷⁶Hf/¹⁷⁷Hf ratios of 0.282689 ± 0.000023 (2 SD, n = 11 for Temora), 0.282012 ± 0.000014 (2 SD, n = 8 for GJ-1). This is 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 (2006–2012) of GJ-1 (0.282010 ± 0.000025; n > 800) reference zircon at GUF.

Zircon oxygen isotopes were measured with the Cameca IMS 1280 multicollector ion microprobe at the Swedish Museum of Natural History (Heinonen et al. 2015), utilizing a ~ 2 nA Cs⁺ primary ion beam together with a normal incidence low-energy electron gun for charge compensation, medium field magnification (~ 80x), and two Faraday detectors (channels L2 and H2) at a typical mass resolution of ~ 2500. Measurements were performed in pre-programmed chain analysis mode with automatic field aperture and entrance slit, centered on the ¹⁶O signal. The magnetic field was locked using NMR regulation for the entire analytical session. Each data acquisition run comprised a 20 × 20 μm pre-sputter to remove the Au layer, followed by the centering steps, and 64 s of data integration performed using a non-rastered, ~ 10 × 10 μm spot. In the measurement chain, every set of four unknowns was followed by two bracketing analyses on the 91500 standard zircon. A δ¹⁸O value of + 9.86‰ (SMOW, Wiedenbeck et al. 2004) was assumed for the 91500 zircon in data normalization, and small linear-drift corrections were applied to each session. External reproducibility of ± 0.3‰ (1 SD) based on measurements on the standards was propagated into the overall uncertainty for each analysis.

Results

Bulk rock geochemistry

Whole-rock geochemical data of 18 samples are presented in Table 2. Classification of the samples based on major element data is presented in Fig. 4. REE- and Multi-element diagrams are shown in Fig. 5.

Table 2 Whole-rock major and trace element contents

Fig. 4

Geochemical classification of the investigated samples based on whole-rock major element data. a Total-alkali-silica (TAS) diagram according to Le Bas et al. (1986), based on analyses recalculated on volatile-free basis, b SiO₂-K₂O plot according to Peccerillo and Taylor (1976). The sample groups are BGK biotite-bearing granite of the Koenigshain pluton, AG amphibole-bearing granite, BGSt biotite-bearing granite of the Stolpen pluton, VR volcanic rocks in dykes, VRW volcanic rocks within the Weissig basin

Fig. 5

Upper continental crust (UCC, Rudnick and Gao, 2013) normalized REE patterns (left column) and multi-element patterns (right column). The data for each sample are shown as a coloured lines. For better comparison between a–e, the range covered by all samples is shown as grey shaded area in all plots

All analysed magmatic and volcanic rocks are rich in SiO₂ (61–79 wt.%, Table 2), plotting in a TAS-diagram (Le Bas et al. 1986) from granodiorite and quartz monzonite to high-SiO₂-granite and from trachydacite and dacite to rhyolite, respectively (Fig. 4a). The biotite-bearing granites of the Koenigshain and the Stolpen plutons plot at the high-SiO₂-end (76–78 wt.% SiO₂) and the amphibole-bearing granitoids generally have lower SiO₂ contents and reveal a greater variability from 64 to 69 wt.% SiO₂. These rocks are mainly classified as granodiorites and quartz-monzonites. The rhyolitic to trachydacitic volcanic dykes vary over a wider range of SiO₂ contents (65–79 wt.%), whereby some rhyolitic dykes are slightly more SiO₂-rich then the granitic rocks (Fig. 4a). The volcanic layers of the Weissig basin reveal similar SiO₂ contents (61–67 wt.%) and are classified as dacite and trachyte (Fig. 4a).

The Na₂O/K₂O-ratio of all samples is ≤ 1.3 for all rocks, classifying them as high-K calc-alkaline to shoshonitic rocks (Fig. 4b: K₂O vs. SiO₂ diagram, Peccerillo and Tailor 1976). In this diagram, biotite-bearing granites are restricted to the high-K calc-alkaline field, but amphibole-bearing granitoids reveal a wider dispersion from the high-K calc-alkaline to the shoshonitic field. Volcanic dykes show the widest scatter with the highest K₂O contents in some samples. All analysed rocks are peraluminous with A/CNK [Al₂O₃/(CaO + Na₂O + K₂O)] ranging from 1.0 to 1.3.

The REE patterns of the biotite-bearing granites of Koenigshain and Stolpen and some of the volcanic dykes show, compared to the upper continental crust (UCC, Rudnick and Gao 2014), a significant negative Eu-anomaly (0.1–0.3, Table 2, Fig. 5a). In these samples, all other REE contents are similar to the upper continental crust with a slight enrichment of the HREE (LaN/YbN-ratios = 0.1–0.8, Table 2). The amphibole-bearing rocks, the volcanic rocks of the Weissig basin, and three volcanic dykes have only weakly negative or absent Eu-anomalies (Eu*/Eu = 0.7–1.1, Table 2) and do not show an enrichment of HREE over LREE compared to UCC (LaN/YbN = 0.6–1.9, Table 2). Their absolute REE contents are similar to that of the upper continental crust, or shifted towards slightly higher contents.

In upper-continental-crust-normalized multi-element diagrams (Fig. 5, Rudnick and Gao 2014), the biotite-bearing granites of Koenigshain and Stolpen and most of the volcanic dykes are slightly enriched in LILE and show significant negative anomalies of Ba, Sr, Ca, Eu, Zr, and Ti. The negative anomalies of Sr and Ca are only weakly pronounced in the amphibole-bearing rocks and one volcanic dyke and show intermediate values for the volcanic rocks of the Weissig basin and two volcanic dykes. The negative Eu, Ba, and Zr anomalies are mostly absent in these rocks.

From bulk rock geochemical data, we calculated crystallization temperatures according to Watson and Harrison (1983) and according to Jung and Pfänder (2007; Table 2). The scarcity of inherited zircons (Tables 3, 4, Supplement 1) in most samples allows for the calculation of zircon saturation temperatures (Watson and Harrison 1983) as melt temperatures estimates. With both methods, the crystallization temperatures for the biotite-bearing granites of the Koenigshain and Stolpen plutons are roughly identical with 734–790 °C for the zircon saturation temperatures (Watson and Harrison 1983) and 658–809 °C according to Jung and Pfänder (2007). The amphibole-bearing granites revealed higher temperatures for both geothermometers with 788–828 °C (Watson and Harrison 1983) and 897–913 °C (Jung and Pfänder 2007). The calculated temperatures of the volcanic dykes roughly vary with SiO₂ content (Fig. 6). For volcanic dykes with SiO₂ < 75 wt.%, the calculated temperatures according to Watson and Harrison (1983) resulted in 847–899 °C indicating even higher temperatures compared to those for amphibole-bearing granitoids, while similar temperatures result from calculations according to Jung and Pfänder (2007, 831–972 °C). Volcanic rocks with SiO₂ > 75 wt.% have zircon saturation temperatures of 746 to 772 °C and magma temperatures of 698–725 °C that overlap with those temperatures calculated with the same methodology for the biotite-bearing granites of the Stolpen and Koenigshain plutons.

Table 3 Single spot zircon U–Pb SHRIMP dates and isotopic data of sample BGK1

Table 4 U–Pb CA-ID-TIMS single-grain zircon dates and isotopic data

Fig. 6

Zircon saturation temperatures (a) calculated according to a Watson and Harrison (1983) and magma temperatures (b) calculated according to Jung and Pfänder (2007) vs. SiO₂ content of the whole-rock samples. The sample groups are BGK biotite-bearing granite of the Koenigshain pluton, AG amphibole nearing granite, BGSt biotite-bearing granite of the Stolpen pluton, VR volcanic rocks in dykes, VRW volcanic rocks within the Weissig basin

Zircon ages

Zircon characterization and Pb–Pb evaporation age determination

Following zircon characterization, single-grain zircon Pb–Pb evaporation was performed on selected grains of each sample. Figure 7 shows representative SE and CL images of zircons from each group of samples (except for amphibole-bearing granites) and Pb–Pb evaporation data of single zircons. Based on these data, we selected samples for additional CA–ID–TIMS and SHRIMP dating. Samples with zircons frequently showing Pb loss and/or elevated Pb_c, and/or inherited ages were excluded from further CA–ID–TIMS dating (Fig. 7). Detailed results of Pb–Pb evaporation analysis are tabulated in Supplement 1.

Fig. 7

Summary of zircon Pb–Pb evaporation data that were used as screening to select samples that are best suited for high-precision zircon U–Pb CA–ID–TIMS dating. Selected samples should have zircons with low amounts of inherited grains, low amounts of common Pb (²⁰⁴Pb/²⁰⁶Pb ≤ 0.0002), and low degree of Pb loss. Grains with ²⁰⁴Pb/²⁰⁶Pb > 0.0002 were discarded and are not presented in the Tables (Supplement 1) and in this Figure. For each sample, the left column (red text) gives the amount of single-grain ages (in %) that have been removed from the dataset because of elevated common Pb (²⁰⁴Pb/²⁰⁶Pb > 0.0002). Representative CL images are given for each sample group, if available. Weighted mean ages were calculated with Isoplot/Ex (Ludwig 2008). Uncertainties on mean ages are 95%-confidence errors. The zircon Pb-Pb evaporation data define two groups of crystallization ages, marked as coloured bands

Zircons of the biotite-bearing granites of the Koenigshain pluton show oscillatory zoning in their interior zones covered by broad black CL rims (Fig. 7a). These zircons contain inclusions of apatite, biotite, and muscovite and, less frequently, quartz, feldspar, and ilmenite. The zircons show dominance in {101} pyramids and some variation in prims from {110} to {100} (Hammer 1996). Zircon Pb–Pb evaporation dating yielded similar mean ages for all four samples from Koenigshain with only very few younger grains (interpreted as Pb loss) and few older, inherited grains. The calculated mean intrusion ages of the four samples from Koenigshain vary from 319 ± 7 Ma to 324 ± 4 Ma (Fig. 7). Despite some single grains with elevated Pb_c (²⁰⁴Pb/²⁰⁶Pb > 0.0002), all four Koenigshain samples seem equally suitable for zircon U–Pb CA–ID–TIMS and SHRIMP dating. The sample BGK1 was chosen for additional SHRIMP and CA–ID–TIMS dating, and sample BGK4 was chosen in addition for CA–ID–TIMS dating to obtain ages from the two small plutons separated by the granodioritic Cadomian basement and a fault (Fig. 2c).

Zircons of the biotite-bearing granites of the Stolpen pluton display oscillatory zoning, similar to those of the biotite-bearing rocks of the Koenigshain pluton (Fig. 7b), but without broad black CL rims. Frequent inclusions are biotite, apatite, quartz, and feldspar. Similar to the biotite-bearing granites of the Koenigshain pluton, zircons show dominance in {101} pyramides and {100} as well as {110} prisms (Hammer 1996). Zircon Pb-Pb evaporation data of these samples revealed minor Pb loss and only few inherited ages. The calculated mean ages of the three samples vary from 304 ± 2 Ma to 308 ± 5 Ma and are therefore significantly younger than the samples from the Koenigshain pluton (Fig. 7b). BGSt11 and BGST14 were chosen for additional CA-ID-TIMS-dating.

Amphibole-bearing granitoids can be distinguished from the biotite-bearing granites by their zircon morphology showing co-existence of {100} prisms and {211} pyramids (Hammer 1996). In zircon Pb–Pb evaporation data, no Pb loss was detected but a significant portion of inherited ages was obtained that often were only slightly older than the calculated intrusion ages (from 322 ± 5 to 325 ± 5 Ma). Sample AG61 was selected for CA-ID-TIMS dating because the inherited ages are much older compared to the intrusion ages and therefore can easily be distinguished, while Pb_c contents in most grains were low (Fig. 7c).

Zircons from volcanic dykes share the morphology and internal structure of both biotite-bearing granitic plutons. Inclusions in zircons are smaller than those in the biotite-bearing granites and have been identified as apatite, biotite, quartz, feldspar, and ilmenite. The zircon evaporation data scatter over a wide range with calculated weighted mean ages from 296 ± 6 to 326 ± 3 Ma. Most of the samples display both inheritance and Pb loss (Fig. 7d). Elevated Pb_c rarely occurs in zircons of most volcanic dykes. None of the volcanic dyke samples was dated with zircon U–Pb CA–ID–TIMS.

Although zircon characteristics of the volcanic rocks of the Weissig basin do not differ from volcanic dykes, slight rounding of zircon grains can be recognized, particularly in sample VRW67 from the Hutberg Formation of the Weissig basin (Fig. 7e). The weighted mean ages for samples VRW65 and VRW67 are 306 ± 4 Ma and 304 ± 4 Ma, respectively. Single grain zircon Pb–Pb evaporation ages reveal inheritance in both samples of the Weissig basin and additional Pb loss in sample VWR67. Nevertheless, this sample was chosen for zircon U–Pb CA–ID–TIMS dating because it contains few inherited grains and few grains with elevated Pb_c.

From the new zircon evaporation age data, the biotite and amphibole-bearing granitic rocks of the northern Lusatian Block can be subdivided into two well-defined groups (Fig. 7). The older group comprises the biotite-bearing granites of the Koenigshain pluton (mean ages from 319 ± 7 to 324 ± 4 Ma), the amphibole-bearing granites (mean ages from 322 ± 5 to 325 ± 5 Ma), and probably one volcanic dyke (VR15, mean age of 326 ± 3 Ma). The younger group comprises the biotite-bearing Stolpen granites (from 304 ± 2 to 308 ± 5 Ma), six out of seven volcanic dykes (VR03, VR15, VR20, VR28, VR34, from 296 ± 6 to 304 ± 4 Ma), and probably the volcanic rocks from the Weissig basin (from 304 ± 4 to 306 ± 4 Ma, although several single measurements yielded ages up to 320 Ma).

Zircon U–Pb age determinations by SHRIMP

Sample BGK1 was dated using the zircon U–Pb SHRIMP method. The results of these analyses are given in Table 3 and Supplement 2. The single measurements range from 321 ± 4 to 327 ± 3 Ma (Table 3). A Concordia age of 323.9 ± 2.1 Ma (MSWD of concordance = 0.42) was calculated from all eleven spot analyses (Figure in Supplement 2). This age is ca. 4 Ma older than the zircon Pb-Pb evaporation date of this sample (319 ± 2 Ma) although both ages overlap within errors.

Zircon U–Pb age determinations by CA–ID–TIMS

Zircon U–Pb CA–ID–TIMS isotopic results for six samples are presented in Table 4 and shown as ²⁰⁶Pb/²³⁸U ranked age plots in Fig. 8. For each sample, 10–23 grains were analysed by this method. Mean sample ages representing the crystallization event were calculated from established age clusters with the software ET_Redux (Bowring et al. 2011). The error includes the internal 2σ measurement error, the tracer calibration uncertainty, and the uncertainty of the decay constant, allowing for a comparison with ages determined by different dating methods (Schoene et al. 2006: z error).

Fig. 8

Single grain zircon U–Pb CA–ID–TIMS analyses as ²⁰⁶Pb/²³⁸U weighted mean dates for granitic and volcanic samples, calculated with the software ET_Redux (Bowring et al. 2011). Each vertical bar represents a single zircon grain analysis including its 2σ analytical (internal) uncertainties whereas data represented by grey bars are not included in the weighted mean calculation. Horizontal black lines represent the weighted mean age

Analysis of 14 zircons from sample BGK1 yielded ages between 308.6 ± 0.7 Ma and 316.2 ± 1.2 Ma (Fig. 8, Table 4). Four zircons yielded identical ages resulting in a mean age of 312.7 ± 0.4 Ma (n = 4) which is interpreted as the intrusion age. Older ages are supposed to be zircons incorporated during intrusion from slightly older magma batches (antecrystic grains, Miller et al. 2007). Single ages that are younger than the cluster used for mean age calculation are interpreted as slight Pb loss.

A second sample (BGK4) was dated from the same pluton. The first attempt to date this sample (17 zircons, analysis numbers 1–12 and14–18 in Table 4) revealed variable ages from 302.4 ± 0.4 to 316.0 ± 0.5 Ma. We interpreted the large age scatter towards younger ages as Pb loss. For this reason, we applied an additional leaching procedure to the remaining zircons (24 instead of 12 h, analysis number b-1 to b-6 in Table 4). The single ages from zircons leached 24 h show less scatter (from 311.2 ± 0.6 to 313.9 ± 0.4 Ma) and support our suggestion of Pb loss for the youngest single dates. An age cluster at 312.9 ± 0.4 Ma (n = 3) is interpreted as the intrusion age of this sample from three zircon ages with 24 h leaching (Fig. 8a). From this data set, two single-grain ages are up to 1 Ma older than this cluster and are interpreted as antecrystic grains. Within error, the intrusion ages of both samples from the Koenigshain granite are equal (312.7 ± 0.4 and 312.9 ± 0.4 Ma).

Two samples were dated from the biotite-bearing granite of Stolpen. Two out of 11 single-grain ages of sample BGSt11 are obviously (~ 20 Ma) older than the main age cluster (Table 4, 318.2 ± 0.9 Ma, 323.3 ± 0.8 Ma). From the remaining 9 single-grain dates (ranging from 294.5 ± 0.3 to 298.5 ± 0.5 Ma), the oldest dates are concordant and form a cluster, while younger ages are slightly (but increasingly) discordant (Supplement 3). Thus, we interpret the oldest group as the intrusion age of this sample (Fig. 8b, 298.3 ± 0.4 Ma, n = 3). Based on their discordancy, the slightly younger ages (up to 4 Ma) are interpreted to result from Pb loss (e.g., Mezger and Krogstad 1997; Schoene 2014).

All age data of sample BGSt14 cluster between 295.3 ± 2.1 Ma and 299.1 ± 0.4 Ma (Fig. 8, Table 4). As in sample BGSt11, the Concordia diagram displays discordant pattern for younger ages (< 298 Ma, Supplement 3). Consequently, we used the oldest age cluster for the calculation of the intrusion age (298.4 ± 0.4 Ma, n = 3). One single slightly older age (299.1 ± 0.4 Ma) is interpreted as antecrystic zircon. The resulting mean ages of both samples from the Stolpen granite are therefore identical within error (298.3 ± 0.4 and 298.4 ± 0.4 Ma).

Five out of 16 single-grain analyses from the amphibole-bearing granite (sample AG61) yielded much older ages than the remaining dates (Table 4, 356 ± 2 to 537.7 ± 0.3 Ma). We interpret these data as inherited grains. Two distinctly younger dates at ca. 309 Ma are interpreted as Pb loss. The remaining seven single-grain ages define a cluster with a weighted mean age of 312.1 ± 0.4 Ma (n = 4, Fig. 8). A tail towards older ages is interpreted as antecrystic zircon grains or mixed ages.

From the volcanic rock from the Weissig basin (sample VRW67) ten single zircon grains were analysed. Besides one antecrystic grain (302.6 ± 1.0 Ma) and slight Pb loss in three grains, an age cluster from six analyses yielded a weighted mean age at 299.1 ± 0.4 Ma. This age is interpreted as the eruption age of this volcanite (Fig. 8, Table 4).

Whole-rock Nd- and Sr-isotope ratios

Nd- and Sr-isotopes were determined for the four biotite-bearing granites of Koenigshain (samples BGK1, BGK2, BGK3, BGK4) and are listed in Table 5. The εNd values of these samples range from − 4.2 to − 0.8 (t = 320 Ma) with two-stage model ages according to Liew and Hofmann (1988) ranging from 1.1 to 1.4 Ga.

Table 5 Sm–Nd isotope data

Zircon geochemistry: Hf- and O-isotope ratios

The measured zircons from sample BGK1 have relatively uniform Hf and very homogeneous O-isotopic compositions (Table 6, Supplement 4). The hafnium isotope ratio ¹⁷⁶Hf/¹⁷⁷Hf varies from 0.282472–0.282576 (corresponding to εHf(t) = − 0.6 to − 4.1 (t = 320 Ma)), and the oxygen isotope value (δ¹⁸O) varies from 6.0 to 6.8‰. Hf model ages range from 1.2 to 1.4 Ga that are identical with the zircon Nd model ages.

Table 6 LA–MC–ICP–MS Lu–Hf isotope data and SIMS oxygen isotope data of zircons from sample BGK1

Discussion

Comparison of zircon ages analysed with different U–Pb methods (evaporation, SHRIMP, TIMS)

The Pb–Pb zircon evaporation data of biotite and amphibole-bearing granites acquired during this study are systematically older than the ages determined for the same samples with the zircon U–Pb CA–ID–TIMS method (Fig. 9). The offset in ages varies between 1.3 and 4.1% and the error bars do not overlap with the CA–ID–TIMS ages (Fig. 9). This systematical bias of the evaporation data can best be explained by the ca. 1% older age of the secondary standard Temora in comparison to the values published by Black et al. (2004, Supplement 4) and the relatively young Phanerozoic ages of our samples. Generally, Phanerozoic ²⁰⁷Pb/²⁰⁶Pb ages are less reliable in comparison to Precambrian ages, because for ages < 1 Ga, the difference of ²⁰⁷Pb/²⁰⁶Pb ratios strongly decreases towards younger ages.

Fig. 9

Summary of new zircon data acquired with different methods. Mean ²⁰⁷Pb/²⁰⁶Pb ages from zircon evaporation are shown as vertical black bars. One weighted mean ²⁰⁶Pb/²³⁸U SHRIMP age is shown as a vertical orange bar. Mean ²⁰⁶Pb/²³⁸U CA–ID–TIMS ages are shown as vertical dark green bars. Light green bars represent the variation range of all ²⁰⁶Pb/²³⁸U CA–ID–TIMS sample ages for a certain rock group (e.g. group a, b, c, and e)

The SHRIMP age of sample BGK1 is with 323.9 ± 2.1 Ma ca. 3.7% older than the CA–ID–TIMS age (312.2 ± 0.4 Ma, Fig. 9). The accuracy of our SHRIMP age cannot be assessed, because it was determined without the parallel measurement of a secondary standard in 2011. Thus, the uncertainty is only a statistical error, excluding external components. The absolute age resolution of SHRIMP has been estimated as 1–2% (Schaltegger et al., 2015). However, even with an uncertainty of 2%, the SHRIMP age is still significantly older than the CA–ID–TIMS age. Tichomirowa et al. (2019a) compared CA–ID–TIMS with SHRIMP ages from the same samples of Variscan granites from the Western Erzgebirge. While single SHRIMP ages based on ca. 10 measurement have shown similar deviations from CA–ID–TIMS ages, the mean of several samples from the same pluton (corresponding to > 30 data points) resulted in a SHRIMP age identical with CA–ID–TIMS data. Therefore, we suggest that the small number of single data points for BGK1 sample (n = 11) may explain the deviation to the CA–ID–TIMS age.

Two-phase igneous activity on the Lusatian Block

Our new CA–ID–TIMS data of Variscan post-collisional granitoids and volcanics of the Lusatian Block indicate two short episodes (1–2 Myr) of magmatic activity (Fig. 8). The older magmatic event records the intrusion of the Koenigshain granites and the amphibole-bearing granitoids and is restricted to 312–313 Ma. The younger magmatic event (the formation of the Stolpen granite and the volcanic rocks of the Weissig basin) occurred between 298 and 299 Ma.

Despite being shifted towards older ages, our Pb-Pb evaporation data in general support the conclusion that Variscan igneous activity on the Lusatian Block occurred in two separate events (Fig. 7). However, some of the evaporation ages overlap due to large errors and cannot clearly be assigned to the younger or older group, respectively (Figs. 7, 9: BGK 4, BGST11, VR03). Nonetheless, taking both CA-ID-TIMS and evaporation data into consideration, the Koenigshain granite, the amphibole-bearing granite and one of the dykes (VR15) can be assigned to the older group. The younger group is represented by the Stolpen granite, most of the dykes, as well as volcanites from the Weissig basin (Fig. 9).

From the Koenigshain pluton, three main granite varieties are described (porphyritic granite, equigranular leucogranite, and fine-grined monzogranite) that are geochemically different and sometimes show internal intrusion contacts (Eidam and Götze 1991; Hecht et al. 1999). The porphyritic (BGK4) and the equigranular varieties (BGK1) yielded within errors identical ages (312.7 ± 0.4 Ma; 312. 9 ± 0.4 Ma). Similarly, the two main granite facies (porphyritic and equigranular granites) of the Karkonosze pluton gave indistinguishable zircon U–Pb CA–ID–TIMS ages at ca. 312 Ma (Kryza et al. 2014b). According to these data, the Koenigshain pluton and the Karkonosze pluton belong to the same magmatic event. These results are in line with further high-precision CA–ID–TIMS ages that record intrusion activity of some large composite plutons—composed of several magma batches—within ca. 2 Myr (e.g., Kryza et al. 2014b; Ratschbacher et al. 2018; Tichomirowa et al. 2019a).

Both the 298–299 Ma and the 312–313 Ma magmatic events of the Lustian Block were probably accompanied by the intrusion of subvolcanic rocks as dykes. Field relationships, like, e.g. dykes that crosscut the granitic rocks (Kozdrój et al. 2001, Fig. 2b, c), suggest that some of the dykes may be younger than the granitic rocks.

In the Weissig basin, biostratigraphic correlations to other, chronostratigraphically well-defined basins can provide insight into the age of volcanic rocks interlayered with fossil-bearing sediments and help to assess age data. Fossil macroflora of the Weissig basin suggests that the Hutberg Formation is possibly older than the Napoleonstein Formation (Barthel et al. 2010; Reichel 2012; Fig. 3). A stratigraphic relationship of the two formations cannot be determined, because they occur in different parts of the Weissig basin. A correlation of the Hutberg Formation with the Acanthodes horizon of the Goldlauter Formation of the Thuringian Forest (Reichel 2012) indicates an age between 296.9 ± 0.4 and 299.3 ± 0.3 Ma (zircon U–Pb CA–ID–TIMS data of igneous rocks in the Thuringian Forest, Lützner et al. 2021). Our new CA–ID–TIMS age of the Napoleonstein Formation (299.1 ± 0.4 Ma; Fig. 8) well corresponds to this time interval for the Acanthodes horizon and may serve as an argument for the nearly contemporaneous formation of the Hutberg and Napoleonstein units of the Weissig basin. All volcanic rocks of the Weissig basin can thus be assigned to the younger magmatic episode (Fig. 9).

Identification of sources for amphibolite-bearing and biotite-bearing granites

In the literature, the subdivision of the Lusatian Variscan igneous rocks into amphibole-bearing and biotite-bearing granitoids in Germany is consensus (e.g., Eidam et al. 1995; Hammer 1996; Hecht et al. 1999; Hammer et al. 1999; Lisowiec et al. 2013, 2014). While biotite-bearing granites exist in both age groups (312–313 Ma for Koenighshain and 298–299 Ma for Stolpen), all dated amphibole-bearing rocks belong to the older group. In addition to their distinct petrographic composition, amphibole- and biotite-bearing rocks have a different whole-rock geochemistry (Figs. 4, 5, Table 2) as well as Sr- and O-isotope composition (Hammer 1996; Hammer et al. 1999), and zircon morphology (Hammer 1996). Our new Nd-isotopic data of samples BGK1, BGK2, BGK3, and BGK4 from the biotite-bearing granites of Koenigshain overlap with those of the amphibole-bearing granitoid data from Hammer (1996) and do not confirm a difference in Nd-isotopic ratio of both rock types (Hammer 1996). However, calculated magma temperatures are higher for amphibole-bearing granites compared to biotite-bearing granites (Fig. 6). Probably, the biotite- and amphibole-bearing granites have different sources.

Hammer (1996) and Hammer et al. (1999) concluded that tholeiitic rocks are the most suitable and dominant source for amphibole-bearing granitoids. The enrichment of LILE and HFSE elements was explained by these authors as being caused by a subduction-induced metasomatic overprint of a mafic source.

The probable source of biotite-bearing granites is less well defined. Hammer (1996) and Hammer et al. (1999) proposed another dominant, probably metapelitic source for these granites. This assumption is based on a comparison with whole-rock geochemical data of the Variscan Eisgarn granite of Austria, which is thought to have a metapelitic educt (Vellmer and Wedepohl 1994). Hammer (1996) also documented enrichment in LILE and HFSE in the biotite-bearing granites and interpreted this as a supply of elements and volatiles during partial melting instead of source enrichment.

Isotope data on zircon are particularly suitable to determine sources of rocks, because these data are not affected by fractional crystallization or hydrothermal overprint and thus represent the composition of the melt (Chen and Zheng 2017; Tichomirowa et al. 2019b). The zircon Hf data from sample BGK1 (biotite-bearing granite of Koenigshain, Table 6) vary from εHf_(t) = − 4.1 to nearly chondritic ratios (εHf_(t) = − 0.6). In addition, zircon δ¹⁸O values (6.0–6.7‰) are very homogeneous but slightly higher than pure mantle values (5.3 ± 0.3‰, Valley et al. 2005). The calculated δ¹⁸O values of the corresponding melts (8.2—8.9‰, Lackey et al. 2008) agree with the analysed whole-rock δ¹⁸O values of these biotite-bearing granites (Hammer 1996, δ¹⁸O from 8.3 to 9.2‰). Accordingly, the Lusatian biotite-bearing granites can be classified as I-type granites (Chappell and White 2001). A dominant sedimentary source for the biotite-bearing granites (δ¹⁸O typically 10–30‰, Valley et al. 2005) can be excluded, so that a metapelitic source rock—as suggested by Hammer (1996) and Hammer et al. (1999)—is implausible. Intermediate to acidic igneous rocks (δ¹⁸O typically 5–10‰, Valley et al. 2005) are the most likely sources for the biotite-bearing granites from Koenigshain.

A direct comparison of zircon Hf and δ¹⁸O values of sample BGK1 with potential source rocks is limited by the small amount of published Hf- and O-isotope data from the Lusatian Block. A comparison with other Cadomian basement rocks and Variscan igneous rocks of the Saxothuringian Zone reveals the difference of Hf and δ¹⁸O values for sample BGK1 (Fig. 11). Both the Variscan igneous rocks of the North German basin and of the Polish lowlands show distinctly higher δ¹⁸O values and mostly lower but scattering εHf (Pietranik et al. 2013; Słodczyk et al. 2018) indicating distinct sources. The Variscan granites of the Western Erzgebirge show similarly homogeneous εHf but slightly higher δ¹⁸O values compared to zircons from sample BGK1 (Tichomirowa et al. 2019b, Fig. 10). Variscan high-grade metamorphic amphibolite facies and granulite-facies gneisses (metamorphism between 360 and 330 Ma, e.g., Schmädicke et al. 1995; Kröner and Willner 1998; Tichomirowa et al. 2005) have been proposed as the most likely source for these Variscan granites from the Western Erzgebirge based on their zircon data (Hf- and O-isotopes, xenocrystic ages) and their homogeneous model ages (Hf, Nd: 1.3–1.2 Ga, Tichomirowa et la. 2019b). A homogenization of the Hf-isotopic composition in zircon requires an almost complete dissolution of inherited zircons (Farina et al. 2014). However, inherited zircons are abundant in all Cadomian basement rocks of the Bohemian Massif (e.g., Tichomirowa et al. 2001, 2012; Tichomirowa 2002; Friedl et al. 2004; Białek et al. 2014; Zieger et al. 2018). In addition to frequent inherited zircons, all rocks from the Cadomian basement have very heterogeneous zircon Hf- and O-isotope composition as shown for the Saxothuringian basement of the Erzgebirge (Tichomirowa et al. 2018) and of the Schwarzburg Antiform (Linnemann et al. 2014). The almost non-metamorphosed Cadomian basement of the Lusatian Block is composed of the same Cadomian basement as in the Erzgebirge. Therefore, the basement rocks in both the Lusatian Block and the Erzgebirge share many similarities (bulk rock composition, zircon morphology, abundance and ages of inherited zircons, Tichomirowa et al. 2001, 2012), although in the Erzgebirge, the rocks later underwent a high-grade Variscan metamorphism. The Hf and Nd model ages of the biotite-bearing Koenigshain granite (sample BGK1) are very homogeneous (ca. 1.3–1.2 Ga, Fig. 10b), excluding exposed Cadomian basement rocks (that usually have model ages > 1.5 Ga, Tichomirowa et al. 2012) as dominant source rocks (Fig. 10a). Consistent with the homogeneous Hf-isotopic composition, the number of randomly dated inherited grains in the biotite-bearing granites was very small in both evaporation and CA-ID-TIMS dating (< 5%; Table 4, Supplement 1).

Fig. 10

Summary of zircon Hf and O data. Data are from: 1—Pietranik et al. (2013), 2—Słodczyk et al. (2018), 3—Tichomirowa et al. (2019b), 4—this study, 5—Linnemann et al. (2014), 6—Tichomirowa et al. (2018). a) εHf (320 Ma) versus δ¹⁸O values from the same spot locations within zircons from sample BGK1 (Koenigshain pluton) in comparison with data from Variscan gneisses and granites of the Erzgebirge, NE German basin, and Polish lowlands. Abbreviations are af-melting—amphibolite-facies melting, gf-melting—granulite-facies melting according to Tichomirowa et al. (2018) b) Hf-isotope evolution diagram with εHf-data as violin plots (i.e., mirrored probability density plots according to Hintze and Nelson, 1998). CHUR = chondritic uniform reservoir; DM = depleted mantle. c Vertical probability density function of Cadomian sediments of the Schwarzburg antiform (ages between 0.5 and 3.5 Ga, data from Linnemann et al. 2014)

In summary, we agree with the supposed dominant mantle source (tholeiites) for the amphibole-bearing granitoids (Hammer 1996; Hammer et al. 1999). Biotite-bearing granites probably had a different source. Based on our Hf- and O-isotopes in zircon, we can exclude a dominant metapelitic source as suggested by Hammer (1996) and Hammer et al. (1999). The homogeneity of these isotopes also excludes most of the exposed Saxothuringian (Erzgebirge or Lusatian) basement rocks as one of the dominant sources. Intermediate to acidic igneous rocks of unknown ages seem to be the most likely source for the Variscan biotite-bearing granites. The old model ages require a contribution from old (subducted) crust, while their low zircon δ¹⁸O, relatively high Hf- and Nd-isotope compositions indicate an additional input from a mantle source. In comparison to other basement rocks of Saxothuringia (Pietranik et al. 2013; Linnemann et al., 2014; Słodczyk et al. 2018; Tichomirowa et al. 2018, 2019b; Fig. 10b), the Hf- and Nd-isotope compositions are higher and Hf and Nd model ages are younger indicating a larger mantle contribution (Fig. 10b). Further studies are necessary combining whole-rock Nd-, Pb- and Sr- and zircon Hf- and O-isotope analyses on the different types of Variscan igneous rocks to get more information on source rocks. A comparison with published and new isotopic ratios of suspected source rocks, e.g. with the Variscan mantle-derived rocks (e.g. lamprophyric dykes, Abdelfadil et al. 2013; Soder and Romer 2018) could further contribute to this discussion.

The relationship of magmatism and faulting

Igneous rocks along the northeastern boundary of the Lusatian Block resulted in ages of 312–313 Ma (biotite-bearing granites from the Koenigshain pluton, amphibole-bearing granitoids, Karkonosze pluton according to Kryza et al. 2014b). This boundary is the Intra-Lusatian Fault, which is a branch of the ISF (Fig. 2). In contrast, igneous rocks with 298–299 Ma ages (the Stolpen pluton and the volcanic rocks of the Weissig basin) prevail along the western boundary of the Lusatian Block, bound to the Lusatian Thrust Fault and the Stolpen-Klotzsche Fault which are part of the EFZ (Fig. 2b). Since there is petrologic evidence that the Variscan granitic rocks of the Lusatian Block are bound to faults (Thomas and Davidson 2016; Lisowiec et al. 2014; Mierzejewski and Oberc-Dziedzic 1990; Oberc-Dziedzic et al. 2015) and intruded within a post-collisional strike-slip environment (Edel et al. 2018; Mazur et al. 2020), these ages may give constraints for faulting activity on the EFZ and the ISF.

Tectonic activity at the EFZ has been determined previously by dating the Meissen Massif, which is bound to the EFZ and was deformed during its intrusion by dextral strike-slip movements (e.g., Hofmann et al. 2009; Linnemann et al. 2010). The rocks of the Meissen Massif, and thus the time of dextral shear, have been dated with zircon U–Pb SHRIMP (326 ± 6, 330 ± 5 Ma, Nasdala et al. 1999) and with zircon U–Pb LA–ICP–MS (334 ± 3 Ma, Hofmann et al. 2009). Ar–Ar and K–Ar ages on hornblende, biotite and muscovite revealed ages between 323.5 ± 1.0 Ma and 334.7 ± 7.0 Ma (Wenzel et al. 1997; Sharp et al. 1997). Probably, the EFZ was re-activated with different kinematics several times during the post-collisional period of the Variscan orogeny (Scheck et al. 2002; Edel et al. 2018). Our new age data of the Stolpen pluton and the volcanic rocks of the Weissig basin established a magmatic phase at 298–299 Ma that might also be linked to strike-slip deformation along the EFZ. This is within error identical to the age of the Leutewitz ignimbrite of the Meissen volcanic complex (Meissen Massif, 303 ± 3 Ma, Hoffmann et al. 2013). Consequently, two-phase tectono-magmatic activity of the EFZ can be supposed from these data.

Periods of Variscan magmatic activity in Saxothuringia

The crystallization ages of most igneous rocks of Saxothuringia have been determined by different dating methods. Extensive compilations of magmatic intrusion and eruption ages have been given e.g., by Förster and Romer (2010) and by von Seckendorff (2012). Because of its high closure temperatures, zircon U–Pb-dating is best suited to date crystallization of igneous rocks. Currently, only the high precision of CA–ID–TIMS data allows to recognize small (< 1%) differences in crystallization ages. For this reason, we compare our data with published zircon U–Pb CA–ID–TIMS data on further igneous rocks of the Saxothuringian and Tepla-Barrandian zones (Fig. 11).

Fig. 11

adopted from Edel et al. (2018) and Mazur et al. (2020). Age data are from 1—this study, 2– Kryza et al. 2014b, 3—Tichomirowa et al. (2019a), 4—Lützner et al. (2021), 5—Breitkreuz et al. (2021), 6—Opluštil et al. (2016a), 7—Opluštil et al. (2016b). Red colours represent granitic plutons, purple colours represent volcanic rocks. BGK—biotite-bearing granitoids of the Koenigshain pluton, AG –amphibole-bearing granitoids, BGSt biotite-bearing granites of the Stolpen pluton, VRW volcanic rocks of the Weissig basin, KP Karkonosze pluton, ASB Aue-Schwarzenberg pluton, BER Bergen pluton, KIB Kirchberg pluton, EIB Eibenstock pluton, MF Möhrenbach Formation, IF Ilmenau Formation, OF Oberhof Formation, RF Rotterode Formation, TC Tharandt Caldera, NBG Niederbobritzsch pluton, KPB Karkonosze piedmont basin, ISB Intra-Sudetic basin, CF Chotĕvice Formation, BF Broumov Formation, ŽF Žacler Formation, LF Línĕ Formation, TF Týnek Formation, SF Slaný Formation, NM Nýřany Member, URM Upper Radnice Member, LRM Lower Radnice Member. As for the Lusatian Block an older and a younger magmatic phase were defined for the Saxothuringian Zone from these data and are marked with a green and an orange bar, respectively

Compilation of CA–ID–TIMS ages of Variscan magmatic rocks of the Saxothuringian Zone and of the adjoining Tepla-Barrandean volcano-sedimentary basins. Tectonic phases have been

This compilation of U–Pb CA–ID–TIMS data (Fig. 11) suggests the existence of two well-established magmatic episodes following the Variscan orogeny (e.g., Tischendorf and Förster 1990; Pietranik et al. 2013). Contemporaneously with the intrusion of the Koenigshain pluton and the amphibole-bearing granites of the Lusatian Block, the Karkonosze pluton (Kryza et al. 2014b), the volcanic rocks of the Tharandt Caldera of the NE Erzgebirge (Breitkreuz et al. 2021), and volcanic rocks interlayered within the Žacléř Formation of the Intra-Sudetic basin (Opluštil et al. 2016a) were formed. The intrusion of the large granite plutons of the Western Erzgebirge and of the Niederbobritzsch pluton of the NE Erzgebirge occurred slightly earlier, between 322.9 ± 0.4 Ma and 314.1 ± 0.6 Ma (Tichomirowa et al. 2019a; Breitkreuz et al. 2021). The intrusion of the Stolpen pluton and the formation of the volcano-sedimentary Weissig basin occurred simultaneously with and subsequently to the volcanic rocks of the Ilmenau Formation and preceding the deposition of the Oberhof Formation of the Thuringian Forest basin (Lützner et al. 2021). Similar, but slightly younger ages have been reported from the Chotěvice Formation of the Karkonosze piedmont basin and the Broumov Formation of the Intra-Sudetic basin (Opluštil et al. 2016a).

The Varican post-orogenic episode is characterized by polyphase strike-slip-faulting at major shear zones (Arthaud and Matte 1977; Elter et al. 2020). According to structural and microtectonic considerations, dextral faulting prevailed at the main EFZ, but at least one subordinate event of sinistral activity can be recognized at marginal faults (Mattern, 1996). From paleomagnetic, structural and geochronologic data, Edel et al. (2018) derived the timing and kinematics of several tectonic phases that affected the Variscan orogenic belt after the main collisional event. Three of these phases happened during the period of magmatic events of the Saxothuringian Zone hitherto dated with zircon U–Pb CA–ID–TIMS (Fig. 11). During a WNW–ESE extensional episode from 325 to 310 Ma, pre-existing NW striking faults were re-activated dextrally (Edel et al. 2018). These fault movements led to the abundant emplacement of granitic plutons in the Saxothuringian Zone (e.g., Linnemann et al. 2010; Edel et al. 2018) and to the less frequent occurrence of volcanic rocks within and outside of intramontane basins (Fig. 11, the Tharandt Forest Caldera and volcanic rocks in the Intrasudetic basin, Opluštil et al. 2016a). The following episode of NE–SW compression with a duration from 310 to 300 reactivated the NW striking faults sinistrally and was apparently without magmatic activity in the Saxothuringian Zone (Fig. 11, Edel et al. 2018). In contrast to the Saxothuringian data, volcano-sedimentary basins of the Tepla-Barrandean Zone do not show this gap in magmatic activity, but additionally cover the age range between 310 and 300 Ma (Opluštil et al. 2016b, Fig. 11). Finally, NNE–SSW extension from 300 to 260 Ma reactivated NW striking faults dextrally. The magmatic episode that occurred mainly at the beginning of this period, between 300 and 297 Ma (Fig. 11), is widespread in the Saxothuringian Zone. Magmatic activity mainly occurs as volcanic rocks that were deposited together with sediments in the frequent intramontane basins. The Weissig basin and the Thuringian Forest basin are two of these basins that have been biostratigraphically correlated to other Saxothuringian basins (e.g., Schneider et al. 2020). Our zircon U–Pb CA–ID–TIMS data demonstrate that this episode of volcanic activity was also accompanied by the emplacement of plutonic rocks such as the Stolpen pluton.

Conclusion

Our new zircon U–Pb CA–ID–TIMS data suggest two distinct magmatic episodes in the Lusatian Block. These two episodes took place at 312–313 Ma and 298–299 Ma. The biotite-bearing granites of Koenigshain and amphibole-bearing granites constitute the older group. The biotite-bearing granites of Stolpen and the volcanic rocks of the Weissig basin are assigned to the younger group.

New whole-rock Nd and zircon O- and Hf-isotopic data of biotite-bearing granites of the Koenigshain pluton give new insights on the sources of these rocks. The Cadomian basement of the Lusatian Block and of the Erzgebirge can be excluded as sources based on the homogeneity of zircon Hf and whole-rock Nd-isotope data and from corresponding model ages of Variscan granites. Instead, from zircon O-isotopic data, we can infer on an acid or intermediate igneous source of these I-type granites.

The new ages together with literature data suggest two temporally separate magmatic episodes probably related to faulting in the Saxothuringian Zone of the Variscan sorogen. The remarkable simultaneity of the granitoid plutons that are bound to the ISF and its branches on the one hand and of the igneous rocks that are bound to the EFZ on the other hand, suggests that the rise of magma might be linked to distinct faulting events on these shear zones. The correlation of magmatic episodes with tectonic phases of extension rather than compression (Fig. 11, Edel et al. 2018) supports this hypothesis.

Additional zircon U–Pb CA–ID–TIMS data from faulting-related Variscan post-collisional igneous rocks will help to better constrain the upper and lower age limits of these magmatic episodes. These limits would allow to reconstruct the differential tectonic setting at different times and to extend this consideration to other parts of the Variscan orogen. Further zircon and whole-rock Hf-, O-, Nd-, Sr-, Pb-isotope data for different petrological/chemical rock types and age groups will provide additional information for potential sources of these rocks.