# Eclogite subduction wedge intruded by arc-type magma: The earliest record of Variscan arc in the Bohemian Massif

Piérig Deiller, Pavla Štípská, Marc Ulrich, Karel Schulmann, Stephen Collett, Vít Peřestý, Bradley Hacker, Andrew Kylander-Clark, Hubert Whitechurch, Ondrej Lexa, et al.

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Piérig Deiller, Pavla Štípská, Marc Ulrich, Karel Schulmann, Stephen Collett, et al. Eclogite subduction wedge intruded by arc-type magma: The earliest record of Variscan arc in the Bohemian Massif. Gondwana Research, 2021, 99, pp.220-246. 10.1016/j.gr.2021.07.005. hal-03514233

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1 Eclogite subduction wedge intruded by arc-type magma: the earliest record of Variscan arc in the Bohemian Massif

2 the Bohemian Massif

3 Piérig Deiller*, Center for Lithospheric Research, Czech Geological Survey, Klárov 131/3

4 11821, Praha 1, Czech Republic.

5 Pavla Štípská, Center for Lithospheric Research, Czech Geological Survey, Klárov 131/3

6 11821, Praha 1, Czech Republic; Université de Strasbourg, CNRS, IPGS, UMR 7516, 67000 Strasbourg, France

7 Strasbourg, France

8 Marc Ulrich, Université de Strasbourg, CNRS, IPGS, UMR 7516, 67000 Strasbourg, France

9 Karel Schulmann, Center for Lithospheric Research, Czech Geological Survey, Klárov 131/3

10 11821, Praha 1, Czech Republic; Université de Strasbourg, CNRS, IPGS, UMR 7516, 67000 Strasbourg, France

11 Strasbourg, France

12 Stephen Collett, Center for Lithospheric Research, Czech Geological Survey, Klárov 131/3

13 11821, Praha 1, Czech Republic.

14 Vít Peřestý, Institute of Petrology and Structural Geology, Charles University, Prague, Czech Republic

15 Republic

16 Bradley Hacker, Department of Earth Sciences, University of California, Santa Barbara,

17 California, CA 93106 USA.

18 Andrew Kylander-Clark, Department of Earth Sciences, University of California, Santa Barbara, California, CA 93106 USA.

19 Barbara, California, CA 93106 USA.

20 Hubert Whitechurch, CNRS UMR7516, Université de Strasbourg, 1 rue Blessig, F-67084, France.

21 France.
22 Ondrej Lexa, Institute of Petrology and Structural Geology, Charles University, Prague,
23 Czech Republic.

24 Eric Pelt, Laboratoire d'Hydrologie et de Géochimie de Strasbourg (LHyGeS), Université de
25 Strasbourg, CNRS, ENGEES, 1 rue Blessig, 67000 Strasbourg, France.

26 Jitka Míková, Czech Geological Survey, Geologická 6, 15200 Praha 5 Czech Republic.

27 *Corresponding author: pierigd@yahoo.fr

28 ABSTRACT

29 Previously unrecognized mafic and felsic plutonic rocks, formerly interpreted as an anatectic
30 part of the amphibolite-eclogite Mariánské Lázně Complex (MLC, Bohemian Massif) were
31 examined together with a possible upper-crustal equivalent represented by the Čistá pluton
32 intruding continental crust farther east. These plutonic rocks were studied by whole-rock
33 geochemistry (major and trace elements, Sr-Nd isotopes) and zircon U-Pb geochronology.
34 The MLC magmatic rocks range from pyroxene-amphibole gabbro to trondhjemite and
35 oligoclasite. The Čistá pluton consists of porphyritic granitoids. The calc-alkaline nature of
36 these rocks, relative enrichment of fluid-mobile elements (including large-ion lithophile
37 elements), strong fractionation of light rare earth elements over heavy rare earth elements and
38 depletion of high field strength elements are evidence for active margin magmatism. U-Pb
39 zircon dating indicates that the magmatism was Mid-Devonian in both the MLC (ca. 385 Ma)
40 and Čistá pluton (ca. 390 up to ca. 365 Ma). Both magmatic units show inherited cores
41 probably linked to recycling of crustal material. Sr-Nd isotopic compositions indicate that the
42 gabbro formed from a depleted MORB mantle source composition with a small contribution
43 of a subduction and crustal component. A three-stage geodynamic model is proposed: 1) ca.
44 390–380 Ma recycling of a high-pressure accretionary wedge and migration of arc melts into
45 more distal parts of the upper continental plate, 2) ca. 380–370 Ma compression and
thickening of the accretionary wedge-arc edifice, and 3) ca. 370–365 Ma development of a giant east-dipping detachment that reworked the whole edifice in the west and also involved emplacement of the supracrustal Čistá pluton along antithetic west-dipping shear zone in the east. The studied magmatic rocks record a Mid-Devonian magmatic arc related to the formation of a subduction wedge stacked beneath the upper plate that was followed by formation of a second and independent Carboniferous continental arc further east.

KEYS WORDS: continental arc magmas; Bohemian Massif; whole-rock geochemistry; Sr, Nd isotopes; U–Pb zircon geochronology

# 1. INTRODUCTION

Subduction of oceanic lithosphere leads to eclogitization of the former oceanic crust, the associated dehydration reactions lead to flux melting of the overlying mantle wedge, resulting in predominantly calc-alkaline magmatism in the upper plate (e.g. Sisson and Grove, 1993; Grove et al., 2012). Therefore, identification of high-pressure (HP) metamorphic rocks associated with calc-alkaline magmatism and accurate constraint of their relative timings are important criteria in the understanding of paleo-subduction zones (e.g. Stern, 2002).

The Paleozoic European Variscan belt resulted from an Early – Mid Devonian subduction of a Cambrian-Ordovician Rheic type oceanic plate followed by Late Devonian - Early Carboniferous collision of the Gondwanan margin with a large ribbon continent in the north (e.g. Stampfli et al., 2013; Edel et al., 2018). The early subduction event was characterized by ca. 400–390 Ma imbrications of HP nappes with intraoceanic arc assemblages well preserved in ophiolitic sequences in Iberia (Sanchéz Martínez et al., 2007), southern Armorica (Cartier and Faure, 2004) and the Münchberg massif (Koglin et al., 2018) (Fig. 1a). In contrast, the later subduction-collisional event is characterized by the development of granulite-migmatite domes, which enclose boudins of ultra-HP and ultra-high-temperature (UHT) mantle
fragments, most spectacularly exposed within the Bohemian Massif (e.g. Medaris *et al.*, 1995; Schulmann *et al.*, 2014; Konopásek *et al.*, 2019; Ackerman *et al.*, 2020; Maierová *et al.*, 2021) but, also within the Black Forest, Vosges and French Massif Central (e.g. Lardeaux *et al.*, 2014). Only the latter event is associated with regionally extensive (ca. 360–335 Ma) arc magmatism firmly established on previously stacked continental crust units (e.g. Lardeaux *et al.*, 2014; Žák *et al.*, 2014; Fig. 1a). In the French Massif Central, it is represented by tonalitic intrusions or andesitic volcanic sequences, while in the Bohemian Massif, it is characterized by a 9000 km² large composite Central Bohemian Plutonic Complex (CBPC, Fig. 1a, b, c).

The lack of a magmatic arc associated with the Mid-Devonian subduction represents an important missing link between the earlier and later stages of the Variscan Orogeny, which therefore hampers our understanding of its tectono-thermal evolution.

The Bohemian Massif represents a valuable object in bridging this gap as both the early subduction and later collisional events are well constrained from petrological-chronological and structural data (e.g. Schulmann *et al.*, 2009; Collett *et al.*, 2018; Martínez Catalan *et al.*, 2020). The target area of this study is the Mariánské Lázně Complex (MLC, Figs 1b, 2a), part of a metamorphic belt (also including the Münchberg and Góry Sowie Massifs) of Early to Mid-Devonian subduction related eclogite- to granulite-facies rocks that were interpreted as remnants of an oceanic suture associated with the subduction of the Saxothuringian Ocean beneath the upper crustal Teplá-Barrandian Unit (TBU) forming the upper continental plate (Franke *et al.*, 2017; Martínez Catalán *et al.*, 2020). Further Devonian age subduction-derived remnants of mantellic and crustal origin were also recently described in the lower crustal Moldanubian domain (Ackerman *et al.*, 2020).

Based on mid-Devonian metamorphic ages of the eclogite, the MLC rock association was interpreted to record the main oceanic subduction-accretion stage preceding a paroxysmal Early Carboniferous collisional event (Schulmann *et al.*, 2009, 2014). It was also shown that
the MLC is formed by two distinct units: 1) an upper gabbro-amphibolite unit of Late Cambrian age, which might be interpreted as the deepest part of the TBU (Jašarová et al., 2016); and 2) a lower eclogite-amphibolite unit that belongs to the subduction-accretionary wedge (Collett et al., 2018). Both units suffered the same granulite-amphibolite facies overprint and therefore represent a key area where the common evolution of subducted and accreted oceanic rocks and upper plate continental crust can be studied. The rock association that comprises the focus of this study consists of mafic gabbro to quartz-gabbro, partly metamorphosed to amphibolite, and felsic magmatic rocks, sometimes deformed into gneiss, spatially associated with eclogite in the eclogite-amphibolite unit. Both the gabbroic and felsic rocks were previously interpreted as products of melting of the eclogite-amphibolite unit—i.e. as metamorphic rocks produced by melting of a broadly unknown protolith or as a result of HT decompressional melting of eclogitic oceanic crust (Timmermann et al., 2004).

Based on field and petrographic study, combined with whole rock geochemical, Sr-Nd isotopic, and U-Pb zircon geochronological analyses, we propose that the studied rocks are in fact of magmatic origin and that they intruded the partially exhumed Mid-Devonian eclogite dominated subduction wedge (Collett et al., 2018). Because of the structural and metamorphic overprint of the MLC eclogite and amphibolite, which can be correlated to that of hanging-wall schists of the TBU (Peřestý et al., 2017, 2020), we can also preliminarily constrain the emplacement age of these magmatic rocks as Devonian. If true, these magmatic rocks represent the first record of an igneous event recorded in the European Variscan belt that was intimately related to the exhumation of a mantle wedge and its emplacement beneath the continental upper plate. To develop this hypothesis we examined also the Čistá pluton, a Late Devonian intrusion emplaced into the TBU upper crustal rocks (Venera et al., 2000) and recently interpreted as one of the earliest intrusions of the CBPC continental magmatic arc further east (Žák et al., 2011). Here, we propose an alternative link between the gabbro of the
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MLC and the granodiorite of the Čistá pluton, which may question the Late Devonian – Early

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Carboniferous continuity of arc magmatism in the Central Bohemian Plutonic Complex.

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Instead, it is proposed here that two temporally and spatially distinct magmatic arcs may have
formed at the western and eastern parts of the TBU, respectively.

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In this frame, we propose a new petrogenetic and geodynamic model in which the
Early Devonian gabbro of the MLC originated from the melting of depleted mantle and
intruded a stacked eclogite-bearing subduction wedge of Early Cambrian protolith age. The
wedge additionally became partially molten, which is indicated by the presence of Cambrian
inherited zircon grains in the majority of the studied samples (only one sample does not show
any inheritance). In our model, the subduction wedge, intruded by the magmatic arc, was
underplated beneath the continental upper plate and part of the magma then rose up in the
form of a supracrustal granodiorite stock forming the Čistá pluton. All these data are
integrated with the existing geodynamic models to propose two spatially and temporally
distinct magmatic arcs, linking the final stages of Early-Mid Devonian building of the
subduction wedge, oceanic subduction and the onset of Carboniferous continental collision in
the European Variscides.

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## 2. GEOLOGICAL SETTING

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### 2.1 The Teplá-Barrandian Unit

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The TBU occurs in the central part of the Bohemian Massif (Fig. 1a, b) and together with the lower-crustal Moldanubian domain it formed the upper plate during the Variscan Orogeny (Schulmann *et al.*, 2009). The basement of the TBU represents a near-complete section across a Late Ediacaran to Cambrian active margin (e.g. Sláma *et al.*, 2008; Hajná *et al.*, 2011, 2014; Ackerman *et al.*, 2019). From SE to NW this sequence is composed of a ca. 620–560 Ma volcanic arc overlain by intra-arc basins (Davle volcanic complex and Štěchovice Group,
respectively), an Ediacaran to Cambrian accretionary wedge (Blovice Accretionary Complex),
and a Cambrian composite metabasite complex (Mariánské Lázně Complex, MLC).

The basement is unconformably overlain by a Mid-to-Late Cambrian continental and
marine sedimentary sequence and by the Prague Basin, where the earliest Ordovician-to-mid-
Devonian marine sedimentation interacted with rift-related submarine volcanism (Tasáryová
et al., 2018). The Prague Basin consists of Ordovician cold-water passive margin siliciclastic
succession, Silurian tuffitic black shales, Devonian warm-water carbonate-dominated reef and
platform successions terminated by the Middle Devonian-Givetian flyshoid siliciclastic
Srbsko formation (see Vacek and Žák, 2019 for review). The Srbsko formation contains
detrital zircon grains dated at 411–391 Ma (Strnad and Hladil, 2001; Strnad and Mihaljevič,
2005). According to these authors, the chemistry and detrital zircon pattern of the Srbsko
formation indicate magmatic activity related to a subduction-related island-arc environment;
although this arc was thus far unidentified in the preserved geological record.

The southeastern boundary between the TBU and the Moldanubian domain is
composed of the calc-alkaline to highly potassic Central Bohemian Plutonic Complex
(CBPC) interpreted as a large internally compositionally zoned continental magmatic arc
(Janoušek et al., 1995; Holub et al., 1997a; Janoušek et al., 2000; Žák et al., 2014). There, the
Staré Sedlo and Mirotice gneiss complexes (calc-alkaline granitoid protolith of ca. 380–365
Ma) are thought to represent the oldest Paleozoic intrusive events in the area prior to
emplacement of the CBPC itself (Košler et al., 1993a,b, 1995). The earliest undeformed
intrusions of the CBPC sensu stricto are the Sázava suite in the north and NW (Fig. 1b)
emplaced ca. 354 Ma during a regional transpression regime (Janoušek and Gerdes, 2003;
Žák et al., 2005a), followed by the high K - calc-alkaline Blatná suite in the central part dated
at ca. 346 Ma and terminated by ca. 343–337 Ma ultrapotassic intrusions in the SE (Fig. 1b;
Holub et al., 1997a, b; Žák et al., 2005b). The structure of the high K – calc-alkaline suite
reflects a transition from NW directed steep magmatic foliation onto an extensional gently dipping magmatic flow regime, while the ultrapotassic intrusions show solid state extensional normal shearing related deformation fabrics (Žák et al., 2005b). Altogether, the temporal and tectonic evolution of the CBPC continental arc is interpreted in terms of an early intrusion of primitive calc-alkaline magma at ca. 354 Ma during general transpressional regime followed by ca. 347–336 Ma high K - calc-alkaline to ultrapotassic magmatism related to large scale syn-extensional exhumation of the core of the Bohemian Massif (Žák et al., 2005b, Janoušek et al., 2010). This tectonic evolution is terminated by intrusions of post-tectonic anatectic granitoids (ca. 342–326 Ma) in the north eastern part of the CBPC (Trubač et al., 2009) and in the western part of the Bohemian Massif (e.g. Trubač et al., 2019).

In contrast, the study area located in the northwestern part of the TBU, is dominated by the Blovice Accretionary Complex (Hajná et al., 2014), locally metamorphosed (to form the Teplá Crystalline Complex, TCC), and composed of phyllite, metagreywacke, schist and paragneiss (Kastl and Tonika, 1984; Dörr et al., 1998). Detrital zircon geochronological study from a banded mica schist of the TCC indicates maximum depositional age of 533±8 Ma (Žák et al., 2020). Late Cambrian granitoids (ca. 528–504 Ma; Dörr et al., 1998; Venera et al., 2000), metapegmatite (ca. 487–470 Ma; Glodny et al., 1998) and the Late Devonian Čistá and Štěnovice granitoids (Fig. 1b; ca. 375–373 Ma; Venera et al., 2000; Žák et al., 2011) intruded the metasedimentary pile of the TCC. In the TCC, three deformation events were recorded (Peřestý et al., 2017, 2020). The first event was associated with an early Cambro-Ordovician HT–LP metamorphism. The second deformation event was a Devonian shortening resulting in westward-increasing Barrovian metamorphism and the development of a vertical schistosity. The third deformation event corresponded to Late Devonian (ca. 370 Ma) extensional shearing (Peřestý et al., 2017), associated with a HT–LP metamorphic overprint that similarly
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affected the underlying MLC (Fig. 2a, b). The Devonian Barrovian and HT-LP events

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affected also the majority of the MLC.

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## 2.2 The Mariánské Lázně Complex

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The MLC itself (Fig. 2a, b) overlies a body of serpentinite, derived from a spinel harzburgite (Medaris *et al.*, 2011), which separates the MLC from the low-grade schist and amphibolite of the Kladská Unit (Kachlík, 1993) and lower plate Saxothuringian domain. The MLC is dominated by amphibolite and migmatitic amphibolite that host bodies of variably retrogressed eclogite mainly in the north-western and central parts (eclogite-amphibolite belt in Fig. 2a, b), bodies of coronitic Cambro-Ordovician (C–O) metagabbro with partly preserved magmatic textures (Štědrá, 2002) in the south-east (gabbro-amphibolite belt in Fig.

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subordinate high pressure (HP) mafic granulite, paragneiss and felsic orthogneiss are also present (O’Brien, 1994, 1997; Jelínek *et al.*, 1997; Štědrá, 2001; Timmermann *et al.*, 2004; Faryad, 2012). The C–O metagabbro share common geochemical, geochronological, and petrological features with similar bodies in the overlying Teplá-Barrandian Unit (Timmermann *et al.*, 2004; Jašarová *et al.*, 2016).

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### 2.2.1 Metamorphism

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Eclogite-facies metamorphism in the MLC is preserved in relict Grt–Omph–Ph–Ky–Rt assemblages with estimated peak conditions of ca. 25 kbar and ca. 650–750° C (Collett *et al.*, 2018). The eclogite is to a large extent overprinted by granulite-facies assemblage at ca. 14 kbar and ca. 800° C (e.g. diopside–plagioclase symplectite replacing omphacite, sapphirine and/or corundum surrounding kyanite, clinopyroxene–orthopyroxene–plagioclase around quartz; Collett *et al.*, 2018), and then partially or completely retrogressed to amphibolite facies assemblages (Amp–Pl–Grt–Ttn–Qtz; Štědrá, 2001; Timmermann *et al.*, 2004; Collett *et al.*, 2018). Furthermore, local relict garnet cores suggest an older phase of metamorphism (Faryad, 2012). The C–O metagabbro does not show evidence for eclogite-facies

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metamorphism. However, *P*–*T* estimates for the garnet-bearing coronas (700 ± 50 °C; 13.5 ± 1.5 kbar; Jašarová *et al.*, 2016) are broadly consistent with the amphibolite to granulite-facies overprint of the eclogite. The amphibolite is considered to represent the thoroughly retrogressed equivalents of the associated eclogite and metagabbro (Kastl and Tonika, 1984; Jelínek *et al.*, 1997; Štědrá, 2001; Timmermann *et al.*, 2004). The magmatic rocks of the present study have been poorly documented in previous work and they were assumed to be migmatitic amphibolite, presumably produced during the HT overprint of the eclogite. The only “igneous” sample investigated in detail was described as a felsic leucosome in the neck of an eclogite boudin (sample HT-99-1 in Timmermann *et al.*, 2004).

## 2.2.2 Geochronology

Geochronological studies distinguish two major protolith groups. Zircon yielded late Ediacaran-Early Cambrian (ca. 550–530 Ma) protolith ages for both the eclogite and amphibolite (Timmermann *et al.*, 2004) and Cambro-Ordovician (ca. 500 Ma) protolith ages for the C–O metagabbro (Timmermann *et al.*, 2006). Zircon lower intercept, monazite, titanite and rutile dates for each of the different lithologies are Variscan (ca. 385–360 Ma) and were interpreted as the age of metamorphism (Timmermann *et al.*, 2004). The only U-Pb zircon metamorphic ages for two eclogites (377±27 Ma, 382±3 Ma) and one amphibolite (373±10 Ma) are presented in Timmermann *et al.* (2004), who inferred that this age reflected the timing of Variscan exhumation and associated decompression melting rather than eclogitization. A felsic leucosome related to the eclogite yielded a titanite date of 378 ± 4 Ma, which was considered to be the magmatic crystallization age of the leucosome and a minimum age for the granulite-facies metamorphism (Timmermann *et al.*, 2004). Lu–Hf garnet geochronology (Collett *et al.*, 2018) indicates a Middle Devonian (390.3 ± 1.4 Ma) age for the eclogite-facies event, whereas a younger (376.1 ± 3.0 Ma) Sm–Nd age is interpreted to date cooling. Similar Sm-Nd garnet rim ages (ca. 377–367 Ma) are additionally presented in
243 Beard *et al.* (1995), while separated garnet cores yielded older (ca. 420–409 Ma) ages. Mid-
244 to Late-Devonian dates have also been recorded by $^{40}$Ar/$^{39}$Ar and K–Ar work on hornblende
245 (ca. 372 Ma; 379 Ma; 380 Ma), biotite (ca. 359 Ma) and muscovite (ca. 366 Ma) (Kreuzer *et
246 al.*, 1992; Zulauf, 1997; Dallmeyer and Urban, 1998; Bowes *et al.*, 2002; Medaris *et al.*,
247 2011).

248 2.2.3 Geochemistry

249 Previous geochemical studies were mostly focused on the eclogite, the C–O metagabbro and
250 the associated amphibolite. The geochemical signature of the eclogite and the amphibolite
251 mainly corresponds to mid oceanic ridge basalts (MORB) with depletion in light rare earth
252 elements (LREE); however, some LREE-enriched, E-MORB-like patterns were also reported
253 (Kastl and Tonika, 1984; Beard et al., 1995; Timmermann et al., 2004). This enrichment,
254 together with mainly positive εNd values (+5.4 to -0.8) may reflect either subduction, or
255 ocean-island basalt components in their source (Beard et al., 1995; Timmermann et al., 2004).
256 Crowley et al. (2002) identified three contrasting geochemical groups within the metabasite,
257 alkali, tholeiitic, and low-Ti tholeiitic and suggested that the metabasites formed by the
258 interaction between oceanic lithosphere generated at a spreading center and an upwelling
259 mantle plume. One sample of intermediate granulite exhibits adakite-like trace-element
260 compositions (Timmermann et al., 2004). The sample of felsic leucosome (HT-99-1 in
261 Timmermann et al., 2004) shows close isotopic affinity to this granulite and some of the
262 amphibolite and was interpreted as originating from eclogite-amphibolite melting. The C–O
263 metagabbro show E-MORB-like REE patterns compatible with an intracontinental rift setting
264 (Štědrá, 2002; Jašarová et al., 2016).

265 2.2.4 Structure

266 The dominant SE-dipping structures within the MLC units were used to interpret that
267 serpentinite lies structurally beneath a central eclogite-amphibolite belt and a hanging-wall of
268 metagabbro-amphibolite (Fig. 2a). Hrouda et al. (2014) measured anisotropy of magnetic
269 susceptibility (AMS) fabrics in mafic and ultramafic rocks and showed that the serpentinite,
270 massive eclogite, gabbro and amphibolite reveal N-S to NE-SW or NW-SE trending steep to
271 sub-horizontal magnetic foliations. These fabrics were interpreted to be entirely the result of
272 deformation, acquired by all the investigated rock types, and to have formed during retrograde
273 metamorphism. Collett et al. (2018) and Peřestý et al. (2017; 2020) distinguished a structural
274 succession from steep to progressively shallower dipping fabrics schematically depicted in
275 Fig. 2b. The steep fabric is defined by NNE-SSW and NW-SE striking eclogite-facies garnet-
276 poor and garnet-rich layers (SEcl in Collett et al., 2018). The migmatitic fabric in amphibolite
277 surrounding the eclogite boudins (S2) is characterized by alternating felsic and mafic bands,
278 with and without garnet. This fabric is steeply to moderately dipping to the east. The S2 fabric
279 is folded and transposed into an east-dipping, generally shallower S3 fabric, interpreted as
280 extensional (Collett et al., 2018; Peřestý et al., 2017, 2020).

281 **2.3 The Čistá pluton**

282 The Čistá pluton (Fig. 2c) intrudes the Blovice Accretionary Complex approximately 40 km
283 east of the contact with the MLC (Fig. 1b). This coarse-grained granodiorite pluton (8.5 x 9.6
284 km) is elongated NE–SW and has generally steep contact with its wall rocks (Žák et al.,
285 2011). To the east, it intruded the Cambrian Tis granite, where the contact coincides partly
286 with a NE-SW striking ductile shear zone (Fig. 2d), in which both plutons are deformed by a
287 sinistral transtensional regime (Venera et al., 2000). The Tis granite was deformed in a solid
288 state at high temperature, whilst the Čistá pluton developed both magmatic and high
temperature solid-state deformation fabrics during cooling. To the west, it intruded
289 Neoproterozoic rocks, producing a narrow contact aureole. The Čistá pluton is dated at 373.1
290 ± 1.1 Ma (Pb–Pb zircon vapo-ation; Váná et al., 2000) and is contemporaneous with the
291 Štěnovice granodiorite pluton (375 ± 2 Ma, U–Pb zircon; Žák et al., 2011). The whole-rock
and trace-element geochemical characteristics of each pluton are consistent with a subduction-
related calc-alkaline magmatism, generated in a continental margin setting (Žák *et al.*, 2011).
These intrusions were interpreted to represent the initial Late Devonian stage of voluminous
early Carboniferous arc-related plutonism of the Central Bohemian Plutonic Complex further
east (Žák *et al.*, 2011).

3. FIELD RELATIONS OF MAGMATIC ROCKS OF THE MLC AND THE ČISTÁ
PLUTON

Within migmatitic amphibolite located in the eclogite-amphibolite belt of the MLC, some of
the rocks display typical magmatic structures and their position and relation to structure are
portrayed in profile in Fig. 2b (a total of 25 samples were collected from such rocks). Mainly
composed of ferro-magnesian minerals, plagioclase and minor quartz, these magmatic rocks
exhibit a coarse-grained texture, with a random distribution of the mineral phases, euhedral
crystals, and alternating leucocratic and melanocratic types showing a progressive transition
in petrographic continuity (Fig. 3a). They are called here gabbro and generally occupy
numerous m-scale outcrops. Some rocks, occurring mostly as dm- to m-scale large dykes
(Fig. 3b, c) or pods are trondhjemites (tonalite, composed of fine to coarse grained
plagioclase, subordinate quartz without or with rare amphibole), tonalite (ML33B2) or rocks
dominated by oligoclase feldspar and therefore referred to as oligoclasite (samples ML28).
They commonly show sharp boundaries to the surrounding gabbro as seen in the lower part of
Fig. 3b. In some places they are in petrographic continuity with the coarse-grained host-rock
gabbro, as in the upper part of Fig. 3b. The gabbro and felsic rocks occur also at the contact
with retrogressed eclogite; their sharp contacts indicate they are intrusive with respect to the
metamorphic rocks (Fig. 3c, e, f). Most of these intrusive contacts are in continuity with the
steep S2 fabric preserved in eclogite and to a certain degree in amphibolite (Fig. 3a, b, c, e).
Locally, the unfoliated gabbro, diorite and trondhjemite show transitional fabric foliation (Fig.
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3d) evolving to a solid-state amphibolite-facies foliation. The felsic rocks are in some places foliated and these deformation fabrics are gently east dipping, parallel to the regional S3 foliation.

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In the Čistá pluton, magmatic rocks are mostly coarse-grained granitic rocks (granodiorite to granite) with no macroscopically observable magmatic foliation (6 samples were studied). The core of the Čistá pluton is concentric, defined by steep, margin-parallel, weak magmatic fabric associated with variably oriented hornblende, whereas the outer rim is strongly foliated and more mafic (Venera et al., 2000; Žák et al., 2011). These observations are compatible with anisotropy of magnetic susceptibility AMS study (Chlupáčová et al., 1975), which also indicates an overall concentric structure and increasing magnetic anisotropy toward the pluton margin. The coarse granodiorite is crosscut by up to kilometer long dykes of a fine grained melanocratic rhyodacite (Fig. 2c).

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## 4. ANALYTICAL PROCEDURES

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Selected samples have been studied for mineral chemistry, whole rock geochemistry, Sr and Nd whole rock isotopes, and U-Pb zircon geochronology. Brief descriptions of these analytical procedures are given below and full documentation including analyzed standards and external errors are provided in the Supplementary Data, Analytical procedures file.

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### 4.1 Mineral chemistry

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The mineral chemistry of gabbroic samples ML31A, ML33C, ML34B, ML27A4, ML34C, ML46A, ML33A, ML44-1, ML44-2 and felsic samples ML33B2, ML28D were analyzed by scanning electron microscope and electron microprobe. Mineral analyses of major elements were obtained at the Institute of Petrology and Structural Geology, Charles University in Prague using a scanning electron microscope (SEM) TESCAN Vega with an X-Max 50 EDS detector from Oxford Instruments and an electron probe microanalyser (EPMA) with a field
emission gun (FEG) electron source JXA-8530F by JEOL, equipped with 5 wave dispersive spectrometers (WDS). Amphibole compositions were recalculated using the spreadsheet of Locock (2014) and amphibole names have been assigned according to Hawthorne *et al.* (2012). Representative mineral analyses are given in Supplementary Data, Table S1.

## 4.2 Whole-rock geochemistry

Twenty-five samples from the MLC, including basic to acidic rocks, and six additional acidic rocks from the Čistá pluton were analysed for major and trace elements using ICP-AES and ICP-MS, respectively. Measurements were performed at the Bureau Veritas laboratory (Vancouver, Canada) and at the Laboratoire d'Hydrologie et de Geochimie de la Surface (LHyGeS, University of Strasbourg, France), and are presented in Supplementary Data, Table S2. CIPW norm (calculated with hornblende) is given in Supplementary Data, Table S3.

Sr and Nd isotopic compositions of 11 samples from MLC and 3 samples from the Čistá pluton were performed at LHyGeS. These isotopic compositions were determined following the analytical procedures of Geagea *et al.* (2007), Pin and Zalduegui (1997) for Nd isotopes, and Deniel and Pin (2001) for Sr isotopes.

## 4.3 U-Pb zircon geochronology

Zircon grains from five samples from the MLC and two samples from the Čistá pluton were separated for U-Pb dating. Ages were determined by LASS (laser ablation-split-stream inductively coupled plasma-mass spectrometry) at the Department of Earth Science of the University of Santa Barbara (California, USA) following the analytical procedure described by Kylander-Clark *et al.* (2013) and by laser ablation ICP-MS at the Czech Geological Survey (data in Supplementary Data, Table S4).
364

# 5. PETROGRAPHY AND MINERAL CHEMISTRY OF MAGMATIC ROCKS OF THE MLC AND PETROGRAPHY OF THE ČISTÁ PLUTON

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## 5.1 The MLC samples

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### 5.1.1 Mafic rocks

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Macroscopically, the majority of the MLC mafic samples have a magmatic aspect with granular to microgranular texture, that is in places porphyritic. The mafic rocks are clinopyroxene–amphibole gabbro (Fig. 4a), amphibole gabbro (Fig. 4c, e) to quartz amphibole gabbro (Fig. 4d) with 50–80% of amphibole and 20–50% of plagioclase (Fig. 4a–d), sometimes with quartz (up to 20%). They are characterized by mm- to cm-scale grains (Fig. 4a, c, d, e) with random to weak preferred orientation of amphibole. Some samples are fine-grained (≤1 mm) (Fig. 4b) and have a strong preferred orientation of amphibole. Garnet is present in one sample (ML34B) taken at the contact with an eclogite lens. Accessory minerals are rutile, titanite, biotite, ilmenite, apatite, zircon and epidote (zoisite). In places, the rocks also contain chlorite and altered plagioclase.

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Microscopically, the gabbro has an equigranular (Fig. 5a, b) or porphyritic texture (Fig. 5c). In places, the grains have straight boundaries and contacts in triple points in the amphibole-plagioclase matrix, interpreted as a result of static re-equilibration by grain boundary reduction and homogenization of the grain size (Fig. 5b). The amphibole is mostly tschermakite to magnesio-hornblende (Fig. 5e), with increasing Si and Mg content from the core to the rim (Fig. 5a, e). Plagioclase is either not zoned or shows increasing Na from core to rim (Fig. 5d, f). The chemistry of amphibole and plagioclase is correlated with whole rock composition (Supplementary Data, Fig. S1), suggesting a magmatic fractional crystallization evolution.
387 **5.1.2 Felsic rocks**

388 The sampled felsic rocks have considerable variation in terms of mineral assemblage and in
389 degree of deformation. Samples from the ML30 locality have a foliation parallel to the
390 regional S3 fabric, and are called here gneiss (ML30, ML30E and ML30H). They are
391 composed of plagioclase (15-40%) and quartz (30-60%) with accessory zircon, biotite,
392 muscovite, garnet, apatite and iron oxides. In contrast, the other felsic samples are without
393 macroscopically visible deformation fabric, they have granular to microgranular texture, and
394 are mainly composed of plagioclase and quartz (Fig. 4f), and typically have tonalitic (15-30%
395 of quartz, 60-75% of plagioclase, 5-10% of amphibole) or trondhjemitic (Fig. 4f, 20-60 of
396 quartz, 30-75% of plagioclase) assemblages. Exceptional are the ML28 samples which lack
397 quartz (Fig. 4g, h) and are composed of more than 90% plagioclase mostly of oligoclase
398 composition, but, with internal patches of anorthite composition. Accessory minerals in all
399 samples include zircon, biotite, apatite and iron oxides. Samples contain also few chlorite and
400 epidote, which are interpreted as reflecting hydration related to metamorphic overprint.

401 **5.2 The Čistá pluton**

402 Samples from the Čistá pluton are coarse grained granodiorite and granite (Fig. 4i, j) with a
403 granular texture, except one sample (CI4D) which is an aphanitic rhyodacite dyke. The rocks
404 are composed of biotite, K-feldspar, plagioclase and quartz and accessory zircon, apatite,
405 titanite, amphibole, muscovite and chlorite.

406 **6. WHOLE ROCK CHEMISTRY**

407 The rocks are assigned to 3 major groups: (1) MLC mafic rocks comprising the coarse to
408 medium grained amphibole-bearing rocks ranging mostly from gabbro to quartz-gabbro, (2)
409 MLC felsic rocks comprising medium to fine-grained trondhjemite-oligoclasite-gneiss dykes
410 and pods, and (3) the coarse grained granite to granodiorite rocks of the Čistá pluton. The
411 MFW diagram (Ohta and Arai, 2007; calculation based on eight oxides, Supplementary Data,
412 Fig. S2) is used to indicate if the samples underwent chemical weathering. It shows that the
413 samples fit in the igneous rock trend, with the exception of samples ML30, ML30E and
414 ML30H (the felsic gneiss), which are gently affected by chemical alteration, most probably
415 related to their metamorphic overprint. As the alteration is negligible in the magmatic
416 samples, to very low in case of the gneiss, all samples were kept for the study. However, the
417 gneiss samples were just studied for chemical description but not considered in any modelling
418 study.

419 **6.1 Major element composition**

420 The MLC mafic rocks range from basic to slightly acidic (SiO2=47.84–63.04 wt %), with
421 total alkalis (Na2O+K2O) between 2.59 and 5.20 wt % (Fig. 6a). The samples display
422 homogeneous Al2O3 (14.2–18.8 wt %), variable MgO concentrations (2.5–8.8 wt %) and Mg#
423 (Mg#=100*MgO/(MgO+FeO<sub>T</sub>) on a molar basis) varying from 40 to 78. The MLC felsic
424 rocks range from intermediate to highly acidic (SiO2=58.7–84.0 wt %), with large variation in
425 total alkalis (Na2O+K2O=2.6–9.2 wt %). The MLC felsic rocks are characterized by large
426 variations in Al2O3 (9.2–24.3 wt %) and low MgO (0.03–3.0 wt %). Mg# varies from 44 to
427 59. Samples from the Čistá pluton are all acidic (SiO2=68.1–71.5 wt %), with total alkalis
428 between 7.4 and 8.5 wt % (Fig. 6a). The MLC mafic and felsic rocks define slightly
429 overlapping groups on TAS and Al2O3-FeO-MgO (AFM) diagrams (Fig. 6). All the samples
430 plot in the subalkaline-tholeiitic field of the TAS diagram, except for the samples from
431 locality ML28 (Fig. 6a), and all correspond to the calc-alkaline series in the AFM diagram
432 (Fig. 6b).

433 Figure 7 shows all major element concentrations vs. SiO2 content in the MLC and
434 Čistá rocks (i.e. Harker diagrams), along with Mg# vs. SiO2. Samples of MLC eclogite and
C–O gabbro are additionally plotted for comparison. In these Harker diagrams, MLC mafic and felsic rocks define slightly overlapping groups on each plot. MgO, FeO<sub>T</sub>, CaO and TiO<sub>2</sub> concentrations negatively correlate with SiO<sub>2</sub>. Conversely, Na<sub>2</sub>O positively correlates with SiO<sub>2</sub> for the MLC mafic rocks whereas the felsic rocks are scattered. K<sub>2</sub>O does not correlate with SiO<sub>2</sub> and the MLC felsic rocks are scattered in the diagram. Al<sub>2</sub>O<sub>3</sub> does not vary with SiO<sub>2</sub> in the MLC mafic rocks and decreases with SiO<sub>2</sub> for the felsic rocks. The Mg# is scattered, with a general decrease with increasing SiO<sub>2</sub>. The Čistá granite and granodiorite mostly merge with the trends of the MLC rocks and overlap the group of the MLC felsic rocks. The K<sub>2</sub>O content is higher in Čistá felsic rocks at a given SiO<sub>2</sub> content, compared to the MLC rocks. The MLC mafic and felsic samples are tholeiitic to normal-K calc-alkaline, while Čistá samples are normal-K to high-K calc-alkaline.

## 6.2 Trace element composition of the MLC mafic rocks

The MLC mafic rocks can be further subdivided into three types regarding their trace-element concentrations (Type-1 to Type-3; Supplementary Data, Table S2, Fig. 8). The main difference from Type 1 to Type 3 is the geochemical behavior of the REE (Fig. 8), characterized by a strong fractionation of LREE ((La/Sm)<sub>N</sub> = 0.7–6.2) compared to middle rare earth elements (MREE, (Sm/Dy)<sub>N</sub>= 0.8–3.1) and heavy rare earth elements ((Dy/Yb)<sub>N</sub> = 0.9–1.6).

The type-1 group is composed of two gabbroic samples (ca. 53 wt % SiO<sub>2</sub>, Mg# >74) that have cumulate-similar signatures. The samples are characterized by low REE contents (ΣREE<sub>N</sub>≤65.6), a U-shaped LREE pattern ((La/Sm)<sub>N</sub>=1.0 and 1.2, (Pr/Sm)<sub>N</sub>=0.5), and a flat MREE-HREE spectra ((Gd/Yb)<sub>N</sub>=1.1). They also have slight to pronounced positive Eu anomalies, Eu/Eu*=1.2 and 1.5 (Eu/Eu*=2Eu<sub>N</sub>/(Sm<sub>N</sub>+Gd<sub>N</sub>)). A primitive mantle (PM) normalized trace element diagram shows LILE enrichment marked by positive anomalies in Sr and Pb, and depletion in HFSE (Zr, Nb, Ta).
Type-2 rocks have more differentiated gabbroic composition (SiO₂=47.8–63.0 wt %, Mg#=40–67). They have flat to fractionated REE spectra ((La/Yb)N=1.1–13; (Gd/Yb)N=1.16–3.45)), with slight progressive LREE enrichment ((La/Sm)N=0.7–3.4). These samples have slightly negative to no Eu anomalies (Eu/Eu*=0.8–1), a flat to fractionated Pr/Sm ratios ((Pr/Sm)N=0.9–2.3) and are more enriched in REE (ΣREEₙ=159–465) compared to Type-1 samples. The trace element diagram shows positive Sr and U, variable Pb, Zr-Hf and small to strong negative Nb-Ta anomalies. The samples are generally enriched in LILE compared to HFSE.

Type-3 rocks have gabbroic composition (SiO₂=49.9–51.8 wt %, Mg#=45 and 59). They have the highest REE fractionation ((La/Yb)N=13–27), LREE over MREE and total REE enrichment ((La/Sm)N≤6.1–6.2; ΣREEₙ=651–1496) of all the mafic samples. They are associated with negative Eu anomalies (Eu/Eu*=0.6), negative Pb, Sr and Nb-Ta anomalies, and positive Th anomalies. The sample ML33C also shows a negative Zr-Hf anomaly.

## 6.3 Trace element composition of the MLC felsic rocks

The MLC felsic rocks can be subdivided into three types, referred to as Type-4 to Type-6 (Fig. 8). Type-4 is one sample with very singular chemistry, Type-5 correspond to the oligoclasite samples while Type-6 groups the other felsic rocks.

The type-4 rock is one trondhjemite sample of acidic composition (SiO₂=69.6 wt%, Mg#=59). Its REE spectrum is gently fractionated ((La/Yb)N=6.9) with a positive Eu anomaly (Eu/Eu*=2.2). The sample has a very low REE content (ΣREEₙ=19). On the extended trace element diagram, the spectrum displays pronounced positive Sr-Pb and negative Nb anomalies.

Type-5 rocks correspond to oligoclasite samples (SiO₂=58.7–60.23 wt %, Mg#=50–52). The rocks are REE fractionated (LaN/YbN=18–37) and show negative to no Eu anomalies
484 (Eu/Eu*=0.6–1). The extended trace element diagram shows enrichment in LILE, strong
485 negative Nb-Ta anomalies, positive Th-U anomalies, slight to strong positive Zr-Hf anomalies
486 and no to positive Pb-Sr anomalies. Sample ML28F has a singular fractionation in HREE
487 compared to the two other samples, which may be due to the presence of eclogitic host rock
488 despite the careful selection of samples.

489 Type-6 rocks correspond to one tonalite, four trondhjemite and three gneiss (dotted
490 lines in Fig. 8) (SiO2=62.6–84.01 wt %, Mg#=44–58). They all show strong LREE
491 fractionation ((La/Sm)N=2.48–5.63) except ML31B sample which is less REE fractionated
492 ((La/Sm)N=1.19). Compared to the trondhjemite HREE fractionation ((Ho/Lu)N=1.63–2.40),
493 gneiss, due to the presence of garnet, are inversely fractionated in HREE ((Ho/Lu)N=0.52–
494 0.92). Type-7 rocks have negative to slight positive Eu anomalies (Eu/Eu*=0.6 to 1.2). In the
495 extended trace element diagram, Type-7 rocks show strong Nb-Ta negative anomalies,
496 variable Pb, Sr and Zr-Hf anomalies. They share similar pattern with the Type-5 rocks.

497 **6.4 Trace element composition of the Čistá pluton**

498 Samples from the Čistá pluton are granodiorite and one rhyodacite (SiO2=68.1–71.5 wt %,
499 Mg#=36–45). They also display (Fig. 8b) a strong REE fractionation ((La/Yb)N=13–42),
500 moderate REE contents compared to the MLC felsic rocks (∑REE N=157–322) and small
501 positive Eu anomalies (Eu/Eu*=1.1 to 1.3). Compared to the MLC felsic rocks, they show
502 flatter MREE-HREE spectra ((Ho/Lu)N=0.73–1.18 for Čistá; 1.09–2.62 for MLC). The
503 extended trace element diagram shows a strong enrichment in U, slight Nb-Ta negative
504 anomalies and positive Zr-Hf and Sr anomalies. Furthermore, there is no chemical difference
505 between the Čistá granodiorite and the rhyodacitic dyke.

506 **7. GEOCHRONOLOGY RESULTS**
507 U-Th-Pb zircon geochronological analyses from selected samples were performed at the
508 laboratories of the Czech Geological Survey and at the University of California, Santa
509 Barbara. Samples ML44-1 (gabbro, Type-1), ML20-14-1 (quartz gabbro, Type-2), CI4E1
510 (granodiorite, Čistá) and CI5D1 (granodiorite, Čistá) were analyzed by LA-ICP-MS (Czech
511 Geological Survey, CGS). Samples ML31B (tonalite, Type-6), ML28F (oligoclasite, Type-5)
512 were analyzed by LASS (University of California, Santa Barbara, USA). While, sample
513 ML33B2 (tonalite, Type-6) was analyzed in both laboratories. Analytical spots were localised
514 in texturally homogenous domains based on cathodoluminescence (CL) imaging of zircon
515 (representative images are shown in Fig. 9). The complete analytical results are presented in
516 the Supplementary Data, Table S4, these data have been parsed through a 95–105%
517 concordance test, and plotted on probability density plots (Fig. 10), Terra-Wasserburg
518 concordia diagrams (Fig. 10), and as weighted mean plots (Fig. S3). Owing to significant
519 dispersion in the dataset, Devonian magmatic emplacement ages reported below are
520 approximated using the TuffZirc algorithm (Ludwig and Mundil, 2002).

7.1 Gabbro sample ML 44-1 (Type-1)

521 Zircon grains are 100 to 200 µm long, euhedral to subhedral, and have short elongate forms
522 with an aspect ratios of 1:1.5–1:2 (Fig. 9). In CL images, some zircon grains present a dark
523 core surrounded by a bright rim; others show the opposite. Zoning is commonly oscillatory,
524 but some grains exhibit homogeneous or more complex structures. 40 concordant analyses on
525 26 zircons were obtained from this sample. Two significant populations are apparent in the
526 probability density plot (Fig. 10a), a minor population at the Ediacaran-Cambrian boundary
527 (n=5) and a larger Devonian population (n=35). Within the Devonian population $^{238}$U/$^{206}$Pb
528 ages spread from 390±11 Ma to 358±15 Ma, the youngest dates are typically from thin
529 overgrowths. The TuffZirc algorithm finds a coherent group of 22 analytical point with a
530 weighted mean age of ca. 383 Ma (Fig. S3), the discarded younger ages are considered the
531
product of lead loss during metamorphic overprint. A concordia age of 384±2 Ma (MSWD = 0.64) is calculated for the 22 analytical points (Fig. 10b) and is considered to approximate the time of emplacement.

## 7.2 Quartz gabbro sample ML 20-14-1 (Type-2)

Zircon from this sample is 70 to 200 µm long, euhedral to sub-euhedral with long to short prismatic or rounded shapes and aspect ratio 1:1–1:3 (Fig. 9). In CL images, most zircon grains have a dark core and bright rim, fewer grains have a bright core and dark rim. The zonation is commonly oscillatory, locally homogeneous. 35 concordant analyses on 27 zircons were obtained from this sample. Three populations are observed in the probability plot (Fig. 10a) at ca. 540 Ma (n=12), 500 Ma (n=4) and 380 Ma (n=18). The older (ca. 540 Ma) dates are found in the dark cores, and in one grain in both dark core and white rim, intermediate dates (ca. 500 Ma) are associated with dark-grey cores, whereas the Devonian dates are associated with more grey and brighter zones. A coherent Concordia age of 385±1 (MSWD = 0.54) is calculated from the Devonian dates and inferred to approximate the timing of emplacement (Fig. 10c).

## 7.3 Oligoclasite sample ML 28F (Type-5)

The zircon grains range from 150 to 300 µm long and show general long-prismatic euhedral morphology of 1:3–1:5 aspect ratio (Fig. 9). CL images show zircon with a dark-grey core and a brighter rim. The zonation is commonly oscillatory to homogeneous. 34 concordant analyses were obtained from 9 zircons. The calculated $^{238}$U/$^{206}$Pb ages give a semi-continuous spread from the Early Cambrian to the Early Carboniferous (Fig. S3) with the most significant cluster in the Mid-Devonian (Fig. 10d). The older dates are most commonly associated with grey cores, while the younger dates are associated with bright rims and grey cores. From the Devonian data the TuffZirc algorithm identifies a population (n=5) with a weighted mean age
556

of 389±4 Ma (Fig. S3). A concordia age of 390±3 Ma (MSWD = 1.7) is calculated from these
grains (Fig. 10e).

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558

## 7.4 Tonalite sample ML 31B (Type-6)

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565

Zircon is 150 to 350 µm long, with euhedral, long-prismatic, short-prismatic and more
rounded shapes with aspect ratios of 1:1–1:3 (Fig. 9). CL images mostly show a dark to grey
core surrounded by brighter rims, a few have the opposite. 67 concordant analyses were
obtained from 22 zircons. Only Devonian dates were recorded in this sample, but, with
significant spread from 413±4 Ma to 367±14 Ma. The TuffZirc algorithm identifies a main
population of 36 analytical points, which yield a concordia age of 386±1 Ma (MSWD = 1.4,
Fig. 10f), inferred to be the approximate timing of emplacement.

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## 7.5 Tonalite sample ML 33B2 (Type-6)

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Zircon is 50 to 200 µm long with euhedral long to short prismatic and rounded shapes, with aspect ratios of 1:1–1:3 (Fig. 9). In CL images, zircon grains usually present a dark core surrounded by a bright rim, whereas some have a bright rim around a dark core. Zonation is commonly oscillatory, but some grains have homogeneous or more complex structures. For this sample, 94 zircons were analyzed, yielding 120 concordant dates. The data exhibit a semi-continuous spread of dates from the Late Ediacaran to the end of the Devonian (Fig. S3). Maxima are identified at the Ediacaran-Cambrian boundary, Late Cambrian, and Mid-Devonian (Fig. 10d). The older dates are associated to CL-darker cores and the intermediate zones, while the younger dates are found in CL-dark to white rims, or in CL-dark and white homogeneous zircons. Within the Devonian data the TuffZirc algorithm identifies 40 data points, which produce a concordia age of 386±1 Ma (MSWD = 0.91, Fig. 10g).

578

## 7.6 Granodiorite sample CI4E1 (Čistá)
579 Zircon is 100 to 200 µm long, euhedral to sub-euhedral and has a rounded to elongated
580 prismatic shape with aspect ratios of 1:1–1:5 (Fig. 9). CL images show oscillatory zonation
581 with usually a dark-grey core and rims. Some brighter cores are also observed. 81 concordant
582 ages were measured from 54 zircons. Three analytical spots give Mid-Cambrian dates, while
583 the bulk of the data corresponds to a Mid-Devonian maxima (ca. 392 Ma) and Late Silurian
584 and Late Devonian shoulders (Fig. 10h). Five data points have consistent Late Silurian age
585 (ca. 426 Ma); however, these are associated with patchy or irregular zoned zircon cores (Fig.
586 9). The Mid-Devonian maxima is comprised of 55 analytical data points, which yield a
587 concordia age of 391±1 Ma (MSWD = 1.1, Fig. 10i). A further 13 data points give a Late
588 Devonian concordia age (364±2 Ma, MSWD = 1.3, Fig. 10i). The Mid-Devonian population
589 is associated with both CL-grey homogenous cores and oscillatory zoned rims; while the Late
590 Devonian population is observed in oscillatory zoned rims. Only a few grains preserve both a
591 Mid-Devonian core and Late Devonian rim (Fig. 9)

592 **7.7 Granodiorite sample CI5D1 (Čistá)**

593 Zircon is 100 to 200 µm in length, mostly euhedral to sub-euhedral with an elongated
594 prismatic shape with aspect ratio 1:2–1:3. CL images show dark homogeneous oscillatory
595 zonation, with in a few cases a black/white core (Fig. 9). This sample is notable for a high
596 degree of discordance, only 20 concordant spots were acquired on 19 zircons from 127
597 analytical spots. Since this discordance plots randomly towards common lead and not along
598 discrete discordia lines, it is likely a product of metamictization as a result of high U content
599 in zircon (mean U content = 3126 ppm). From the concordant data two analytical spots gave
600 Paleo-Proterozoic (ca. 1.8 Ga) dates (Fig. 10j). The remaining data produces two maxima in
601 the probability plot (Fig. 10h) at ca. 390 and 370 Ma. The Mid-Devonian population (n=6)
602 yields a coherent concordia age of 394±3 Ma (MSWD = 0.79, Fig. 10k), while the late
603 Devonian data (n=12) yields a concordia age of 369±2 Ma (MSWD = 0.39, Fig. 10k). Some
604

grains exhibit clear core and rim systematics with Mid-Devonian core and oscillatory zoned
Late-Devonian rim (Fig 9).

605

606

8. Sr AND Nd WHOLE ROCK ISOTOPIC COMPOSITION

607

Figure 11a shows initial isotopic ratios ¹⁴³Nd/¹⁴⁴Nd vs ⁸⁷Sr/⁸⁶Sr of the MLC and Čistá pluton
samples (data is presented in Supplementary Data, Table S6) compared to previously
published data of other MLC lithologies. An age of 386 Ma (mean of the MLC mafic and
felsic samples according to histograms in Fig. 10) and 370 Ma (last crystallization age in the
Čistá samples according to our data in Fig. 10) were used to calculate initial ratios of the
MLC and the Čistá magmatic rocks, respectively. The depleted mantle (TDM) Nd model ages
have been calculated using the composition of DMM presented in Workman and Hart (2005)
for the ¹⁴³Nd/¹⁴⁴Nd ratio and Salters and Stracke (2004) for the trace elements used to
recalculate the ¹⁴⁷Sm/¹⁴⁴Nd ratio.

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The MLC mafic rocks (Type-1 to Type-3) exhibit some variability in initial Sr ratios
($^{87}$Sr/$^{86}$Srᵢ=0.704167–0.705051) and comparatively limited range for Nd ratios
($^{143}$Nd/$^{144}$Ndᵢ=0.51226–0.51244); corresponding to moderately positive εNdᵢ values (+2.3–
+5.8; Fig. 11a, b and Supplementary Data, Table S6). The studied felsic rocks have
overlapping but, even more limited variation in Sr ratios ($^{87}$Sr/$^{86}$Srᵢ=0.704168–0.704921) and
greater variability for Nd ratios ($^{143}$Nd/$^{144}$Ndᵢ=0.512187–0.512555). This results in a spread
of εNdᵢ values from near unity to strongly positive (+0.8–+8.0). For the most juvenile sample
(ML14-14-1, Type-4) the calculated TDM age approaches its inferred emplacement age (384
Ma). For the rest of the felsic samples Neoproterozoic TDM ages (ca. 670–900 Ma) are
calculated. Four of the studied mafic samples have similar range of TDM ages (ca. 500–900
Ma); while the three remaining samples have spurious older (ca. 1300 Ma) or negative ages as
a result of high ¹⁴⁷Sm/¹⁴⁴Nd ratios approaching DMM values.
Samples from the Čistá pluton have lower initial Sr ratios ($^{87}$Sr/$^{86}$Srᵢ=0.7035–0.7037) than the MLC rocks, and Nd values corresponding to the mid to lower part of the domain occupied by the MLC rocks ($^{143}$Nd/$^{144}$Ndᵢ=0.512179–0.512383). Calculated εNdᵢ values are weakly to moderately positive (+0.4–+4.3) and Neoproterozoic TDM ages (600–850 Ma) overlap with the majority of the MLC samples.

## 9. DISCUSSION

The MLC gabbroic and felsic rocks investigated in the present study have previously been interpreted as a migmatitic part of the amphibolite and eclogite assemblage resulting from their granulite facies reequilibration and melting (Timmermann et al., 2004). However, this study shows that the petrography, geochemistry and geochronology of the mafic magmatic rocks differ substantially from the host rock eclogite and amphibolite, and the C–O metagabbro. In the following sections we present petrological, geochemical and geochronological arguments for an igneous magmatic arc nature of the studied MLC rocks and their possible affinity to the Čistá pluton. After discussing the petrogenesis of the MLC mafic and felsic magmatic rocks we propose a new geodynamic model involving a distinct Mid-Devonian arc related to stacking of an eclogitic wedge beneath the continental margin.

### 9.1 Petrography and microstructural arguments for igneous evolution of the MLC and Čistá rocks

Unlike the typical MLC amphibolites, the studied MLC gabbros show magmatic features, for example a magmatic compositional layering marked by alternating leucocratic and melanocratic types of gabbro or quartz-gabbro in petrographic continuity (Fig. 3a). Fine- to coarse-grained trondhjemite and oligoclasite commonly exhibit sharp contacts to the surrounding gabbro (Fig. 3b), indicating the mostly late intrusive nature of the dykes, but in some places they are in petrographic continuity with the host-rock gabbro (Fig. 3b), suggesting that some of them crystallized together. The sharp contacts of the gabbro and
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trondhjemite and oligoclasite with retrogressed eclogite indicate intrusive relationships with regards to the metamorphic edifice (Fig. 3e, f).

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Compared to the C–O metagabbro, MLC mafic rocks have a different mineralogy as they contain mostly amphibole and only exceptionally subordinate pyroxene, pointing to a hydrous magma. In contrast, the C–O metagabbro is composed of an anhydrous assemblage of Opx–Cpx–Pl±Amp (Jašarová et al., 2016). It should also be noted that the coronitic and symplectite textures that overprint the primary magmatic assemblage of the C–O metagabbro (Štědrá, 2002; Jašarová et al., 2016) are not present in the studied MLC mafic rocks.

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The grains in the gabbro are subhedral but also form straight grain boundaries with triple points and are similar in size (Fig. 5), suggesting static re-equilibration related to grain boundary reduction due to high temperatures after magma intrusion. The composition of amphibole, ranging mostly from Mg-hornblende to tschermakite, points to its HT nature and is compatible with a magmatic origin. The magmatic to solid-state foliation of the MLC magmatic rocks is gently east dipping, and parallel to the regional S3 foliation (e.g. Peřestý et al., 2017, 2020; Collett et al., 2018), suggesting that these magmatic rocks were emplaced before or at the beginning of the major phase of exhumation. Because of these petrographic and structural differences, the MLC mafic rocks studied here are considered to be independent from the eclogite-amphibolite and C–O metagabbro, and have their own petrogenesis and emplacement history.

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## 9.2 Geodynamic context based on whole rock chemistry

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The eclogite-amphibolite major element chemistry (Figs 6, 7) is broadly similar to the MLC mafic samples, while the C–O metagabbro composition is close to the most basic MLC mafic rocks, but with higher MgO and CaO contents. On trace element diagrams (Fig. 8), Types 1 and 2 mafic rocks show broad similarity to the MLC eclogite-amphibolite and C–O

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metagabbro. However, Type-1 mafic rocks show a notably bigger positive Sr anomaly

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Sr/Sr*=18 and 50 (Sr/Sr*=2*SrN/(PrN+NdN)) in comparison to the typical composition of the

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eclogite-amphibolite (Sr/Sr*=0.4–14.0, mean of 2.3) and C-O metagabbro (Sr/Sr*=0.5–3.0, mean of 1.31). Type-3 mafic rocks exhibit relative enrichment in LREE over MREE

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(LaN/GdN=9 and 10) compared to the eclogite-amphibolite (LaN/GdN=0.3–7.1, mean of 1.4)

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and C-O metagabbro (LaN/GdN=1.5–2.4, mean of 2.0).

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Geochemical data can also provide insight into the tectonic setting of the MLC and

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Čistá magmatic rocks. The major elements indicate a clear calc-alkaline trend, for both the

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MLC and Čistá magmatic rocks (Fig. 6). Moreover, the discriminant diagrams of Pearce *et al.*

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(1984) for felsic rocks and of Wood (1980) for basic rocks presented in Fig. 12 show that

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most of the samples fall in the volcanic-arc related fields. In addition, the normalized multi-

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element plots (Fig. 8) show typical patterns of arc-related rocks. This is indicated by the

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enrichment in mobile relative to immobile elements (LREE vs HREE, LILE vs HFSE) which

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is independent of the nature of the overlying mantle wedge and is assumed to be subduction

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derived. This pattern is attributed to metasomatism of the mantle source of arc basalts by

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fluids released from the subducted slab (Pearce, 1983). Also, a negative Nb-Ta anomaly is

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ascribed to many processes (Kelemen *et al.*, 2003), but most likely is linked to the

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fractionation of niobium and tantalum from other highly incompatible elements via partial

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melting of subducting, eclogite facies basalt or sediment with residual rutile (Ryerson and

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Watson, 1987; Kelemen *et al.*, 1993, 2003; Brenan *et al.*, 1994; Elliott *et al.*, 1997; Turner *et

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*al.*, 1997; Foley *et al.*, 2000). This last assumption supposes that all arc magmas include a

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subduction component derived from partial melting of subducted material in eclogite facies,

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with residual rutile. If only dehydration of the slab occurs, Nb-Ta would be fractionated

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together with the HFSE couple Zr-Hf (Foley *et al.*, 2000). However, the important amount of

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inherited zircon in our samples may have altered the Zr-Hf chemistry. Figure 13 also shows a
trend toward arc volcanism chemistry. In the Th/Yb vs Nb/Yb diagram, basic rocks of Type-1 to -3 are spread between MORB and the volcanic arc array. The MLC felsic rocks and Čistá samples also plot close to the volcanic arc array. In the TiO₂/Yb vs Nb/Yb diagram, the samples generally plot within the shallow melting array, which is compatible with a volcanic arc tectonic context. Finally, εNd-Sr isotopic results (Fig. 11b) also provide evidence of an arc affinity. The studied MLC rocks exhibit less radiogenic Nd compositions than published data from MLC eclogite and amphibolite (Beard *et al.*, 1995; Timmermann *et al.*, 2002) or the Depleted Mantle composition in the Mid-Devonian (Workman and Hart 2005); but have comparatively more radiogenic Sr compositions. One potential explanation for these compositional trends is mixing between a Depleted Mantle source and subducted sediments (i.e. GLOSS (Plank and Langmuir, 1998) in Fig. 11b); however, the samples are only weakly correlative with such a mixing line, suggesting that there are other potential sources involved. These could include the felsic orthogneiss at the base of the MLC (Timmermann *et al.*, 2002) which has a slightly negative εNd value and has more radiogenic Sr than the studied samples (Fig. 11b) or the meta-sediments at the base of the Blovice Accretionary Complex, which are similarly depleted in radiogenic Nd and enriched in radiogenic Sr (Drost *et al.*, 2007). In contrast the Čistá samples appear to plot on a coherent mixing line between a Depleted Mantle and a lower crustal (e.g. Willbold and Stracke, 2010) source.

## 9.3 Timing of the events

### 9.3.1 MLC magmatic rocks

The geochronological results of the MLC magmatic rocks show consistently a main Mid-Devonian date peak in both the mafic and felsic samples (Fig. 10). The magmatic aspect of these zircons (Fig. 9), combined with the field, petrographic, and micro-structural arguments presented above suggest that the Mid-Devonian data approximates the emplacement ages. The calculated concordia ages spread from 393 to 382 Ma indicating that this magmatic event
occurred contemporaneously or immediately following the ca. 390 Ma eclogitization of the
surrounding oceanic crust (Collett et al., 2018). Notably, the magmatic event is in
contemporaneity with the HT overprint of the eclogites dated on zircon by Timmermann et al.
(2004), and occurs before the exhumation of all the MLC system at ca. 375 Ma, inferred from
Ar/Ar cooling ages and partial resetting of the Sm-Nd and Lu-Hf systems Collett et al., 2018).
Late Devonian data from the studied magmatic samples is inferred to represent Pb loss during
this exhumation event. This is in accordance with preservation of the magmatic features of
these rocks in some places, with the absence of the eclogite-facies metamorphism and with
the local development of magmatic- to solid-state deformation exhumation fabric.

The second feature is the large quantity of Early to Late Cambrian dates, typically
found in inherited zircon cores (Fig. 9), the dates fall into two age groups: an Ediacaran-
Cambrian peak around 540 Ma and a Late Cambrian peak around 500 Ma. The ca. 500 Ma
peak overlaps with the age of the nearby C–O gabbro (Timmermann et al., 2006), and the ca.
540 Ma peak correlates well with the protolith ages of the MLC eclogite (Timmermann et al.,
2004) as well as to the main age peak in the metasediments of the Teplá Crystalline Complex
(Žák et al., 2020). Therefore, these zircon cores may represent xenocrystic zircon inherited
from the host rocks, reflecting an important recycling process that occurred during the arc
magmatism event. This observation is also compatible with the isotopic data, showing low Nd
values that can be explained by contribution of the lower crust by recycling (Fig. 11).

### 9.3.2 *The Čistá pluton*

Zircon data from the Čistá pluton show a Middle Devonian main peak at ca. 392 Ma and two
groups of dates at ca. 426 Ma and ca. 370 Ma. We interpret the youngest date of ca. 370 Ma
as the age of emplacement of the Čistá pluton in the upper crust, this is in agreement with the
previous Pb–Pb zircon evaporation age reported by Venera et al. (2000). However, a certain
degree of caution has to be applied when interpreting the older data from the Čistá samples
owing to the high degree of discordance observed in the data (only 16 % of analytical spots in CI5D1 were concordant and 58 % in CI4E1). As stated above, the discordance is likely connected to metamictization as a result of high U contents in zircon (mean U content of 3126 ppm in zircon from CI5D1 and 1385 ppm in zircon from CI4E1). Since small degrees of metamictization can go unnoticed in analytical spots that plot close to, or on, the concordia (e.g. Mezger and Krogstad, 1997) it is important to carefully consider the geological significance of the analytical data. This is evident from the Silurian population in sample CI4E1 where even though the analytical data is concordant, the spots were located in patchy or irregular zones in zircon cores (zircon 23 in Fig. 9). Silurian ages also have spurious geological significance since, although similarly aged volcanics do occur within the Prague Basin (e.g. Tasáryová et al., 2018) these volcanics are dominantly basic in composition and unlikely to be associated with significant zircon growth. However, this is not the case for the Mid-Devonian age peak, which is dominant in sample CI4E1 and subordinate in CI5D1. These data points were measured on homogeneous CL-grey cores and oscillatory zoned zircon rims (Fig. 9) and the acquired ages overlap with those from the MLC magmatic rocks. Therefore, these ages are considered to represent antecrysts that formed following partial crystallization of magma in a deep magma chamber associated with the arc magmatism in the MLC.

Considering the older inheritance, the three zircons from sample CI4E1 with Mid-Cambrian age are interpreted as xenocrystic inherited from the host Cambrian Tis granite (Venera et al., 2000) or from deep crustal migmatites of this age reported by Peřestý et al. (2017). Two zircons from sample CI5D1 yielding Paleo-Proterozoic dates may also be xenocrystic associated with assimilation of the Teplá-Barrandian crust that commonly shows such dates in detrital zircon spectra (Hajná et al., 2017; Žák et al., 2020).

## 9.4 Gabbro setting and chemical variability of Type-1 to Type-3 samples
The elementary and isotopic chemical diversity of magmatic rocks in arc magmas is usually assigned to two main components (Dhuime et al., 2009). The first is the mantle wedge, which can have various chemical characteristics, from depleted to more enriched zones, due to recycling process (Morris and Ryan, 2003). The second component is linked to the subduction process because it corresponds to fluids or melts derived from the subducting slab and migrating into the mantle wedge (Ellam and Hawkesworth, 1988; Plank and Langmuir, 1993, 1998; Miller et al., 1994; Elliott et al., 1997; Hawkesworth et al., 1997; Class et al., 2000; Johnson and Plank, 2000; Turner et al., 2000, 2003; Elliott, 2003; Kelemen et al., 2003). The subduction component is associated with an enrichment in more mobile (LILE, U, Pb, Th and LREE) relative to more immobile elements (HFSE and HREE).

Type-1 gabbros are the most primitive rocks sampled in this study, as they show the smallest contribution of subduction. They have cumulate like signature, which may suppose a change in their chemistry with respect to a real magmatic liquid composition. However, isotopic data indicates that the DMM was the main source in this magmatic system (Fig. 11b). In the following trace elements modelling, we try to model the Type-1 rock chemistry starting from a DMM source. The gabbro setting can be modeled by dehydration of the slab and associated melting of the mantle wedge by hydration (Fig. 14a). Dehydration was modeled using fractional melting of the slab (with both altered oceanic crust and sediments as a source, with temperature dependent partition coefficients corresponding to a fluid context; from Kimura, 2017). Melting of the mantle was modeled by batch melting of the hydrated DMM mantle fluxed by slab derived fluids (using partition coefficients corresponding to a melt context, from Kimura, 2017). Also, the dry DMM melting model (batch melting) has been added for comparison. Fig. 14a shows that the ML44-1 sample pattern is generally similar to the one of the dry DMM model except the La-Ce-Pr LREE that are enriched in the
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ML44-1. This LREE enriched pattern can be approached by addition of 5% of slab derived fluid into the DMM prior to its melting (DMM + fluids model in Fig. 14a).

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Furthermore, the chemistry of Types-1 to 3 shows a clear enrichment in fluid mobile elements relative to immobile ones (LREE vs HREE, LILE vs HFSE), together with the development of a pronounced Nb-Ta negative anomaly. This may be assigned to an increasing contribution of the subduction component in the source, which is in accordance with what was previously highlighted in Fig. 13, where an evolution from MORB to a volcanic arc array is observed from Type-1 to Type-3. This evolution can also be observed in the Supplementary Data, Fig. S4 where the different mafic rock Types, based on increasing LREE fractionation (Pr/Sm), can be correlated with an arc signature enrichment (Th/Yb). It is possible to discriminate between aqueous fluids or melt contribution using trace elements that have contrasting behavior according to the transport system. The contribution of a fluid and/or melt component in the Type-1 to Type-3 groups of the MLC can be shown using the U/Th and Sr/Th ratios as a function of Th (Fig. 14b; McCulloch and Gamble, 1991). The evolution on these plots indicates a main contribution of fluids to Type-1, slightly transitory in Type-2 while the Type-3 component is slab-melt derived. However, isotopic data (Fig. 11b) indicate that the contribution of the GLOSS into the system is limited (less than 2%). Chemical variations can be also derived from contamination by the material present in the accretionary wedge or the lower crust (LC) as pointed out by the isotopic data and the inherited zircons. According to the isotopic data, if the LC is chosen as a main contamination source, its contribution to the system is at maximum 10%. However, as the contaminant material composition (the amphibolite-eclogite and metagabbro) is quite similar to the MLC mafic rocks composition (as shown by isotopic or major and trace elements data), its contribution is less visible, even if bigger than the one of GLOSS. Nevertheless, the DMM remains the main component in the MLC magmatic rocks.

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## 9.5 Petrogenesis of the felsic rocks

A first hypothesis concerning the origin of the felsic rocks was provided by Timmermann *et al.* (2004), who interpreted the sample HT-99-1 (equivalent to our felsic rocks) to come from melting of the host eclogite during exhumation, because of a close isotopic affinity to that of a MLC HP granulite and the amphibolite. The major and trace elements of the sampled felsic rocks show a clear arc magmatism signature, however, inviting reconsideration of this question. In the context of an arc, several hypotheses are feasible and presented in Fig. 15.

First, the felsic rocks may be the result of arc gabbro differentiation by fractional crystallization. According to the mineralogy of the gabbro, amphibole and plagioclase were the main fractionating minerals, with accessory titanite. Fractional crystallization was modeled using the Petromodeler Excel™ sheet (Ersoy, 2013) and performed with two different initial gabbro compositions: the most primitive ML44-1 sample and the gabbro mean chemistry. Results compared to the different felsic rock types are presented in Fig. 15a. The modelling does not fit well to the felsic compositions, especially for the HREE. Furthermore, this process implies parallelism between the spectra and unchanging ratios between different elements during differentiation (Fig. 15a, in a closed system). However there is a striking difference in the HREE fractionation between the mafic rocks which show flat spectra ((Dy/Yb)$_{N}$=0.9–1.6, mean 1.2) and the felsic rocks (except gneiss) which are fractionated ((Dy/Yb)$_{N}$=1.6–3.4, mean 2.2). It could be explained by the growth of garnet, but this is unlikely because garnet has not been observed in the studied mafic rocks.

The second hypothesis follows the suggestion of Timmermann *et al.* (2004), that the felsic magmatic rocks are linked to dehydration melting. This process is capable of generating granitic melts at lower crustal depth (Johannes and Holtz, 1996). It has already been suggested in an island arc setting (Garrido *et al.*, 2006), where granitic melts are interpreted to be produced by dehydration-melting of arc root plutonic rocks. For instance, in the Kohistan
arc, Garrido *et al.* (2006) interpreted granitic rocks to come from dehydration melting of amphibole-bearing rocks in the lower part of the magmatic arc, leaving garnet-bearing restites. In the MLC, arc related amphibole-bearing lithologies are represented by Type-1 to Type-3 rocks. Other amphibole-bearing lithologies are present in the area and could be the source for dehydration melting. These are the retrogressed eclogite/amphibolite, as proposed by Timmermann *et al.* (2006), and the C–O metagabbro. Modelling of dehydration melting was carried out using batch melting equations and temperature dependent partition coefficients (in a melt context) of Kimura (2017), trying the different amphibole bearing lithologies as a source, with the purpose to fit the geochemistry of the three different felsic types, gave the results shown in Fig. 15b–d. Type-4 sample pattern can be approached by modelling the melting of Type-1 gabbroic rocks. Type-5 (oligoclasite) and Type-6 (trondhjemite/tonalite) have similar trace elements patterns that fit well to the modelled liquid composition resulting from MLC arc mafic rocks melting. Furthermore, Fig. 15e presents both dehydration melting residue and liquids modelled from the MLC arc mafic rocks source. The modelled residue composition is similar to the garnet rich pyroxenites chemistry observed in Kohistan (Garrido *et al.*, 2006), that are interpreted to represent residues from amphibole bearing dehydration melting process, generating granitic melts. Such a residue could explain the unusual chemistry of the felsic material present within the MLC. As melting textures of these gabbro are not observed in the field, the source of the felsic rocks may correspond to a deeper part of the same Devonian arc system or to chemically similar but older rocks that may for instance correspond to the arc volcanics of the Jílové belt (Hajná *et al.*, 2014). Considering these results with the isotopic data (Fig. 11) where not all the MLC magmatic samples were measured, the MLC felsic rocks cover an area larger than the one defined by the analyzed mafic rocks. To a first order, the fact that the MLC mafic and felsic rocks are not separated in distinct groups does not preclude either of the two hypotheses, but
the incomplete isotopic data prevents further interpretation. Dating of the felsic rocks shows also an important inheritance of Cambrian to Ordovician zircon (except sample ML31B). This inheritance is clearly a result of contribution of the metamorphic basement and the C–O metagabbro in the genesis of felsic magma but could also stem from the inheritance already present in the arc-gabbro, which could be their alternative source.

## 9.6 Geodynamic model of magmatic arc emplacement

Based on the combination of our new and existing knowledge from the Mariánské Lázně Complex, Teplá Barrandian Unit, and Central Bohemian Plutonic Complex (see section 2.), we propose a new model for the geodynamic development of the Devono-Carboniferous active margin in the Bohemian Massif. This model is detailed in four stages reflecting the evolution of material entering the subduction zone and the geodynamic response in the upper plate (Fig. 16).

### *Stage 1 (Early- to Mid-Devonian, Fig 16a)*

The first stage in the geodynamic model is characterized by the entrance of oceanic or sub-oceanic material into the subduction zone, its transformation to eclogite and exhumation to the base of the upper plate (e.g. Collett et al., 2018; Ackerman et al., 2020). Eclogitized oceanic crust is not buoyant and therefore is generally exhumed beneath the MOHO, where imbricated thrust sheets are stacked in the form of a subduction wedge (Angiboust et al., 2012; Vogt and Gerya, 2014; Ruh et al., 2015; Menant et al., 2020). The earliest geochronological record of this stage is recorded by the Sm-Nd garnet core ages of Beard et al. (1995), which gave Early Devonian ages. Although, the peak of eclogite-facies metamorphism was likely later corresponding to the ca. 390 Ma Lu-Hf garnet ages of Collett et al. (2018). The earliest phase of arc magmatism identified in this study overlaps within error with the peak of eclogite-facies metamorphism and is further recorded in inherited
zircon from the Čistá pluton (Figs. 9, 10) and in detritus in Givetian siliciclastics in the uppermost part of the Prague Basin (i.e. Srbsko Formation, Strnad and Mihaljevič, 2005).

Stage 2 (Mid- to Late-Devonian, Fig. 16b)

The second stage of the model coincides with the introduction of a more-continental type material into the subduction zone (Schulmann et al., 2014; Konopásek et al., 2019) that caused compression and thickening of both the MLC eclogitic wedge and the hangingwall TBU upper plate. This triggers the intrusion of gabbro and associated felsic rocks parallel to the vertical S2 mechanical anisotropy (Fig. 16b) and ca. 380 Ma granulite/amphibolite-facies re-equilibration of the eclogite-facies assemblages at pressures up to 14 kbar and temperatures up to 850 °C (Štědrá, 2001; Collett et al., 2018). These conditions are geothermally consistent with those of the metamorphosed Late Cambrian gabbro reported by Jašarová et al. (2016) and from the overlying TBU (Peřestý et al. 2017), indicating that lower crust of the upper TBU plate experienced heating and thickening together with the eclogitic part of the subduction wedge. The intruding gabbroic magma was also contaminated with host rocks of the eclogitic wedge, as documented by the presence of abundant Early Cambrian zircons.

Stage 3 (Late-Devonian – Early Carboniferous, Fig. 16c)

The third stage of the model relates to extensional exhumation in the upper plate along detachment shear zones, which juxtaposed the TBU mid-crustal rocks against the MLC subduction wedge – arc complex (Fig. 16c; Peřestý et al., 2017; Collett et al., 2018). This event was characterized by a transition from a sub-solidus fabric to HT solid state gently dipping S3 foliation (Fig. 3c, d) associated with progressive cooling (Dallmeyer and Urban, 1998) and rapid exhumation of the whole subduction wedge. The emplacement age of the Čistá pluton into upper crust occurred at ca. 369–364 Ma along a west-dipping detachment zone (Venera et al., 2000), which fits also with ca. 380–365 Ma crystallization ages of the Staré Sedlo and Mirotice sheeted sill-like intrusions that were equally emplaced and solidified
925 in the trans-tensional regime (Košler et al., 1993a, b; Tomek et al., 2015). Finally, the MLC,
926 Čistá and Staré Sedlo -Mirotice arc type rocks show similarly high εNd values ranging from +1
927 to +6 indicating their juvenile source (Fig. 16e). This regional extensional event was most
928 likely triggered by the exhumation of a previously subducted continental bulge during
929 ongoing subduction of a remnant oceanic basin (Fig. 16c, Peřestý et al., 2017).

930 Stage 4 (Early- to Mid-Carboniferous, Fig. 16d)

931 The final stage of the model is related to the main collisional event and is characterized by: 1)
932 relamination of previously subducted continental crust beneath the TBU upper plate
933 (Schulmann et al., 2014; Maierová et al., 2021), 2) exhumation of granulitized relaminant in
934 form of crustal scale diapiric structures at the east (Franěk et al., 2011; Nahodilová et al.,
935 2014, 2020), 3) subduction of the remaining oceanic crust and arrival of main body of the
936 Saxothuringian continent in the west (e.g. Konopásek and Schulmann, 2005; Konopásek et
937 al., 2019). The latter was responsible for growth of continental composite magmatic arc at the
938 eastern termination of the TBU – the Central Bohemian Plutonic Complex between 354 and
939 337 Ma (Fig. 16d). The presence of deeply subducted continental crust in the source, later
940 interpreted as relaminant (Lexa et al., 2011; Schulmann et al., 2014), was documented by
941 geochemical signature of CPBC granitoids and their low +1 to -8 εNd values (Janoušek et al.,
942 1995, 2000; Janoušek and Holub, 2007). Importantly, the Staré Sedlo and Mirotice arc meta-
943 granitoids previously emplaced and gneissified in trans-tensional regime were, together with
944 their Proterozoic and Early Paleozoic host rocks, intruded by CPBC granitoids that were
945 emplaced in transitional transpressional – extensional regime (Fig. 14d, e). The above data
946 show that the two magmatic arcs are characterized by two distinct ca. 390–370 Ma and ca.
947 354–335 Ma periods related to independent compression-extension cycles (Fig. 16e). The
948 most likely reasons for these two magmatic-tectonic cycles (panels in Fig. 16) are complex
949

tectono-thermal processes related to subduction of previously rifted Gondwanan continental

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margin beneath the active margin of the TBU plate (Konopásek et al., 2019).

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## 10. CONCLUSION

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Mafic and felsic rocks of previously unrecognized magmatic nature, and/or formerly

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interpreted as being derived from the amphibolite-eclogite unit of the Mariánské Lázně Complex

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were studied together with the Čistá granitoid pluton intruding the hangingwall

955

TBU.

956

* Petrography and geochemistry of the studied rocks present within the MLC amphibolite-eclogite unit support their magmatic origin, and are compatible with generation in a primitive magmatic arc.

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* Isotopic results and trace element modelling suggests that the MLC amphibole gabbro derived from a DMM mantle composition. Chemical variation in the mafic rock types can be explained by an increasing contribution of the slab-derived component into the system.

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962

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* Modelling the chemistry of the felsic rocks shows that they are not the result of fractional crystallization of the gabbroic rocks but rather issued from dehydration melting of mafic rocks with arc type chemistry.

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* Zircon dates indicate that the arc was emplaced in mid-Devonian time for both the MLC and Čistá pluton, with a two-phase history for the Čistá pluton. Zircon age inheritance in the MLC implies that the magma was contaminated by the rocks from the accretionary wedge into which it was emplaced, whereas in the Čistá pluton the inherited zircons are from the TBU crust.

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* The studied magmatic rocks record distinct magmatic arc event related to Early-Mid Devonian stacking of subduction wedge beneath TBU upper plate that was followed by formation of second Carboniferous continental arc further east.

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973

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## 11. Funding

975

This work was supported by the Czech Science Foundation [GAČR 17-22207S] and Research Project no. 310760 which is a part of the Strategic Research Plan of the Czech Geological Survey [DKRVO/ČGS 2018-2022]; scholarship of V.P. was supported by the Center for Geosphere Dynamics grant [UNCE/SCI/006] and B.H. and A.K.C. also acknowledge the support of the National Science Foundation [EAR-1551054].

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## 12. Acknowledgments

981

Martin Racek and Radim Jedlička are thanked for assistance with collection and processing of SEM and EPMA data at Charles University. Thierry Perrone is also thanked for assistance with collection of geochemical data at Laboratoire d'Hydrologie et de Geochimie de la Surface.

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## 13. REFERENCES

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Ackerman, L., Krňanská, M., Siebel, W., Strnad, L., 2010. Geochemistry of the Drahotín and Mutěnín intrusions, West Bohemian shear zone, Bohemian Massif: contrasting evolution of mantle-derived melts. Mineral. Petrol. 99, 185–199.

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Ackerman, L., Hajná, J., Žák, J., Erban, V., Sláma, J., Polák, L., Kachlík, V., Strnad, L., Trubač, J., 2019. Architecture and composition of ocean floor subducted beneath northern Gondwana during Neoproterozoic to Cambrian: A palinspastic reconstruction based on Ocean Plate Stratigraphy (OPS). Gondwana Res. 76, 77–97.

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1282 Štědrá, V., Kryza, R., Kachlík, V., 2002. Coronitic metagabbros of the Mariánské Lázně Complex and Teplá Crystalline Unit: inferences for the tectonometamorphic evolution of the western margin of the Teplá-Barrandian Unit, Bohemian Massif. In: Winchester, J.A., Pharaoh, T.C., Verniers, J. (eds). Palaeozoic Amalgamation of Central Europe. Geol. Soc. Spec. Publ. 201, 217-236.

1287 Stern, R.J., 2002. Subduction zones. Rev. Geophys. 40, 3-1-3-38.
1288 Strnad L., Hladil J., 2001. Geochemistry and composition of the Middle Devonian Srbsko

1289 Formation in Barrandian Area, Bohemian Massif: a trench or fore-arc strike-slip basin fill

1290 with material from volcanic arc of continental margin? Geolines 13, 111–114.

1291 Strnad, L., Mihaljevič, Martin., 2005. Sedimentary provenance of Mid-Devonian clastic

1292 sediments in the Teplá-Barrandian Unit (Bohemian Massif): U-Pb and P-Pb geochronology of

1293 detrital zircons by laser ablation ICP-MS. Mineral. Petrol. 84, 47–68.

1294 Tasáryová, Z., Janoušek, V., Frýda, J., 2018. Failed Silurian continental rifting at the NW

1295 margin of Gondwana: evidence from basaltic volcanism of the Prague Basin (Teplá–

1296 Barrandian Unit, Bohemian Massif). Int. J. Earth Sci. 107, 1231–1266.

1297 Timmermann, H., Štědrá, V., Gerdes, A., Noble, S. R., Parrish, R. R., Dörr, W., 2004. The

1298 Problem of Dating High-pressure Metamorphism: a U–Pb Isotope and Geochemical Study on

1299 Eclogites and Related Rocks of the Mariánské Lázně Complex, Czech Republic. J. Petrol. 45,

1300 1311–1338.

1301 Timmermann, H., Dörr, W., Krenn, E., Finger, F., Zulauf, G., 2006. Conventional and in situ

1302 geochronology of the Teplá Crystalline unit, Bohemian Massif: implications for the processes

1303 involving monazite formation. Int. J. Earth Sci. 95, 629–647.

1304 Tomek, F., Žák, J., Chadima, M., 2015. Granitic magma emplacement and deformation

1305 during early-orogenic syn-convergent transtension: The Staré Sedlo complex, Bohemian

1306 Massif. J. Geodyn. 87, 50–66.

1307 Trubač, J., Žák, J., Chlupáčová, M., Janoušek, V., 2009. Magnetic fabric of the Říčany

1308 granite, Bohemian Massif: a record of helical magma flow? J. Volcanol. Geoth. Res. 181, 25–

1309 34.

1310 Trubač, J., Janoušek, V., Gerdes, A., 2019. Petrogenesis of fractionated nested granite

1311 intrusions: the Sedmihoří Composite Stock (Bohemian Massif). J. Geosci. 64, 271–294.
Turner, S., Hawkesworth, C., Rogers, N., Bartlett, J., Worthington, T., Hergt, J., Pearce, J., Smith, I., 1997. 238U-230Th disequilibria, magma petrogenesis, and flux rates beneath the depleted Tonga-Kermadec island arc. Geochim. Cosmochim. Acta 61, 4855-4884.

Turner, S., Bourdon, B., Hawkesworth, C., Evans, P., 2000. 226Ra-230Th evidence for multiple dehydration events, rapid melt ascent and the time scales of differentiation beneath the Tonga-Kermadec island arc. Earth. Planet. Sci. Lett. 179, 581-593.

Turner, S., Bourdon, B., Gill, J., 2003. Insights into Magma Genesis at Convergent Margins from U-series Isotopes. Rev. Mineral. Geochem. 52, 255-315.

Vacek, F., Žák, J., 2019. A lifetime of the Variscan orogenic plateau from uplift to collapse as recorded by the Prague Basin, Bohemian Massif. Geol. Mag. 156, 485-509.

Venera, Z., Schulmann, K., Kröner, A., 2000. Intrusion within a trans-tensional tectonic domain: the Čistá granodiorite (Bohemian Massif)—structure and rheological modelling. J. Struct. Geol. 22, 1437-1454.

Vogt, K., Gerya, T. V., 2014. From oceanic plateaus to allochthonous terranes: Numerical modelling. Gondwana Res. 25, 494-508.

Willbold, M., Stracke, A., 2010. Formation of enriched mantle components by recycling of upper and lower continental crust. Chem. Geol. 276, 188-197.

Wood, D. A., 1980. The application of a ThHfTa diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary Volcanic Province. Earth. Planet. Sci. Lett. 50, 11-30.

Workman, R. K., Hart, S. R., 2005. Major and trace element composition of the depleted MORB mantle (DMM). Earth. Planet. Sci. Lett. 231, 53-72.

Žák, J., Schulmann, K., Hrouda, F., 2005a. Multiple magmatic fabrics in the Sázava pluton (Bohemian Massif, Czech Republic): a result of superposition of wrench-dominated regional transpression on final emplacement. J. Struct. Geol. 27, 805-822.
1337 Žák, J., Holub, F. V., Verner, K., 2005b. Tectonic evolution of a continental magmatic arc from transpression in the upper crust to exhumation of mid-crustal orogenic root recorded by episodically emplaced plutons: the Central Bohemian Plutonic Complex (Bohemian Massif). Int. J. Earth Sci. 94, 385–400.

1341 Žák, J., Kratinová, Z., Trubač, J., Janoušek, V., Sláma, J., Mrlina, J., 2011. Structure, emplacement, and tectonic setting of Late Devonian granitoid plutons in the Teplá–Barrandian unit, Bohemian Massif. Int. J. Earth Sci. 100, 1477–1495.

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1353 Zulauf, G., 2001. Structural style, deformation mechanisms and paleodifferential stress along an exposed crustal section : constraints on the rheology of quartzofeldspathic rocks at supra- and infrastructural levels (Bohemian Massif). Tectonophysics, 332, 211–237.

1356
1357 **FIGURE CAPTION**

1358 Fig. 1. (a) Sketch of the Variscan belt (modified from Martínez Catalán et al., 2020) with
1359 location of arc-type rocks within the Variscan belt. (b) Simplified geological map of the
1360 Teplá-Barrandian Unit. Published geochronological and $\mathcal{E}_{Nd}$ data shown in the figure were
1361 selected to highlight and constrain the magmatism during the Devonian to Carboniferous
1362 Variscan evolution. The map was redrafted from the 1:500000 geological map of the Czech
1363 Republic, published by the Czech Geological Survey in 2007. Additional structural elements
1364 are from Chlupáčova et al. (1975), Zulauf (2001) and Peřestý et al. (2017). Source of
1365 geochronological and isotopic data: (1) Dörr, Zulauf (2010); (2) Janoušek et al. (2010); (3)
1366 Košler et al. (1993a); (4) Janoušek, Gerdes (2003); (5) Holub et al. (1997b); (6) Janoušek et
1367 al. (2004); (7) Janoušek et al. (1997); (8) Venera et al. (2000); (9) Žák et al. (2011); (10)
1368 Siebel et al. (1999); (11) Ackerman et al. (2010); (12) Janoušek et al. (2000); (13) Košler
1369 (1993b); (14) Trubač et al. (2019). (c) A-B interpretative cross-section through the
1370 Saxothuringian-MLC-Teplá-Barrandian-Moldanubian units showing tectonic styles related to
1371 Devonian and Carboniferous events affecting the Teplá-Barrandian Unit.

1372 Fig. 2. (a) Geological map of the Mariánské Lázně Complex (modified after Timmermann et
1373 al., 2004 and Collett et al., 2018). (b) Schematic NW-SE cross-sections across the MLC (A-
1374 B) presenting the MLC and TBU architecture after D3 detachment, showing relics of vertical
1375 D2 fabrics in the eclogite-amphibolite and gabbro-amphibolite belts and the deformation
1376 associated to D3 extensional detachment deformation juxtaposing the deep MLC to the
1377 hangingwall TBU crust (after Peřestý et al., 2017, 2020; Collett et al., 2018). Schematic
1378 position of magmatic rocks with respect to the structural succession and relations to
1379 surrounding rock types is based on this study. (c) Geological map of the Čistá pluton
1380 (modified after Venera et al., 2000 and Žák et al., 2011). (d) Schematic NW-SE cross-
sections across the Čistá pluton (C-D) presenting the structure of the Čistá pluton intruding the Late Cambrian Tis granite and weakly metamorphosed metasediments of the TBU. The D3 west-dipping detachment developed along the eastern part of the Čistá pluton (after Venera *et al.*, 2000).

Fig. 3. Photographs showing field relations of individual rock types. (a) Coarse-grained gabbro with domains of variable proportions of amphibole and plagioclase. Boundaries between leucocratic and mesocratic domains are diffuse, suggesting magmatic origin. (b) Trondhjemite with partly sharp and partly diffuse boundary towards a coarse grained gabbro, suggesting magmatic relationships. (c) Tonalite within an eclogite retrogressed to amphibolite. Both lithologies have strong S2 foliation. (d) Outcrop-scale foliation in a gabbro that is parallel with S3 regional foliation. (e, f) lithological contact of a coarse-grained gabbro and trondhjemite with an eclogite retrogressed to a foliated amphibolite. These relations are interpreted as gabbro intruding the eclogite during amphibolite-facies deformation. Detail of contact is shown in (f).

Fig. 4. Photographs of representative hand specimen samples used in this study, sorted in three major groups: (a–e) MLC deformed and undeformed mafic to intermediate rocks, (f–h) MLC felsic rocks of acidic composition and (i,f) Čistá plutonic rocks represented by biotite granitoid. Indicated Types are related to geochemical/petrological characteristics as described in the results.

Fig. 5. Microphotographs of two representative samples showing rock textures and microstructures. (a) Equi-micro-granular texture and random grain orientation typical of HT magmatic textural equilibration (ML44-2). Arrow indicates amphibole profile shown as inset. (b) triple points (highlighted by red arrows) in amphibole indicating static re-equilibration. (c) hetero-micro-granular magmatic texture, with random grain orientation (ML46A). (d) detail of plagioclase grain showing zonation – underlined by white dotted line. (e) amphibole core-
rim compositions following classification scheme of Hawthorne *et al.* (2012) (ML44-2, ML46A). (f) core-rim anorthite content of plagioclase (ML44-2, ML46A).

Fig. 6. Geochemical classification of the studied rocks in (a) the total alkali vs. silica (TAS) diagram (after Middlemost, 1994) and (b) the AFM diagram (after Irvine and Baragar, 1971).

The C–O metagabbro and retrogressed eclogite-amphibolite of the MLC are also shown for comparison (also used in Figs. 7 & 8; data in Supplementary Data, Table S5, and from Beard *et al.*, 1995; Crowley *et al.*, 2002; Timmermann *et al.*, 2004; Jašarová *et al.*, 2016).

Fig. 7. Harker variation diagrams for the MLC and Čistá magmatic rocks. The C–O metagabbro and retrogressed eclogite-amphibolite of the MLC are also shown for comparison.

Fig. 8. Primitive mantle-normalized multiple trace element diagrams and chondrite-normalized REE patterns for the MLC and Čistá magmatic rocks. Three types are distinguished for the gabbroic to quartz-gabbroic rocks (Type-1 to 3) and three types for the felsic rocks in the MLC (Type-4 to 6, dotted line are gneiss). The Čistá magmatic rocks show relatively homogeneous chemistry. Normalization values after McDonough and Sun (1995) for the primitive mantle and after Anders and Grevesse (1989) for the chondrite. For Types 1–3, data of the MLC retrogressed eclogite-amphibolite and the MLC C–O metagabbro are shown for comparison.

Fig. 9. Representative cathodoluminescence images of zircon from dated samples of the MLC mafic and felsic rocks, and the Čistá pluton, sorted by the rock type and sample. The ²⁰⁶Pb/²³⁸U dates are shown in million years (Ma) and with 2σ in-run uncertainty.

Fig. 10. Probability density plots of ²⁰⁶Pb/²³⁸U dates and Tera-Wasserburg concordia diagrams with 2σ error ellipses for MLC mafic rocks (a, b, c), MLC felsic rocks (d, e, f, g) and Čistá pluton (h, i, j, k). Color for each group of samples (purple-pink-red) corresponds to the color used in geochemical diagrams. The different shades of color are attributed to the
1431 main populations of ages corresponding to all the samples identified in the probability density
1432 plots (300–440 Ma; 440–520 Ma and 520–600 Ma). Only the magmatic rocks are reported in the Tera-Wasserburg concordia diagrams (ages are approximated using the TuffZirc algorithm).

1435 Fig. 11. (a) Sr-Nd whole rock isotope diagram for the MLC magmatic rocks (initial
1436 composition recalculated at t₀=386 Ma) and for the Čistá pluton (initial composition
1437 recalculated at t₀=370 Ma). (b) εNd vs Sr isotopic ratio variations in the MLC magmatic rocks
1438 (t₀=386 Ma) and in the Čistá pluton (t₀=370 Ma). Each mark on the mixing curves is a 2%
1439 contribution of either GLOSS (GLObal Subducting Sediments) or LC (Lower crust) to the
1440 DMM (Depleted MORB Mantle) composition. The composition of the DMM, GLOSS, LC
1441 end members are from Workman and Hart (2005); Plank and Langmuir (1998) and Willbold
1442 and Stracke (2010) respectively and recalculated at t₀ = 386 Ma. Other MLC isotopic data
1443 from Beard et al. (1995) and Timmermann et al. (2004) and from the Blovice Accretionary
1444 Complex (Drost et al., 2007) are also shown for comparison.

1445 Fig. 12. (a) Trace element discrimination diagrams for the tectonic setting of granitic rocks
1446 after Pearce et al. (1984). Abbreviations: VAG, Volcanic Arc Granites; syn-COLG, syn-
1447 Collisional Granites; ORG, Ocean Ridge Granites; WPG, Within Plate Granites. (b) Trace
1448 element discrimination diagrams for the tectonic setting of basaltic rocks after Wood (1980).
1449 Abbreviations: CAB, Calc-Alkaline Basalts; IAT, Intra-Arc Tholeites; N-MORB, Normal
1450 Mid-Oceanic Ridge Basalts; E-MORB, Enriched Mid-Oceanic Ridge Basalts; WPT, Within
1451 Plate Basalts; WPA, Within Plate Alkaline basalts. When samples are under detection limit
1452 for one used element in the diagrams, the symbol is color filled, and the concentration in that
1453 element is fixed to its detection limit. Black arrows indicate the location of these samples with
1454 a fixed concentration at half the detection limit.
Fig. 13. (a) Nb/Yb vs. Th/Yb discrimination diagram after Pearce (2008). E- and N-MORB=Enriched and Normal Mid-Oceanic Ridge Basalts; OIB=Oceanic island basalts; SZ=Subduction-derived fluids/melts. Addition of fluids and/or melts derived from the slab to a melt chemically similar to the MORB-OIB array will bring the melt towards the volcanic arc array. Note the evolution of the MLC rocks, from the MORB-OIB array for the more mafic rocks to the volcanic arc array for the more felsic ones. (b) Nb/Yb vs. TiO₂/Yb plot (after Pearce, 2008) indicating the depth of source melting process. When samples are under detection limit for one used element in the diagrams, the symbol is color filled and plotted at the detection limit with an arrow indicating the location of these samples at a fixed concentration of half the detection limit.

Fig. 14. (a) Primitive mantle-normalized multiple trace element diagrams showing ML44-1 composition compared to the modeled dry DMM and metasomatized DMM by slab-derived fluids (Temperature of 1000 °C, melting rate of 20% with 5% contribution of the subduction derived fluid). (b) Sr/Th vs. Th and U/Th vs. Th diagrams for MLC gabbroic samples Type 1 to Type 3 (note the increasing contribution of the slab melting component in the system). When samples are under detection limit for one used element in the diagrams, the symbol is color filled and the concentration in that element is fixed to its detection limit.

Fig. 15. (a) Primitive mantle-normalized multiple trace element diagrams showing modeling results of ML44-1 and MLC mean gabbro composition 10 to 100 % of fractional crystallization (FC, modeled using Petromodeler excel sheet; Ersoy, 2013) compared to the chemistry of the felsic types. (b-d) Primitive mantle-normalized multiple trace element diagrams showing dehydration melting modeling (batch melting) results of different amphibole bearing rocks present in the MLC (eclogite-amphibolite, arc amphibole gabbro, C-O metagabbro) compared to (b) the ML14-14-1 Type-4 sample, (c) the oligoclasite Type-5 samples and (d) the trondhjemites Type-6 samples. Parameters were set to T=750 °C for all
1480 the modelling, 20 % of partial melting for Type-4 modelling and 3% of partial melting for

1481 Types-5 and -6 modelling. (e) Chondrite-normalized REE concentrations of the modeled

1482 felsic melt and restite composition resulting from the arc gabbros dehydration melting model,

1483 compared to the MLC plagiogranite and Kohistan garnet rich hornblende pyroxenite (Garrido
et al., 2010) (T=950 °C and 20% of partial melting).

1485 Fig. 16. Four-stage geodynamic evolution of the Paleozoic active margin from 390 Ma to 340

1486 Ma based on geodynamic scenario of Peřestý et al. (2017): (a) rollback of the Saxothuringian

1487 subduction zone associated with early emplacement of the deep arc chamber beneath the

1488 accretionary wedge and the TBU upper plate; (b) regional thickening of accreted wedge and

1489 upper plate provoked by entry of continental crust into the subduction zone associated with

1490 massive emplacement of arc magmas along S2 anisotropy; (c) development of major east

1491 dipping detachment affecting the MLC accretionary-arc edifice and emplacement of the Čistá

1492 pluton along antithetic west dipping shear zone due to vertical rise of previously subducted

1493 crustal bulge; (d) relamination and exhumation of granulitized crust related to subduction of

1494 remnant basins and arrival of main Saxothuringian continent. This stage is related to

1495 formation of Central Bohemian Plutonic Complex magmatic arc partly digesting earlier

1496 Devonian arc edifice, (e) chronological chart showing tempos of two magmatic arcs, their

1497 εNd signatures and two main compression-extension tectonic cycles related to their

1498 emplacements.

1499
<figure><img src="image_3.png" /></figure>

(b) Mariánské Lázně Complex
Teplá-Barrandian Domain

Serpentinite
A

Eclogite-amphibolite belt
D3

Gabbro-amphibolite belt
D3

<figure><img src="image_4.png" /></figure>

<figure><img src="image_5.png" /></figure>

<figure><img src="image_6.png" /></figure>

<figure><img src="image_7.png" /></figure>

<figure><img src="image_8.png" /></figure>

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

<figure><img src="image_11.png" /></figure>

<figure><img src="image_12.png" /></figure>

Eclogite/amphibolite
(MLC)

<figure><img src="image_13.png" /></figure>

Cambro-Ordovician
metagabbro (MLC)

<figure><img src="image_14.png" /></figure>

Mafic granulite/
amphibolite (MLC)

<figure><img src="image_15.png" /></figure>

Devonian intrusive rocks
(mostly gabbro)
with magmatic texture/
with solid-state foliation

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

Čistá granitoid

<figure><img src="image_18.png" /></figure>

Carboniferous
cover

<figure><img src="image_19.png" /></figure>

Aplite, syenite and
granitoid porphyrie

<figure><img src="image_20.png" /></figure>

Neoproterozoic

<figure><img src="image_21.png" /></figure>

Fault

<figure><img src="image_22.png" /></figure>

Magmatic foliation

<figure><img src="image_23.png" /></figure>

Sample location

<figure><img src="image_24.png" /></figure>

Tis granite; covered by sediments

<figure><img src="image_25.png" /></figure>

<figure><img src="image_26.png" /></figure>

S1-S2

<figure><img src="image_27.png" /></figure>

<figure><img src="image_28.png" /></figure>

S3
<figure><img src="image_29.png" /></figure>

<figure><img src="image_30.png" /></figure>

<figure><img src="image_31.png" /></figure>

<figure><img src="image_32.png" /></figure>

<figure><img src="image_33.png" /></figure>

<figure><img src="image_34.png" /></figure>
<figure><img src="image_35.png" /><figcaption>(a)</figcaption></figure>

ML44-2 - Type 1

1 cm

<figure><img src="image_36.png" /><figcaption>(c)</figcaption></figure>

ML46A
Type 2

1 cm

<figure><img src="image_37.png" /><figcaption>(d)</figcaption></figure>

ML33A
Type 2

MLC mafic rocks

(b) ML31C - Type 2

<figure><img src="image_38.png" /></figure>

1 cm

<figure><img src="image_39.png" /><figcaption>(e)</figcaption></figure>

ML31A
Type 3

1 cm

MLC felsic rocks

<figure><img src="image_40.png" /><figcaption>(f)</figcaption></figure>

ML14-14-1
Type 4

1 cm

<figure><img src="image_41.png" /><figcaption>(g)</figcaption></figure>

ML28F
Type 5
1 cm

<figure><img src="image_42.png" /><figcaption>(h)</figcaption></figure>

ML31B - Type 6

1 cm

Čistá pluton

(i) Cl4E1

<figure><img src="image_43.png" /><figcaption>(j)</figcaption></figure>

<figure><img src="image_44.png" /></figure>

Cl6A
1 cm
<figure><img src="image_45.png" /></figure>

<figure><img src="image_46.png" /></figure>

<figure><img src="image_47.png" /></figure>

<figure><img src="image_48.png" /></figure>

<figure><img src="image_49.png" /></figure>

<figure><img src="image_50.png" /></figure>
<figure><img src="image_51.png" /></figure>

<figure><img src="image_52.png" /></figure>

<table><thead><tr><th></th><th>MLC mafic rocks</th></tr></thead><tbody><tr><td>○ ML44-1</td><td>ML20-14-1</td></tr><tr><td>△ ML44-2</td><td>ML34C</td></tr><tr><td>⁺ ML31C</td><td>ML33E</td></tr><tr><td>× ML46A</td><td>ML33D</td></tr><tr><td>◇ ML34B</td><td>ML33A</td></tr><tr><td>▽ ML27A4</td><td>ML31A</td></tr><tr><td>☒ ML20-14-4</td><td>ML33C</td></tr></tbody></table>

<table><thead><tr><th></th><th>MLC felsic rocks</th></tr></thead><tbody><tr><td>○ ML14-14-1</td><td>ML28E</td></tr><tr><td>△ ML28F</td><td>ML33B2</td></tr><tr><td>⁺ ML30C</td><td>ML30E</td></tr><tr><td>× ML30</td><td>ML31B</td></tr><tr><td>◇ ML14-14-6</td><td>ML30H</td></tr><tr><td>▽ ML28D</td><td></td></tr></tbody></table>

<table><thead><tr><th></th><th>Čistá pluton</th></tr></thead><tbody><tr><td>○ CI4E1</td><td>CI6A</td></tr><tr><td>△ CI3A</td><td>CI5D1</td></tr><tr><td>⁺ CI4D</td><td>CI5A</td></tr><tr><td colspan="2">Others</td></tr><tr><td>◇ C.O. metagabbro</td><td></td></tr><tr><td>☒ MLC eclogite/amphibolite</td><td></td></tr></tbody></table>
<figure><img src="image_53.png" /></figure>

<figure><img src="image_54.png" /></figure>

<figure><img src="image_55.png" /></figure>

<figure><img src="image_56.png" /></figure>

<figure><img src="image_57.png" /></figure>

<figure><img src="image_58.png" /></figure>

<figure><img src="image_59.png" /></figure>

<figure><img src="image_60.png" /></figure>

<figure><img src="image_61.png" /></figure>

<figure><img src="image_62.png" /></figure>

<figure><img src="image_63.png" /></figure>

<figure><img src="image_64.png" /></figure>
Primitive mantle normalized

<figure><img src="image_65.png" /></figure>

<figure><img src="image_66.png" /></figure>

<figure><img src="image_67.png" /></figure>

Chondrite normalized

<figure><img src="image_68.png" /></figure>

<figure><img src="image_69.png" /></figure>

<figure><img src="image_70.png" /></figure>
Primitive mantle normalized

<figure><img src="image_71.png" /></figure>

<figure><img src="image_72.png" /></figure>

<figure><img src="image_73.png" /></figure>

<figure><img src="image_74.png" /></figure>

Chondrite normalized

<figure><img src="image_75.png" /></figure>

<figure><img src="image_76.png" /></figure>

<figure><img src="image_77.png" /></figure>

<figure><img src="image_78.png" /></figure>
MLC mafic rocks

<figure><img src="image_79.png" /></figure>

<figure><img src="image_80.png" /></figure>

<figure><img src="image_81.png" /></figure>

<figure><img src="image_82.png" /></figure>

MLC felsic rocks

<figure><img src="image_83.png" /></figure>

<figure><img src="image_84.png" /></figure>

<figure><img src="image_85.png" /></figure>

<figure><img src="image_86.png" /></figure>

<figure><img src="image_87.png" /></figure>

<figure><img src="image_88.png" /></figure>

Čistá pluton

<figure><img src="image_89.png" /></figure>

<figure><img src="image_90.png" /></figure>

<figure><img src="image_91.png" /></figure>

<figure><img src="image_92.png" /></figure>

<figure><img src="image_93.png" /></figure>
MLC mafic rocks

<figure><img src="image_94.png" /></figure>

MLC felsic rocks

<figure><img src="image_95.png" /></figure>

Čistá pluton

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(a)

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<table><thead><tr><th colspan="4">MLC mafic rocks</th></tr></thead><tbody><tr><td>▽</td><td>Type-1</td><td>◇</td><td>Type-3</td></tr><tr><td>☒</td><td>Type-2</td><td></td><td></td></tr></tbody></table>

<table><thead><tr><th colspan="4">MLC felsic rocks</th></tr></thead><tbody><tr><td>△</td><td>Type-4</td><td>◇</td><td>Type-6</td></tr><tr><td>+</td><td>Type-5</td><td></td><td></td></tr></tbody></table>

<table><thead><tr><th colspan="2">Čistá pluton</th></tr></thead><tbody><tr><td>☒</td><td>Granitoid</td></tr></tbody></table>

(b)

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<figure><img src="image_125.png" /><figcaption>(b) Mid- to Late-Devonian, entry of continent in subduction, wedge and arc compression</figcaption></figure>

<figure><img src="image_126.png" /><figcaption>(c) Late-Devonian – Early Carboniferous, arc magma in extension zones, exhumation of continental bulge</figcaption></figure>

<figure><img src="image_127.png" /><figcaption>(d) Early- to Mid-Carboniferous, CBPC arc magma emplacement in compression zones, relamination of lower plate and its extrusion into upper plate</figcaption></figure>

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● MLC mafic rocks

● MLC felsic rocks

● Čistá pluton

SZ : Subduction zone enrichment

C : Crustal contamination

F : Fractionnal crystallisation

W : within plate enrichment

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Continental crust
(future relaminant)

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Oceanic crust/
mafic lower crust

<figure><img src="image_137.png" /></figure>

Subduction
wedge

<figure><img src="image_138.png" /></figure>

Devonian arc
magmatic rocks

<figure><img src="image_139.png" /></figure>

Lower orogenic
crust

<figure><img src="image_140.png" /></figure>

Teplá-Barrandian
Unit

<figure><img src="image_141.png" /></figure>

Carboniferous arc
magmatic rocks

(a) Early- to Mid-Devonian, retreating subduction zone, stacking of eclogite wedge, invasion of arc magma in subduction wedge

<figure><img src="image_142.png" /></figure>

(b) Mid- to Late-Devonian, entry of continent in subduction, wedge and arc compression

<figure><img src="image_143.png" /></figure>

(c) Late-Devonian – Early Carboniferous, arc magma in extension zones, exhumation of continental bulge

<figure><img src="image_144.png" /></figure>

(d) Early- to Mid-Carboniferous, CBPC arc magma emplacement in compression zones, relamination of lower plate and its extrusion into upper plate

<figure><img src="image_145.png" /></figure>