# Middle Devonian, late Carboniferous, and Triassic magmatic flare-ups in eastern Armorica (Sakarya Zone, Turkey) as revealed by detrital zircon U-Pb-Hf isotopic data

Gültekin Topuz, Osman Candan, Oscar Laurent, Ali Mohammadi, Cengiz Okuyucu, Ömer Faruk Çelik, Jia-Min Wang

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Gültekin Topuz, Osman Candan, Oscar Laurent, Ali Mohammadi, Cengiz Okuyucu, et al., Middle Devonian, late Carboniferous, and Triassic magmatic flare-ups in eastern Armorica (Sakarya Zone, Turkey) as revealed by detrital zircon U-Pb-Hf isotopic data. Geological Society of America Bulletin, 2024, 136 (7-8), pp.3053-3068. 10.1130/B36950.1. hal-04793406

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1 Middle Devonian, Late Carboniferous and Triassic magmatic flare-ups in the

2 Eastern Armorica (Sakarya Zone, Turkey) as revealed by detrital zircon U-

3 Pb-Hf isotopic data

4

5 Gültekin Topuz¹*, Osman Candan², Oscar Laurent³,⁴, Ali Mohammadi¹, Cengiz Okuyucu⁵,

6 Ömer Faruk Çelik⁶, Jia-Min Wang⁷

7

8 ¹ İstanbul Teknik Üniversitesi, Avrasya Yer Bilimleri Enstitüsü, TR34469 Maslak, İstanbul,

9 Turkey; E-mail: topuzg@itu.edu.tr (https://orcid.org/0000-0001-8690-4614);

10 mohammadiali@itu.edu.tr

11 ² Dokuz Eylül Üniversitesi, Jeoloji Mühendisliği Bölümü, Tınaztepe Yerleşkesi, Buca, İzmir,

12 Turkey; E-mail: osman.candan@deu.edu.tr

13 ³ ETH Zurich, Department of Earth Sciences, Institute for Geochemistry and Petrology,

14 Clausiusstrasse 25, 8092, Zurich, Switzerland

15 ⁴ CNRS, Géosciences Environnement Toulouse, Observatoire Midi-Pyrénées, 14 avenue Edouard

16 Belin F-31400 Toulouse, France. E-mail: oscar.laurent@get.omp.eu

17 ⁵ Konya Teknik Üniversitesi, Jeoloji Mühendisliği Bölümü, 42250 Selçuklu, Konya, Turkey. E-

18 mail: cokuyucu@ktun.edu.tr

19 ⁶ Kocaeli Üniversitesi, Jeoloji Mühendisliği Bölümü, Umuttepe Yerleşkesi, TR41001 Kocaeli,

20 Turkey; E-mail: celikfrk@gmail.com

21 ⁷ State Key Laboratory of Lithospheric Evolution, Institute of Geology and Geophysics, Chinese

22 Academy of Sciences, 19 Beitucheng West Road, Chaoyang District, 100029 Beijing, China, E-

23 mail: wangjm9595@gmail.com
24 ABSTRACT

The Sakarya Zone (northern Turkey) contains a well-preserved Early-Middle Jurassic and Late Cretaceous submarine magmatic arc constructed over pre-Jurassic bedrocks that are considered to be the eastward extension of the Armorican Terrane Assemblage in Europe. In this study, we present U-Pb-Hf isotopic data on the detrital zircons from Middle Permian and Lower Jurassic sandstones to reveal episodes of Paleozoic–Early Mesozoic magmatic flare-ups. Detrital zircon ages together with data from the literature define three major age groups at 400–380, 326–310 and 250–230 Ma, indicating three distinct magmatic flare-ups. In addition, there are minor age clusters at 460–430 and 215–195 Ma. Initial εHf values of the detrital zircons indicate a significant juvenile input during the Triassic flare-up, the involvement of significant reworked crustal material during the Late Carboniferous magmatic flare-up, and both juvenile and reworked crustal material during the Middle Devonian magmatic flare-up. Within the pre-Jurassic continental basement rocks of the Sakarya Zone, the Late Carboniferous igneous rocks are well documented and most voluminous, and the Middle Devonian rocks are known locally, while the Triassic igneous rocks apart from those in Triassic accretionary complexes are hardly known. Because the Sakarya Zone is a Gondwana-derived continental block that was later involved in the Variscan and Alpine orogenies, these magmatic flare-ups cannot be explained by subduction-related processes along a single subduction zone. We propose that the Sakarya Zone rifted from the northern margin of Gondwana during Late Ordovician-Silurian, the Devonian magmatic flare-up (400–380 Ma) was related to the southward subduction of the Rheic ocean beneath the Sakarya Zone during its northward drift and the Late Carboniferous magmatic flare-up (326–310) occurred following the collision of the Sakarya Zone with Laurussia, and the Triassic one (250–230 Ma) resulted from northward subduction of the Tethys ocean beneath the
Sakarya Zone. Comparison with literature data shows that the Triassic and Late Carboniferous magmatic flare-ups are also characteristic features of neighboring Armorican domains such as the Balkans and the Caucasus; however, the Middle Devonian flare-up appears to be restricted to the Sakarya Zone. Along with the Late Carboniferous flare-up, the Late Ordovician-Silurian flare-up which is locally recorded in the Sakarya Zone is typical of the Armorican Terrane Assemblage as a whole.

**Key-Words:** Sandstone, provenance, U-Pb zircon age, zircon Hf isotopy, magmatic flare-up, Sakarya Zone, Armorica, Turkey

55

## INTRODUCTION

In orogenic belts, the intensity of magmatic activity shows temporal and spatial variations. In general, episodes of high-volume magmatism alternate with those of low-volume or no magmatism (e.g., DeCelles et al., 2009; Kirsch et al., 2016; Paterson and Ducea, 2015; Zhang et al., 2020; Chapman et al., 2021). Episodes of high-volume magmatism are referred to as magmatic flare-ups and episodes of low-volume magmatism as magmatic lulls. In general, magmatic flare-ups last 5 to 20 Ma, while magmatic lulls occur over 30 to 70 Ma. Although it is difficult to determine the exact volume of magma produced during flare-up events, estimated production rates range from 140 to 215 km³/km·Ma during flare-ups, and less than 15 km³/km·Ma during lulls (Chapman et al., 2021). The proposed drivers for such magmatic flare-up events are multiple, including (i) changes in convergence rate, (ii) slab roll-back, (iii) slab breakoff, slab tear and lithospheric delamination, (iv) repeated magmatic underplating of the lower crust, and (v) lithospheric thickening in the fore-arc and back-arc domains (e.g. Ducea and Barton, 2007; Paterson and Ducea, 2015; Shellnutt et al., 2014; Ardila et al., 2019; Zhang et al., 2019).
Episodes of the magmatic flare-ups and lulls are best constrained by combining the relative abundance of detrital zircon ages from clastic sedimentary rocks with the ages of exposed igneous rocks and their estimated volumes (e.g. Paterson and Ducea, 2015; Zhang et al., 2019; Chapman et al., 2021). In several orogenic belts such as the Sierra Nevada arc in California and the Neo-Tethyan arc system in south Asia (Paterson and Ducea, 2015; Zhang et al., 2019; Balgord et al., 2021), detrital zircon age clusters match well the timing of high-volume magmatism, except for the case of dominant basic magmatism. Although clusters of the detrital zircon ages are affected by various factors such as sampling bias and relative abundance of zircons in different rock types (e.g. Moecher and Samson, 2006; Dickinson, 2008; Malusa et al., 2013; Anderson et al., 2019), in ancient and/or long-lived orogenic belts, detrital zircon ages are the primary means of revealing magmatic flare-ups (e.g., Balgord et al., 2021; Pastor-Galan et al., 2021), as the igneous rocks related to former flare-ups may be largely eroded, metamorphosed, or buried under younger volcanic-sedimentary strata.

The Sakarya Zone (northern Turkey) contains a well-preserved Early-Middle Jurassic and Late Cretaceous submarine magmatic arc that was constructed over pre-Jurassic bedrock mostly buried under Jurassic and younger volcanic-sedimentary strata (Figs. 1 and 2) (Şen, 2007; Okay and Nikishin, 2015). Pre-Jurassic basement outcrops account for only about 10% of the total outcrops and probably represent the eastern extension of the Armorican Terrane Assemblage in central Europe (e.g. Winchester and the Pace TMR Network Team, 2002; Okay et al., 2006; Okay and Topuz, 2017). The Armorican Terrane Assemblage or simply Armorica sensu Franke includes Armorican Massif, Vosges, Black Forest and Bohemian Massif. In this study, we present U-Pb-Hf isotope data on detrital zircons from Middle Permian sandstones in a Permo-Triassic accretionary complex and from Lower Jurassic sandstones unconformably overlying the pre-Jurassic basement to reveal the pre-Jurassic magmatic flare-ups. The U-Pb zircon age clusters
together with literature data define three major flare-up events at 400–380 (latest Early to earliest Late Devonian), 326–310 (latest Early to Late Carboniferous) and 250–230 Ma (Early to earliest Late Triassic; after Cohen et al. 2013 updated). For the sake of brevity, the flare-ups at 400–380, 326–310 and 250–230 Ma are hereinafter referred to as Middle Devonian, Late Carboniferous and Triassic, respectively. Furthermore, we present an evolutionary model for the Paleozoic and early Mesozoic evolution of the Sakarya Zone beginning with the rifting from the northern margin of Gondwana during Late Ordovician-Silurian, continuing with northward drift and accretion into Laurussia during the Early Carboniferous and ending with the Permo-Triassic subduction of the Paleo-Tethys beneath the Laurasia.

## GEOLOGICAL FRAMEWORK

The Sakarya Zone (northern Turkey) is a Gondwana-derived continental fragment bound by the Intra-Pontide suture and the Black Sea to the north and the Izmir-Ankara-Erzincan suture to the south (Fig. 2). Its pre-Jurassic bedrock can be divided into two parts: (i) a continental domain mainly of Carboniferous to Middle Permian age, and (ii) Permo-Triassic sedimentary and low-grade metamorphic rocks interpreted as accretionary complexes, commonly known as the Karakaya Complex (Figs. 2 and 3). The continental domain consists of several units such as (i) Early Carboniferous high-temperature / middle- to low-pressure (HT-M to LP) metamorphic rocks intruded by various types of Early Carboniferous to Early Permian acidic to basic intrusions (Nzegge et al., 2006; Topuz et al., 2004b, 2010, 2023; Dokuz, 2011; Kaygusuz et al., 2012, 2016; Ustaömer et al., 2012, 2013; Karslı et al., 2016), (ii) Upper Carboniferous to Lower Permian sedimentary rocks (Okay and Leven, 1996; Dokuz et al., 2017), and (iii) Early to Middle Devonian non-metamorphic granites intruding undated low-grade metamorphic rocks (Okay et
al., 2006; Aysal et al., 2012; Sunal, 2012). The Early Carboniferous HT-M to LP metamorphic units comprise metasedimentary rocks mainly of Late Neoproterozoic to Early Paleozoic protolith ages and acidic to ultrabasic meta-igneous rocks with Ordovician to Early Carboniferous igneous crystallization ages (e.g., Dokuz et al., 2022; Karslı et al., 2020; Topuz et al., 2020). The Permo-Triassic accretionary complexes include (i) greenschist to blueschist- and epidote-amphibolite-facies metamorphic rock assemblages, consisting of metabasite, marble, phyllite and minor metachert and serpentinite (the Lower Karakaya Complex), and (ii) non-metamorphic to very low-grade sandstone, basalt and limestone (the Upper Karakaya Complex) (Pickett and Robertson, 1996; Okay et al., 2002, 2020; Okay and Göncüoğlu, 2004; Topuz et al., 2004a, 2014, 2018; Robertson and Ustaömer, 2012). Ar-Ar phengite dating on the Lower Karakaya Complex gave metamorphic ages ranging from 262 to 195 Ma (Middle Permian to earliest Jurassic) (Okay and Monié, 1997; Okay et al., 2002, 2020; Topuz et al., 2004a, 2014, 2017). The metaophiolite fragments record Early to Middle Permian igneous crystallization ages (274–263 Ma) (Topuz et al., 2018; Okay et al., 2020). The metabasic and basaltic rocks in both the Lower and Upper Karakaya complexes exhibit mainly anorogenic alkaline affinities similar to those found in seamounts, suggesting seamount accretion during the Late Triassic (Pickett and Robertson, 1996; Sayit and Göncüoğlu, 2009). In places, the Permo-Triassic accretionary complexes are tectonically juxtaposed with the Early to Middle Jurassic accretionary complexes, ophiolites and Late Cretaceous ophiolitic mélanges. This suggests a history of subduction and accretion from Middle Permian to Late Cretaceous time.

All the pre-Jurassic basement units are overlain with a major unconformity by Early to Late Jurassic sedimentary and volcaniclastic rocks and locally with Ammonitico Rosso limestone horizons in the eastern to central part of the Sakarya Zone (Fig. 2; Altıner et al., 1991; Şen, 2007;
Kandemir and Yılmaz, 2009; Vincent et al., 2018). Exposures of the Jurassic volcaniclastic rocks in the western part of the Sakaraya Zone are minor (Genç and Tüysüz, 2010). The volcaniclastic rocks formed in a time interval from the Pliensbachian to the latest Oxfordian-early Kimmeridgian (183–158 Ma; Cohen et al. 2013 updated), and exhibit lateral facies and thickness variations. These features led to the interpretation that they formed in a rifting environment (Şen, 2007, and the references therein). The volcaniclastic rocks pass up into neritic carbonates which are Late Jurassic to Early Cretaceous in age (Altıner et al., 1991; Vincent et al., 2018). The basal levels of the volcaniclastic sequence occupy a crucial position in determining the major pre-Jurassic magmatic flare-ups, as they overlay various bedrocks with a major unconformity.

In clear contrast to the Sakarya Zone, the Anatolide-Tauride Block is characterized by Paleozoic to Late Mesozoic sedimentary rocks (mainly carbonates) with local unconformities and lacks any Early Carboniferous HT-M to LP metamorphism and deformation (e.g. Özgül and Kozlu, 2002; Özgül, 1983; Robertson et al., 2021; Fig. S1). Paleozoic magmatism is minor and local. Nevertheless, Triassic magmatism is more widespread. The Triassic unconformities in the Anatolide-Tauride Block as well as the lateral facies and thickness changes in the Jurassic volcaniclastic rocks of the Sakarya Zones have been attributed to a rifting event that resulted in the opening of the Neo-Tethyan ocean (e.g. Şengör and Yılmaz, 1981; Şengör et al. 2019). However, several lines of evidence such as (i) the presence of the Permo-Triassic, Jurassic, and Late Cretaceous accretionary complexes in the Sakarya Zone, (ii) blocks or slivers of Permian ophiolite as well as Middle Devonian, Carboniferous, Permian and Triassic cherts in the Permo-Triassic accretionary complexes, and (iii) Middle Triassic to Upper Cretaceous radiolarites in the Late Cretaceous accretionary complexes conclusively suggest that the Izmir-Ankara-Erzincan suture represents the trace of a long-lived oceanic domain from at least Silurian to Late
164 Cretaceous (e.g. Pickett and Robertson, 1996; Tekin et al., 2002; Robertson and Ustaömer, 2012; Okay et al. 2013, 2020; Topuz et al. 2013a, 2014, 2017).

166

167 THE TOKAT MASSIF

168 The Tokat Massif is a large metamorphic area (ca. 18 000 km²) to the southwest of the
169 Eastern Pontides, and consists of Permo-Triassic to Late Jurassic low-grade to non-metamorphic
170 rocks tectonically interleaved with several east-west trending sheets of Late Cretaceous ophiolitic
171 mélange, 0.2 to 2 km across and up to 60 km long (Figs. 2 and 4). The low-grade metamorphic
172 rocks comprise transitional greenschist- to blueschist-facies metabasite, marble, phyllite and
173 minor metachert and serpentinite (e.g., Rojay and Göncüoğlu, 1997; Yılmaz et al., 1997; Yılmaz
174 and Yılmaz, 2004; Eyüboğlu et al., 2011, 2018; Catlos et al., 2013, Topuz et al., 2022). Within
175 the metamorphic rocks, there are Triassic anorogenic tholeiitic to alkaline metagabbros as well as
176 Early Devonian and Permian limestone olistoliths in predominantly Triassic metaclastic rocks
177 (Çapkınoğlu and Bektaş, 1997; Eyüboğlu et al., 2018). The anorogenic tholeiitic to alkaline
178 metagabbros with igneous crystallization ages of 245–203 Ma are dominant meta-igneous
179 lithologies in the Tokat Massif (Eyüboğlu et al., 2011, 2018). On the other hand, the Late
180 Cretaceous ophiolitic mélanges comprise basalt, gabbro, radiolarite, serpentinite, sandstone,
181 shale, limestone, trondhjemite and metamorphic rocks (e.g., Bozkurt et al., 2007; Çelik et al.
182 2019a-b). The Tokat Massif is interpreted either as a Triassic marginal basin that developed on
183 the northern margin of Gondwana (Şengör and Yılmaz, 1981; Şengör et al., 2019; Eyüboğlu et al.
184 2018) or as a Permo-Triassic accretionary complex related northward subduction of Paleo-Tethys
185 beneath the Laurasia (Okay and Göncüoğlu, 2004; Topuz et al., 2014).
186

187

THE MIDDLE PERMIAN AND LOWER JURASSIC SANDSTONES

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206

In the northern part of the Tokat Massif, there is a dismembered formation composed of blocks/olistoliths of Lower to Middle Permian neritic limestone in a clastic matrix, herein referred to as the Middle Permian blocky unit (Figs. 4, S2a,b). Limestone olistoliths range in size from tens of meters to kilometers in diameter. Similar blocky units are known from the Upper Karakaya Complex in the western part of the Sakarya Zone (Pickett and Robertson, 1996; Okay and Göncüoğlu, 2004; Robertson and Ustaömer, 2012). The matrix in all these blocky units of the Upper Karakaya Complex in the western part of the Sakarya Zone is Late Triassic in age, as can be inferred from the ages of the youngest detrital zircons (Ustaömer et al., 2016). However, all available detrital zircon ages from the matrix of the blocky unit are ≥ 294 Ma (earliest Permian, n=360, this study). Hence, we have named this unit the Middle Permian blocky unit. This blocky unit may be a discrete older part of the Upper Karakaya Complex. The largest exposure of this unit, approximately 30 km long and up to 5 km wide, is found ca. 20 km to the north of Amasya. Another exposure, ca. 10 km long and ca. 3 km wide, occurs around the town of Ladik. To the north of the North Anatolian right-lateral strike-slip fault, there are also two discrete outcrops of this blocky unit. Our descriptions here relate to two outcrops south of the North Anatolian Fault. The base of the unit is not exposed. The internal stratigraphy of the blocky unit is obscure (Fig. S2a). Layering is locally recognizable only in limestone. In the limestones there are clastic interlayers up to 20 cm thick. The Middle Permian blocky unit is unconformably overlain by either Lower Jurassic clastic rocks or Ammonitico-Rosso facies limestones (Fig. 3).

207

208

The Jurassic clastic rocks unconformably overlie the low-grade Permo-Triassic metamorphic rocks of the Tokat Massif, and grade upward into Upper Jurassic and Lower
209 Cretaceous limestones (Figs. 3 and 4). The Early to Late Jurassic volcanic and volcanoclastic rocks are confined to the Jurassic clastic successions located north of the Jurassic magmatic front (Fig. 4). The total thickness of the clastic series south of Amasya is over 100 m. The lowermost horizons (up to 15 m thick) consist of conglomerate grading upwards to sandstone and minor siltstone. Individual layers are 20 to 80 cm thick. The clasts in the conglomerates range in diameter from 2 to 20 cm and are composed mainly of chert, phyllite and metabasite. To the northwest of Göynücek, the Lower Jurassic sandstones (ca. 60 m. thick) unconformably overlie the retrogressed blueschist-facies rocks, and grade conformably upward to Upper Jurassic-Lower Cretaceous chert-bearing, thin-bedded limestone.

218

219 **PETROGRAPHY**

220 In this study, we have studied detrital zircons from six sandstone samples collected over an area of about 1100 km². Three of them come from the Middle Permian sandstones (T290A, T320 and T356B), and three from the Lower Jurassic sandstones (samples T239, T267 and T295) (Figs. 3 and 4; Table S1). The Lower Jurassic sandstone samples were taken from the lowermost horizons of the Lower Jurassic clastic sequence. The stratigraphic position of the Middle Permian sandstones is difficult to deduce because the internal structure is unclear, and the base of the Middle Permian blocky sequence cannot be observed.

221

222

223

224

225

226

227

228 **The Middle Permian Sandstones**

229 *Sample T290A* is a poorly sorted lithic arenite (matrix < 10 vol.%), comprising quartz (50-60 vol.%) and minor K-feldspar (5 vol.%) and plagioclase (5 vol.%), and lithic fragments (ca. 30
vol.%; Fig. S3a). Lithic fragments consist mainly of fossiliferous limestone and minor amounts of andesite and granite. Metamorphic rock fragments are rare. Myrmekitic and micrographic microstructures are preserved in granite fragments. Quartz grains are predominantly polycrystalline. The grains have an angular to subangular shape, and the grain size ranges from 50 to 300 µm, which indicates the poorly sorted nature of sandstone. All these features suggest poor textural and mineralogical maturity. *Sample T320* is a poorly sorted, matrix-supported (ca. 30 vol.% matrix) lithic greywacke, comprising quartz (ca. 30 vol.%) and minor biotite (partially chloritized), muscovite, K-feldspar, plagioclase and lithic fragments (ca. 30 vol.%; Fig. S3b). Lithic fragments are predominantly limestone and minor andesite and granite. Metamorphic rock fragments are minor. The grains mostly have an angular to subangular shape. Therefore, the textural and mineralogical maturity is poor. *Sample T356B* is from an about 10 cm thick loose sand interlayer within a Middle Permian limestone block. Thus, the sandy interlayer is interpreted as contemporaneous with the limestone, in marked difference to samples T290A and T320, which represent the surrounding matrix of limestone olistoliths.

245

246 The Lower Jurassic Sandstones

247 *Sample T239* is a poorly sorted Jurassic lithic arenite (matrix < 15 vol%), comprising quartz (15-20 vol%), calcite (15-20 vol.%), and rock fragments (50-60 vol.%; Fig. S3c). Lithic fragments were derived predominantly from limestone and subordinately phyllite, mica-schist, andesite, granophyre and granite. Lithic fragments tend to be more rounded than quartz grains.
248
249
250
*Sample T267* is a poorly sorted lithic arenite similar to sample T239 (Fig. S3d). The grains are dominated by lithic fragments (ca. 60 vol%), which are represented by limestone, greenschist, phyllite/mica schist, calc phyllite, serpentinite, andesite, granite and polycrystalline quartz.
251
252
253
254 Metamorphic rock fragments are more common than igneous ones. Quartz (15-20 vol.%),
255 epidote, calcite, biotite and muscovite occur as discrete minerals. *Sample T295* is a poorly sorted
256 lithic greywacke (matrix ca. 20 vol%), comprising lithic fragments (ca. 35 vol.%), quartz (ca. 20
257 vol.%), biotite, muscovite, K-feldspar and plagioclase in a calcite cement/matrix (ca. 30 vol.%;
258 Figs. S3e-f). Lithic fragments are from phyllite/mica schist, andesite, granophyre, granite and
259 serpentinite. Overall, the grains have an angular to subangular shape.

260 In summary, both the Middle Permian and Lower Jurassic sandstones are poorly sorted
261 and lack textural and mineralogical maturity. The lithic fragments ranging from granitic to
262 granophyric rocks, and the abundance of biotite flakes particularly in the Lower Jurassic
263 sandstones suggest that a major part of the grains came from a continental area, as granitic to
264 granophyric rocks are missing in the Tokat Massif. Metamorphic rock fragments are more
265 common in the Lower Jurassic sandstones.

266

267 ANALYTICAL TECHNIQUES

268 Detrital zircons were extracted from sandstone samples weighing approximately 3 to 5 kg
269 using standard mineral separation techniques such as cracking, sieving, washing, drying,
270 magnetic and heavy mineral separation at the Eurasian Institute of Earth Sciences (Istanbul
271 Technical University). Handpicked zircons under a binocular microscope were mounted in epoxy
272 resin and polished down to roughly half thickness. Polished mounts were then photographed
273 using a scanning electron microscope equipped with a cathode luminescence (CL) detector at the
274 Geological Engineering Department of Hacettepe University in Ankara. Cathode luminescence
(CL) and backscattered electron images helped visualize the internal structures of the zircon grains and the presence of inclusions prior to isotopic analysis.

U-Pb analyses and trace element compositions of zircon were determined simultaneously by laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS). The analyses were performed at the Institute of Geochemistry and Petrology, ETH Zurich (Eidgenössische Technische Hochschule Zürich, ETHZ), Switzerland, using a RESOlution (ASI/Applied Spectra) excimer (ArF) 193 nm laser ablation system attached to an Element XR (Thermo) sector-field ICP-MS. Lu-Hf analyses were obtained by laser ablation-multi-collection-inductively coupled plasma-mass spectrometry (LA-MC-ICP-MS) at ETHZ, using a RESOlution (ASI/Applied Spectra) excimer (ArF) 193 nm laser ablation system attached to a Nu Plasma II (Nu Instruments) multi-collector ICP-MS. Details about the instrumentation and analytical conditions can be found in the Appendix and in Tables S2 and S3. Data from unknowns and reference materials are presented in Tables S4 and S5. Concordia and frequency probability plots were produced using ISOPLOT v.4.15 (Ludwig, 2003). The concordance of the ages was calculated using the ratio of $^{206}$Pb/$^{238}$U and $^{207}$Pb/$^{235}$U ages. In this study, only the analyses with concordance between 90–110% are considered. In the histograms, $^{206}$Pb/$^{238}$U ages are used for zircons with ages <1 Ga, while $^{207}$Pb/$^{206}$Pb ages are used for zircons with ages >1 Ga. The probability curve shows the mean peak age of the age group in a sample.

293

## 294 RESULTS

295 Morphology and Internal Structure
The zircons in the studied samples are characterized by (i) grain sizes ranging from 50 to 296 200 µm in the longitudinal direction, (ii) colors ranging from clear colorless, yellowish to violet, 297 and (iii) external shapes ranging from euhedral prismatic to wholly rounded (Fig. S4). Euhedral 298 prismatic zircon grains account for less than 15 % of the dated detrital zircon grains. The vast 299 majority of zircon grains (ca. 90 %) show well-developed oscillatory zoning along with sector 300 zoning. Metamict zircons are not recognized. Some grains appear to have inherited core-like 301 domains. Feldspar and apatite inclusions are common in zircon grains. The marginal domains of 302 the zircon grains were preferred for the measurements. 303

## Trace Element Compositions 304

A total of 720 spot analyses have been performed on the detrital zircons from six 305 sandstone samples. In each sample, 120 spots were analysed (Table S4) to maximize 306 representativeness of the sample population (Vermeesch, 2004). A total of 616 spot analyses 307 yielded concordant U-Pb ages at 90–110 % [(²⁰⁷Pb/²³⁵U age)/(²⁰⁶Pb/²³⁸U age) *100]. We only 308 consider the age values and trace element compositions of concordant spot analyses. 309

The zircon grains are mainly characterized by a steep heavy rare earth element-enriched 310 (REE) pattern, a positive Ce anomaly, and a negative Eu anomaly (Figs. S5a-f). There is no 311 apparent relationship between the age of the detrital zircons and REE characteristics, such as the 312 magnitude of the negative Eu anomaly (not shown). Out of 616 concordant analyses, only twelve 313 zircon grains show REE patterns with heavy REE depletion and varying degrees of negative Eu 314 anomaly, a feature of zircons equilibrated with garnet (Fig. S5g) (e.g., Rubatto, 2002). Some 315 zircon grains (16 grains) exhibit a steep REE pattern with minor Eu anomaly, suggesting growth 316 from magmas that did not experience significant plagioclase fractionation (Fig. S5h). 317
318 U-Pb Zircon Ages

319 In the Middle Permian sandstone samples, the $^{206}$Pb/$^{238}$U ages of the detrital zircon grains 320 range from 294 to 2913 Ma (earliest Permian to Mesoarchean after Cohen et al., 2013 updated) 321 (Figs. 5a-c; Table S4). The youngest detrital zircon population in all three samples is earliest 322 Permian in age (299–294 Ma) consistent with the paleontological dates of the limestone 323 olistoliths. In all Middle Permian sandstone samples, the $^{206}$Pb/$^{238}$U ages of the detrital zircon 324 grains define marked age clusters peaking at 318, 324 and 325 Ma (Late Carboniferous). The 325 main age population in samples T290A and T320A (53 and 57% of all concordant zircons) is 326 Neoproterozoic to latest Mesoproterozoic with noticeable clusters at 562–555, 634–629 and ca. 327 1025 Ma (Figs. 5a-b). In contrast, zircons older than Paleozoic are subordinate (ca. 25%) in 328 sample T356B, where the dominant age peak is at 392 Ma (Middle Devonian) (Fig. 5c).

329 In the Lower Jurassic sandstone samples, the age of the detrital zircons ranges from 195 330 to 2698 Ma (Early Jurassic to Neoarchean) (Figs. 5d-f; Table S4). All samples consistently 331 display noticeable age clusters at 250–233 and 318–316 Ma. In addition, there are marked age 332 clusters peaking at 380 and 383 Ma in samples T239 and T267, respectively. Only sample T239 333 has an additional age peak located at 441 Ma. The Paleozoic to early Mesozoic zircons make up 334 66% of all the concordant zircon grains in sample T295, and 37 and 44% in samples T239 and 335 T267, respectively. Similar to the Middle Permian sandstones, the Lower Jurassic sandstones 336 have age clusters at 581, 650–604 and 969–1027 Ma (late Neoproterozoic to latest 337 Mesoproterozoic). Samples T239 and T267 are further characterized by non-negligible mid- 338 Paleoproterozoic to Neoarchean populations (ca. 15-29% of all concordant zircons; Figs. 5d-e). 339 In sample T295, only 6 out of 105 grains have $^{207}$Pb/$^{206}$Pb ages >1500 Ma (Fig. 5f).
In summary, the detrital zircon age spectra from both the Lower Jurassic and Middle Permian sandstones contain variable amounts of Phanerozoic zircons that define distinct age clusters in the range of 250-233, 325-316 and 392-380 Ma. The combination of all detrital zircon age data suggests that the most pronounced clusters are at 325-316 Ma (mainly Late Carboniferous), followed by 250-230 Ma (mainly Early to Middle Triassic). The age cluster at 393-380 Ma (mainly Middle Devonian) is the least pronounced one, and a subdued age peak at ca. 441 Ma (Early Silurian) is present only in one sample.

## Lu-Hf Isotopic Compositions

Lu-Hf isotopic compositions of the 478 detrital zircon grains from six samples were measured (Table S5). In each sample, at least 70 spots on the zircon grains with concordant U-Pb ages were analyzed. The results are shown on $^{206}$Pb/$^{238}$U zircon ages (Ma) versus time-corrected $\epsilon$Hf values in Figures 6a-b.

The initial $\epsilon$Hf values of the dated zircon grains show a large variation from -26 to +14. Although detrital zircons with Triassic ages exhibit initial $\epsilon$Hf values from -3 to +10, the negative values make up only 12.5% of this age population. On the other hand, Carboniferous zircons have initial $\epsilon$Hf values in the range of -12 to +13, where negative initial $\epsilon$Hf values make up ca. 94 % and positive values are only shown by the zircons with ages $\le$ 314 Ma. Middle Devonian zircons display initial $\epsilon$Hf values ranging from -9 to +7 with negative initial $\epsilon$Hf values accounting for 76 %. Silurian zircons are characterized by initial $\epsilon$Hf values of -5 to +8 with positive initial $\epsilon$Hf values making up ca. 60%. It is noteworthy that the detrital zircons from the Middle Permian and Lower Jurassic sandstones show different ranges of $\epsilon$Hf values for a given age for the Paleozoic era. For example, Devonian zircons from the Middle Permian sandstones are characterized mainly by negative initial $\epsilon$Hf values, and those from Lower Jurassic
sandstones mainly by positive εHf values. The same trend is also observed in the latest 363
Carboniferous zircons. In addition, one Permian sample shows a very clear population at ca. 550 364
Ma with positive εHf (around +10), which is lacking in the Lower Jurassic samples. Archean and 365
Proterozoic zircons from both Middle Permian and Lower Jurassic sandstones, on the other hand, 366
show the same range of initial εHf. 367

To sum up, initial εHf values of the detrital zircons vary widely from positive to negative 368
in both the Middle Permian and Lower Jurassic sandstones. The Triassic and Silurian zircons are 369
characterized primarily by positive initial εHf values, while the Late Carboniferous and Middle 370
Devonian zircons are characterized primarily by negative initial εHf values. The detrital zircons 371
from the Middle Permian and Lower Jurassic sandstones show different ranges of initial εHf 372
values for the same age range during the Paleozoic era. 373

374

375 DISCUSSION

376 Geochemical Constraints on Possible Sources of Detrital Zircons

Th/U ratios of the dated zircon grains range from 0.01 to 1.54 (Figs. S6a-b; Table S4). 377
Only 16 zircon grains among 616 concordant analyses (ca. 2.6 %) have Th/U ratios less than 378
0.10. Zircons with Th/U ratios <0.10 are concentrated at 495–467 (late Cambrian-Ordovician) 379
and 573–745 Ma (middle to late Neoproterozoic). Commonly, zircons with a Th/U ratio <0.10 380
are considered to be metamorphic in origin, and zircons with a Th/U ratio >0.10 are igneous in 381
origin (e.g., Rubatto, 2002; Hoskin and Schaltegger, 2003). Based on the compilation of the 382
metamorphic zircon compositions and phase equilibrium modeling in metapelitic systems, 383
Yakymchuk et al. (2018), however, showed that the Th/U ratio of suprasolidus metamorphic 384
zircon is controlled by (i) U and Th concentrations in the system, (ii) the amount of monazite, and 385
386 (iii) the timing of zircon crystallization. Metamorphic zircons that crystallized at or near wet
387 solidus in the presence of monazite are predicted to have Th/U ratios < 0.1, while metamorphic
388 zircons that crystallized under open-system conditions at elevated temperatures in the absence of
389 monazite are predicted to have Th/U ratios > 0.1. To summarize, all zircon grains with Th/U
390 ratios <0.1 are probably of metamorphic origin grown in the presence of monazite, but not all
391 metamorphic zircons have Th/U ratios <0.1.

392 U/Yb ratios of the zircon grains range from 0.1 to 47.2, with average values between 0.8
393 and 1.3 in different samples (Figs. S6c-d). Overall, zircons grown in magmas from enriched
394 mantle and crustal sources mostly show elevated U/Yb ratios (ca. > 0.1) compared to those from
395 normal mid-ocean ridge magmas (Grimes et al., 2015). Nb/Yb ratios range from 0.001 to 0.483,
396 with average values between 0.008 and 0.013 (Figs S6e-f). Thus, the majority of the dated zircon
397 grains plot above the mantle array on the Nb/Yb versus U/Yb diagram after Grimes et al. (2015)
398 (Fig. 7a). Forty zircon grains out of 616 concordant grains (7%) fall within the field of mantle
399 zircon array which is defined by the zircons of mid-ocean ridges, ocean islands and plume-
400 influenced settings. Only eight zircon grains out of 40 yielded Phanerozoic ages of 225-211 and
401 433-428 Ma (Late Triassic and Silurian). Thus, the overwhelming majority of dated zircon
402 grains are compositionally similar to those from convergent margin granitoids. On the Hf versus
403 Ti diagram after Laurent et al. (2022) (Fig. 7b), the zircons fall within the fields defined by the
404 zircons from magmatic arc to post-orogenic I-type granites.

405 Based on the above considerations and the internal structures of dated zircon grains (Fig.
406 S4), we conclude that the dated zircon grains, regardless of their age, were predominantly derived
407 from convergent margin igneous rocks, and only a small fraction might be of metamorphic and
408 mid-ocean ridge and ocean island basaltic origin (ca. < 10 %).
409 Paleozoic-Early Mesozoic Magmatic Flare-ups

410 The combination of all detrital zircon ages from the Middle Permian and Lower Jurassic
411 sandstones defines three major age peaks at 240, 321 and 391 Ma, and smaller age peaks at 212,
412 430 and 475 Ma (Figs. 8a-b). Comparable age peaks (238, 312 and 393 Ma) are delineated by the
413 detrital zircons of (meta-)sandstones, collected over an area of 28 000 km², within the Late
414 Triassic accretionary complexes in the western part of the Sakarya Zone, ca. 600 km west of the
415 Tokat Massif (Figs. 7c-d; Ustaömer et al., 2016). A limited set of detrital zircon ages from Lower
416 to Upper Jurassic clastic rocks in the Gümüşhane area, ca. 300 km east of Tokat Massif,
417 uniformly define age clusters at 204–199 and 320–310 Ma (Figs. 2 and S7a-b) (Akdoğan et al.,
418 2018; Liu et al., 2022). Although the Carboniferous clusters entirely overlap with those in this
419 study and in Late Triassic accretionary complexes from the western part of the Sakarya Zone
420 (Ustaömer et al., 2016), the Triassic zircon age peak occurs at 204–199 Ma (latest Triassic-
421 earliest Jurassic) rather than at 250–230 Ma (Early to Late Triassic), which is comparable to the
422 minor peak in sample T239 (Fig. 5d). Furthermore, there is no Middle Devonian age peak.
423 However, detrital zircons from an Early Carboniferous low-grade metapsammite from the
424 Demirözü area, ca. 40 km to the southeast of Gümüşhane, displays a pronounced Middle
425 Devonian age cluster peaking at 393 Ma (sample BP36 in Dokuz et al., 2022) (Fig. S7c). Detrital
426 zircon U-Pb zircon ages from the central part of the Sakarya Zone and Crimea show some
427 marked differences from the eastern and western parts of the Sakarya Zone such as (i) the
428 presence of a pronounced Early Permian age peak (291–289 Ma) (Figs. S7d-e), (ii) a significant
429 abundance of Paleoproterozoic and Meso- to Neoarchean zircons, and (iii) the sporadic
430 occurrence of the zircons in the age range 1100 and 1700 Ma (Okay et al., 2013; Akdoğan et
431 al., 2017; Kuznetsov et al., 2019; Nikishin et al., 2020). These differences are also reflected in the
age data of the detrital zircons from Lower Cretaceous turbidites in the region which have been
attributed to a source in the East European Craton before the opening of the Black Sea during the
Late Cretaceous (Okay et al., 2013; Akdoğan et al. 2017, 2019). Liu et al. (2022), on the other
hand, suggest that there was no exchange of detritus across the Eastern Black Sea basin during
the Middle Jurassic-early Eocene, and the Eastern Black Sea basin was probably rifted by the
Middle Jurassic as a back-arc basin. The source of Early Permian zircons can be ascribed to the
local granites of Early Permian age in the Central Pontides (e.g. Nzegge et al., 2006). The other
age clusters are comparable to those defined by the detrital zircons from the Middle Permian and
Lower Jurassic clastic rocks of the Tokat Massif.

To sum up, available detrital zircon U-Pb age data from the Permian, Late Triassic and
Jurassic clastic rocks from the Sakarya Zone consistently define major age clusters at 250-230,
326-310 and 400-380 Ma, as well as less pronounced clusters at 215-195, 455-430 and
480-470 Ma. The presence of these major age clusters is also supported by the detrital zircon
ages from Upper Cretaceous and Middle Eocene clastic rocks throughout the Sakarya Zone (Figs.
8e-f, S7) (e.g. Okay et al., 2013; Akdoğan et al., 2017, 2019; Mueller et al., 2019, 2022; Liu et
al., 2022). These age clusters correspond to the magmatic flare-ups in the Sakarya Zone. Given
the uncertainties in the age data, these major magmatic flare-ups lasted ca. 15 to 20 Ma and were
separated by magmatic lulls lasting ca. 15 to 60 Ma.

## Possible Provenance of Early Mesozoic-Paleozoic Zircons

There are no well-documented and confirmed Triassic magmatic arc rocks in the pre-
Jurassic basement outcrops in the Sakarya Zone, despite the common presence of Triassic detrital
zircons in the Late Triassic accretionary complexes and in the Lower Jurassic sandstones (Okay
and Nikishin, 2015, Ustaömer et al., 2016, this study). The only reported Triassic igneous rocks
are anorogenic alkaline to tholeiitic (meta-)gabbros in the Late Triassic accretionary complexes and Late Cretaceous ophiolitic mélanges (e.g. Eyüboğlu et al., 2011, 2018; Çelik et al., 2019b). We, however, consider it unlikely that the Triassic zircons in our data set were derived from Triassic gabbroic rocks within the Permo-Triassic accretionary complexes, on the basis of zircon-poor nature of the gabbroic rocks and trace element compositional features of the dated zircons (Fig. 7). Based on the presence of Triassic igneous rocks from the bore-hole data from the Scythian platform, Okay and Nikishin (2015) proposed that the Triassic magmatic arc lies buried under the Cenozoic cover of the Scythian platform (Fig. 2), which was adjacent to the Pontides before the Late Cretaceous opening of the Black Sea. The Middle to Late Permian time (264–252 Ma) represents a short phase of magmatism in the Istanbul Zone (Bozkurt et al., 2013; Aysal et al., 2018; Babaoğlu et al., 2023) (Fig. 2). This magmatism is somewhat older than the postulated Triassic magmatic flare-up (250–230 Ma). Hence, we regard it improbable that the Triassic zircons in the Latest Triassic accretionary complexes and Lower Jurassic sandstones originated from the Istanbul Zone during Early Jurassic time. In conclusion, the derivation of the Early to Middle Triassic zircons from the northern range prior to the opening of the Black Sea Basin seems to be a conceivable possibility, as suggested by Okay and Nikishin (2015).

The Late Carboniferous magmatic flare-up (326–310 Ma) represents the most voluminous and relatively well-documented one. Magmatic rocks showing similar ages as this flare-up are batholiths and stocks throughout the Sakarya Zone and Caucasus, comprising mainly middle- to high-K calc-alkaline I-type granites (Fig. 2; e.g., Topuz et al., 2010; Ustaömer T. et al., 2013; Dokuz, 2011; Ustaömer P.A. et al., 2012; Kaygusuz et al., 2012, 2016; Rolland, 2017). Late Carboniferous gabbroic rocks (326–322 Ma) are minor, and represented by relatively small mafic to ultramafic intrusions, and were derived from middle to high-K calc-alkaline basic magmas
(Topuz et al., 2023). The magmatic flare-up event started ca. 5–8 Ma after the peak of the regional HT-M to LP metamorphism. In some provenance studies, the Carboniferous age cluster is located at 334–330 Ma (e.g., Mueller et al., 2019; Kuznetsov et al., 2019), which corresponds to the peak of the HT-M to LP metamorphism throughout the Sakarya Zone and Caucasus and NW Iran (Figs. 2 and 3) (Topuz et al., 2004b; Mayringer et al., 2011; Ustaömer T. et al., 2013; Rolland et al., 2016; Roland, 2017; Moazzen et al., 2020). The initial εHf values of both the Late Carboniferous detrital zircons and the Late Carboniferous granites entirely overlap (Fig. 6b). This similarity is also reinforced by fully negative values of the initial εNd values of the Late Carboniferous granites.

In contrast to the Late Carboniferous igneous rocks, the Middle Devonian igneous rocks are documented locally from the Sakarya Zone (Okay et al., 2006; Aysal et al., 2011; Sunal, 2011; Topuz et al., 2020; Karslı et al., 2020; Dokuz et al., 2022). In recent years their number has steadily increased with the dating of metaigneous rocks within the Early Carboniferous HT-M to LP metamorphic units. The Middle Devonian rocks are represented by biotite ± hornblende- bearing and muscovite ± biotite granites with U-Pb zircon igneous crystallization ages of 404–380 Ma. The relatively pronounced age clusters at 400–380 Ma in the detrital zircon age histograms (Figs. 8a-d) suggest that the Early to earliest Late Devonian igneous rocks must be more common than documented. The predominantly negative values of the initial εHf values of the Middle Devonian detrital zircons (ca. 76%) are consistent with the widespread occurrence of the Middle Devonian muscovite ± biotite granites. The Devonian zircons in the Middle Permian sandstones have predominantly negative initial εHf values, while those in the Lower Jurassic sandstones have predominantly positive initial εHf values (Fig. 6b). This quite contrasted isotopic signature suggests that the Middle Permian and Lower Jurassic sandstones do not have the same
provenance and/or that the Middle Devonian magmatism was characterized by bimodal processes
involving either reworking of older crust or new growth from depleted mantle-like sources. In
any case, there were predominantly juvenile magmatic rocks in the catchment area when the
Lower Jurassic sandstone was formed.

The Late Ordovician-Silurian detrital zircon age peak is apparent only in some
provenance studies (Figs. 8b and 8f and S7d-f) (Okay et al., 2013; Akdoğan et al., 2017;
Kuznetsov et al., 2019; Mueller et al., 2022, this study). So far, Silurian igneous rocks are known
from the western part of the Sakarya Zone, where they are represented mainly by anorogenic
alkaline and tholeiitic amphibolite and subordinately metagranite with igneous crystallization
ages of 440-420 Ma (Topuz et al., 2020; Karslı et al., 2020). The basic rocks were derived
mainly from juvenile material, while acidic rocks were derived from magmas which involved
substantial reworked material (Fig. 6b). Local appearance of the Silurian age peak can be a bias
mainly due to the basic nature of the Late Ordovician-Silurian igneous rocks.

Detrital zircon ages from the Middle Permian and Lower Jurassic sandstones define a
minor subdued peak at 475 Ma (Early Ordovician). Similar age clusters are also found in the
detrital zircon age data of Akdoğan et al. (2017) and Mueller et al. (2022). To date, Ordovician
igneous rocks have been reported only from the Armutlu Peninsula in the western Sakarya Zone
(Okay et al., 2008) and from the Kurtoğlu metamorphic area in the eastern Sakarya Zone (Dokuz
et al., 2022). Based on Th/U ratio ≤ 0.1, some detrital zircon grains with Early Ordovician ages
are thought to have a possible metamorphic origin, which is consistent with the data presented in
Ustaömer et al. (2013). However, Early Ordovician metamorphism in the Sakarya Zone is still
unknown.

## Possible Causes of the Paleozoic-Early Mesozoic Flare-ups
There are only a few studies dealing with the Paleozoic to Early Mesozoic evolution of the Sakarya Zone (e.g. Okay et al., 2006; Okay and Topuz 2017; Topuz et al., 2020; Ustaömer et al. 2012, 2013; Dokuz et al., 2017, 2022; Eyüboğlu et al., 2010, 2018; Karslı et al., 2020). There is a general consensus that the Sakarya region is a block of Gondwana origin, based on the widespread presence of Neoproterozoic zircons in pre-Jurassic metaclastic rocks (Ustaömer et al., 2012, 2013). However, the timing of rifting from Gondwana in the south, and accretion to Laurussia in the north are highly debated. Several studies (e.g., Şengör and Yılmaz, 1981; Şengör et al., 2019; van Hinsbergen et al., 2020) suggest that the Sakarya Zone first separated from the northern margin of Gondwana during Permo-Triassic time. On the other hand, Karslı et al. (2020) and Dokuz et al. (2022) argue that the pre-Jurassic basement of the Sakarya Zone contains domains that detached from the northern margin of Gondwana at different times such as the Ordovician and the Late Devonian, and the pre-Jurassic basement of the Sakarya Zone contains domains derived both Avalonia and Armorican Terrane Assemblage. Both proposals are considered unlikely based on the following arguments. As mentioned above, the Izmir-Ankara-Erzincan suture represents the trace of a long-lived oceanic domain at least from the Silurian to Late Cretaceous on the basis of the age of the accretionary complexes and deep-marine sedimentary blocks therein (e.g. Okay et al., 2013, 2020; Topuz et al., 2013). Moreover, the Middle Paleozoic-Early Mesozoic stratigraphy of the Sakarya Zone and the Anatolide-Tauride Block, the Arabian Platform, differ significantly (Fig. S1). The latter proposal suggests an opening of Paleo-Tethys to the north of the Rheic Ocean, which contradicts the areal distribution of the Paleo-Tethyan accretionary complexes in the Sakarya Zone and with configurations of the oceanic domains in the commonly accepted paleogeographic reconstructions (e.g. Stampfli and Borell, 2002). So far, no area has been identified that has a large number of detrital zircons with
547 ages ranging from 1100 to 1750 Ma inside the basement units of the Sakarya Zone, that could
548 represent part of Avalonia (Figs. 8 and S7).

549 Below we discuss the possible drivers of Paleozoic to early Mesozoic magmatic flare-ups
550 in the Sakarya Zone in the context of the following constraints: (i) The Sakarya Zone is a
551 Gondwanian continental fragment, and forms the eastern extension of the Armorican Terrane
552 Assemblage (Winchester and the Pace TMR Network Team, 2002; Okay et al., 2006; Ustaömer
553 et al., 2012; Okay and Topuz, 2017), (ii) the Sakarya Zone rifted from the northern margin of
554 Gondwana probably during the Late Ordovician-Silurian, and accreted to Laurussia during the
555 Carboniferous (e.g., Ustaömer et al., 2013; Okay and Topuz, 2017; Topuz et al., 2020), and (iii)
556 the limited Paleozoic and early Mesozoic igneous rocks in the Anatolide-Tauride Block are ruled
557 out as potential source since the Izmir-Ankara-Erzincan suture represents the trace of an oceanic
558 domain which existed from the middle Paleozoic to the end-Mesozoic (e.g., Topuz et al., 2013;
559 Okay et al., 2020). Therefore, the Paleozoic to early Mesozoic magmatic flare-ups cannot be
560 explained by processes along a single subduction zone. The processes leading to detachment
561 from the northern margin of Gondwana, northward drift, and accretion to the southern margin of
562 Laurussia should be considered. As a reference point, we use the Late Carboniferous-Permian
563 paleogeographic reconstruction at the end of the Variscan Orogeny, which consistently shows an
564 easterly widening oceanic embayment (Fig. 9a) (e.g. Stampfli and Borell, 2002; von Raumer et
565 al. 2013). As the continental collision is succeeded by large-scale strike-slip faulting and the
566 formation of intracontinental basins in Central Europe (e.g. Corsini and Rolland, 2009), the
567 Sakarya and Istanbul zones were located at the southern margin of Laurasia. Below, we discuss
568 the possible causes of the magmatic flare-ups from the youngest one to the oldest one (Figs. 9b-
569 g).
Compelling evidence for the existence of subduction during Permo-Triassic times comes from the widespread Permo-Triassic accretionary complexes including Triassic eclogites, blueschists and Permian ophiolites (e.g. Pickett and Robertson, 1996; Okay and Monie, 1997; Okay et al., 2002, 2020; Okay and Göncüoğlu, 2004; Topuz et al., 2004b, 2014, 2018; Robertson and Ustaömer, 2011). Thus, the Triassic magmatic flare-ups at 250–230 and 215–198 Ma are related to the northward subduction of Paleo-Tethys beneath the Sakarya Zone (Figs. 9b-c). The widespread presence of anorogenic alkaline to tholeiitic metagabbroic rocks in the Late Triassic accretionary complexes point to the accretion of a seamount chain or an oceanic plateau (Pickett and Robertson, 1996; Sayit and Göncüoğlu, 2009). Consequently, the latest Triassic-earliest Jurassic minor flare-up might be related to slab steepening following the accretion of a seamount chain or an oceanic plateau. This situation would lead to the migration of a magmatic front from the north to the south.

Late Carboniferous magmatic flare-up occurred 5-8 Ma after the peak of HT-M to LP metamorphism in the Sakarya Zone, and took place during the exhumation of HT-M to LP metamorphic rocks in an extensional setting (Topuz et al., 2023). The drastic change in U-Pb zircon age spectra of the Istanbul Zone (Avalonia) as of Carboniferous time provides conclusive evidence that the accretion of the Sakarya Zone to Laurussia during Early Carboniferous by the consumption of the Rheic ocean (Okay and Topuz, 2017; Akdoğan et al., 2022). The Late Carboniferous magmatic flare-up is tentatively ascribed to the removal of the lowermost part of the lithosphere by convective thinning (Topuz et al., 2022) (Figs. 9d-e). Recent discoveries of Early Carboniferous high-pressure metamorphic rocks in northwestern Iran indicate that the southern margin of the Sakarya Zone and the Caucasus was active during Early Carboniferous (e.g. Rosetti et al., 2017; Wan et al., 2021) (Fig. 2). However, it is unclear whether this
subduction was short-lived or continuous from the Early Carboniferous to the Middle Permian. It
is difficult to assess the relative contributions of active subduction and accretion to Late
Carboniferous magmatic flare-up. Seeing that the HT-M to LP metamorphism throughout the
Sakarya Zone, the Caucasus and NW Iran occurred at 335-330 Ma (e.g. Topuz et al., 2004a;
Mayringer et al., 2011; Ustaömer et al., 2013; Rolland et al., 2016; Roland, 2017; Moazzen et al.,
2020), and the high-volume magmatism occurred during the exhumation of the HT-M to LP
metamorphic rocks (Topuz et al., 2010, 2023), we argue that the removal of the lowermost parts
of the thickened lithosphere after the accretion to Laurussia seems a more conceivable trigger for
the Late Carboniferous flare-up (Figs. 9d-e). A similar scenario has been proposed for the Late
Carboniferous magmatism in the French Massif Central, part of the Armorican Terrane
Assemblage in Central Europe (Fig. 1; Laurent et al. 2017; Vanderhaeghe et al. 2020).

To date, no Devonian accretionary complexes in the form of high-pressure rocks and
ophiolites have been reported. Middle Devonian igneous biotite ± hornblende granites display
subduction-related geochemical affinity (Aysal et al., 2012). Paleogeographic reconstructions
consistently show that the northern margin of the Armorican Terrane Assemblage was bound by
a south-vergent subduction zone (Stampfli and Borell, 2002; von Raumer et al., 2013). Hence, we
cautiously ascribe the Middle Devonian flare-up to the southward subduction of the Rheic ocean
beneath the Sakarya Zone during the northward drift (Fig. 9f).

Middle Ordovician-Silurian (460-430 Ma) probably represents the time of the rifting of
the Sakarya Zone from the northern margin of Gondwana (Fig. 9g), which is inferred from (i) the
within-plate geochemical features of amphibolite with Silurian protolith ages in the Sarıcakaya
Massif (Fig. 2), and (ii) the presence of Late Silurian to Devonian deep-sea sedimentary blocks in
the Paleo-Tethyan accretionary complexes in the western part of the Sakarya Zone (Topuz et al., 2020, and the references therein).

In conclusion, the locally recorded Late Ordovician-Silurian magmatic flare-up was likely caused by the rifting of the Sakarya Zone from the northern margin of Gondwana. The Middle Devonian magmatic flare-up and subsequent accretion to Laurussia during the Early Carboniferous were caused by the northward drift associated with the southward subduction of the Rheic ocean beneath the northern margin of Sakarya Zone. The most voluminous magmatic flare-up occurred during the Late Carboniferous and was most likely related to the loss of the lowermost sections of the thickened lithosphere. The southern margin of the Sakarya Zone became active following the accretion to Laurussia to the north. The Triassic magmatic flare-up, the most juvenile one among the documented ones in this study, were connected to the Paleo-Tethys' northward subduction beneath the Sakarya Zone.

## Comparison with the Armorican Terranes

Below, we make a preliminary comparison of the Paleozoic-Early Mesozoic detrital zircon age clusters of the Sakarya Zone with those of the Greater Caucasus in the east, the Balkans in the west and southern France in Europe (Fig. 1). All of these areas are part of the Armorican Terrane Assemblage. Similar to the Sakarya zone, the Greater Caucasus and the Balkans were also involved in Alpine orogeny, although southern France was outside the Alpine orogenic belt. The main difficulty in making the comparisons is that provenance studies based on U-Pb-zircon dating have mainly been performed on young sedimentary rocks such as the Oligocene to Miocene and recent clastic rocks (Allen et al., 2006; Vincent et al., 2013; Abbo et
al., 2020), rather than basement rocks or older clastic rocks unconformably overlying basement rocks.

A Middle to Upper(?) Jurassic sandstone sample from the Greater Caucasus shows U-Pb zircon age clusters peaking at 164, 242, 319 and 444 Ma (Allen et al., 2006) (Fig. 10a). Abbo et al. (2020) dated the detrital zircons from beach placer samples taken from the estuaries of three major rivers such as Strymon, Nestos and Evros that drain Serbo-Macedonian and Rhodope massifs (Fig. 1). The detrital zircons yield zircon age clusters peaking at 245, 309 and 460 Ma apart from the Mesozoic and Cenozoic ones (Fig. 10b). The Rhone River (southern France) drains a significant part of the European Armorican terranes, only a small portion of which was involved in Alpine orogeny (Fig. 1) (Avigad et al., 2022). The delta sands of the Rhone river yield U-Pb zircon age clusters peaking at 309, 456 and 553 Ma (Fig. 10c).

The Late Carboniferous magmatic flare-up is the most characteristic feature of the entire Armorican Terrane Assemblage. Likewise, Late Ordovician-Silurian magmatic flare-up which is recorded in some provenance studies in the Sakarya Zone is widespread. The Triassic magmatic flare-up seems to be restricted to the parts of the Armorican Terrane Assemblage that was later involved in Alpine orogeny. The Middle Devonian flare-up is local, confined to the Sakarya Zone.

## CONCLUSIONS

U-Pb detrital zircon ages from Middle Permian and Lower Jurassic sandstones from the Tokat Massif in conjunction with the literature data define the presence of three episodes of Paleozoic-Early Mesozoic high-volume magmatism in the Sakarya Zone: Triassic (250–230 Ma), Late
659 Carboniferous (325–310 Ma) and Middle Devonian (400–380 Ma). Among the magmatic flare-ups, the Late Carboniferous one is the most voluminous one, followed sequentially by the 660 Triassic, and the Middle Devonian ones. As the Sakarya Zone is a Gondwana-derived block that 661 was successively involved in Variscan and Alpine orogenies, these magmatic flare-ups cannot be 662 explained by processes along a single subduction zone. Detachment from the northern margin of 663 Gondwana and collage with the Laurussia should also be considered. The Late Carboniferous 664 flare-up event is well-documented in outcrop, and is marked by the formation of batholiths which 665 postdate the peak of HT-M to LP metamorphism by at least 5 to 8 Ma. It is ascribed to the 666 convective thinning of the lowermost part of the thickened lithosphere following the accretion of 667 the eastern Armorica with Laurussia. The Triassic magmatic flare-up event represents the most 668 juvenile crust formation. The widespread occurrence of Permo-Triassic accretionary complexes 669 suggests that the Triassic magmatic flare-up occurred in an active subduction setting, ascribed to 670 the northward subduction of the Tethys. On the other hand, related Triassic magmatic arc rocks in 671 the basement exposures of the Sakarya Zone have not been defined so far. The Middle Devonian 672 magmatic flare-up event is tentatively ascribed to the southward subduction of the Rheic oceanic 673 lithosphere beneath the Sakarya Zone. The Triassic and Late Carboniferous high-volume 674 magmatism is also a characteristic feature of the Balkans and the Caucasus. The Middle 675 Devonian one seems confined to the Sakarya Zone. Along with the Late Carboniferous flare-up, 676 the Late Ordovician-Silurian flare-up, which is locally recorded in the Sakarya Zone, is 677 representative of the Armorican Terrane Assemblage as a whole. 678

679 ACKNOWLEDGMENTS

680 We have benefited from the constructive reviews of Yann Rolland, Nurettin Kaymakçı and two 681 anonymous referees, and discussions with Aral I. Okay. We gratefully acknowledge the help of
Cemre Ay, İnal Demirkaya, Merve Dilbaz and Mutlu Özkan during mineral separation. This study has been financially supported by a research grant from the scientific and technological research council of Turkey (TÜBİTAK #116Y234) and the research foundation of the Istanbul Technical University (İTÜ BAP #39758).

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977 Pulur complex, NE Turkey: implications for the pre-Liassic evolution of the Eastern
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995 Middle Permian ophiolite fragment in Late Triassic greenschist-to blueschist-facies rocks in
996 NW Turkey: An earlier pulse of suprasubduction-zone ophiolite formation in the Tethyan

997 belt: Lithos, v. 300, p. 121–135, https://doi.org/10.1016/j.lithos.2017.12.005.

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1004 mafic-ultramafic intrusions in the Eastern Pontides (Pulur Complex): Implications for the

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1007 Ustaömer, P.A., Ustaömer, T., and Robertson, A.H.F., 2012, Ion probe U-Pb dating of the Central

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1057 Yılmaz, Y., Serdar, H.S., Genc, C., Yigitbas, E., Gürer, Ö.F., Elmas, A., Yıldırım, M., Bozcu, M.,

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1062 6000-km-long Neo-Tethyan arc system with coherent magmatic flare-ups and lulls in South

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1064 **FIGURE CAPTIONS**

1065 **Figure 1.** Simplified tectonic map of Europe and neighboring areas, showing main Phanerozoic

1066 sutures (modified after Topuz et al., 2020, and the references therein). Also are shown the

1067 drainage basins of the rivers Rhone, Evros, Nestos and Strymon.

1068 **Figure 2.** Distribution of the pre-Jurassic basement outcrops in northern Turkey (modified after

1069 Okay and Nikishin, 2015). IAES Izmir-Ankara-Erzincan suture; IPS Intra-Pontide suture;

1070 ITS Inner Tauride suture.

1071 **Figure 3.** Columnar section showing main pre-Jurassic basement units of the Sakarya Zone and

1072 its Jurassic-Cretaceous stratigraphy (modified after Topuz et al., 2020).

1073 **Figure 4.** Geological map of the Tokat Massif and the locations of the samples studied (modified

1074 after Hakyemez and Papak, 2002).

1075 **Figure 5.** U-Pb age distribution pattern of the detrital zircons from Middle Permian and Lower

1076 Jurassic sandstones from the Tokat Massif. Time scale is after Cohen et al. (2013, updated).

1077 **Figure 6.** *a* $^{206}$Pb/$^{238}$U zircon ages (Ma) versus initial $\epsilon$Hf(t) values of the detrital zircons from

1078 Lower Jurassic and Middle Permian sandstones from the Tokat Massif. *b* $^{206}$Pb/$^{238}$U zircon

1079 ages (Ma) versus initial $\epsilon$Hf(t) values of the Phanerozoic detrital zircons. In addition, the
1080 initial εHf values of zircons (dark blue lines; Karslı et al., 2016, 2020; Liu et al., 2018;
1081 Şengün et al., 2020; Topuz et al., 2020) and initial εNd values of whole rocks (green bars;
1082 Nzegge et al., 2006; Topuz et al., 2010; Dokuz, 2011; Aysal et al., 2012; Sunal, 2012;
1083 Kaygusuz et al., 2012, 2016; Karslı et al., 2016) are shown.

1084 **Figure 7.** Compositional variation of the dated detrital zircon grains in (a) Nb/Yb versus U/Yb
1085 after Grimes et al (2015), and (b) Hf versus Ti after Laurent et al. (2022).

1086 **Figure 8.** U-Pb detrital zircon age histograms with probability density curves. Binwidth is set to
1087 10 Ma. **a-b** This study, **c-d** the Late Triassic accretionary complexes in the western part of
1088 the Sakarya Zone (data from Ustaömer et al., 2016), **e** the Upper Cretaceous sandstones
1089 from the Sakarya Basin (data from Mueller et al. (2022), and **f** the Upper Cretaceous
1090 sandstones from Samsun (data from Liu et al. 2022).

1091 **Figure 9.** Possible drivers of the Paleozoic and early Mesozoic magmatic flare-ups in the Sakarya
1092 Zone. **a** Paleogeographic reconstruction for Late Carboniferous-Early Permian (modified
1093 after Stampfli et al., 2013), **b** 212–199 Ma magmatic flare-up probably caused by slab
1094 rollback and detachment after the accretion of an oceanic plateau or a seamount chain (Sc).
1095 Bc Baltica, Av Avalonia, Iz Istanbul Zone, Sz Sakarya Zone. **c** 250–230 Ma magmatic
1096 flare-up probably resulted from the northward subduction of the Paleo-Tethys (Pt) under
1097 Laurasia. **d** 325–310 Ma magmatic flare-up possibly originated by the removal of
1098 lowermost lithosphere after the accretion of Armorica (Arm) to Laurussia. **e** 400–380 Ma
1099 magmatic flare-up is probably related to the southward subduction of the Rheic ocean
1100 beneath the Sakarya Zone. **f** 440–420 Ma Silurian magmatic flare-up is ascribed to rift-
1101 related magmatism leading to the detachment of Armorica from the northern margin of
1102 Gondwana.
1103 **Figure 10.** U-Pb detrital zircon age histograms with probability density curves from the adjacent
1104 and distant Armorican terranes. **a** Middle-Upper? Jurassic sandstone from Greater
1105 Caucasus (data from Allen et al., 2006), **b** placer sand samples from the rivers Strymon,
1106 Nestos and Evros in the Balkans (data from Abbo et al., 2020), and **c** sand samples from the
1107 Rhone river delta (data from Avigad et al. 2022).

1108 **Supplemental Material** [Analytical Techniques, Figure S1. Comparative columnar sections of
1109 the Anatolide-Tauride Block and the Sakarya Zone, Figure S2. Field pictures, Figure S3. Thin
1110 section micrographs showing microtextural features of the studied sandstone samples, Figure S4.
1111 CL images of dated zircon grains, Figure S5. REE patterns of the dated zircon grains, Figure S6.
1112 Variation of U/Th, U/Yb and Nb/Yb ratios of the dated detrital zircon grains with age. Figure S7.
1113 U-Pb detrital zircon age histograms and probability density curves for literature data. Table S1
1114 Coordinates of the dated samples, Tables S2-S3 U-Pb and Hf metadata. Table S4 U-Pb-Th
1115 isotopic data and trace element compositions of zircons; Table S5 Hf isotopic data of zircons]
1116 Please visit https://doi.org/10.1130/XXXX to access the supplemental material, and contact
1117 editing@geosociety.org with any questions.
<figure><img src="image_2.png" /><figcaption>Figure 1</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>Figure 2</figcaption></figure>
<figure><img src="image_4.png" /><figcaption>Figure 3</figcaption></figure>

5. Permo-Triassic accretionary complex
<figure><img src="image_5.png" /></figure>

Middle Permian to Late Jurassic
low-grade metamorphic rocks
<figure><img src="image_6.png" /><figcaption>Figure 5</figcaption></figure>

Sample T-290A
Permian sandstone
n = 103 zircons
Youngest zircon: 295 ± 8 Ma
Oldest zircon: 2664 ± 57 Ma

Sample T-320A
Permian sandstone
n = 87 zircons
Youngest zircon: 299 ± 6 Ma
Oldest zircon: 2913 ± 60 Ma

Sample T-356B
Permian sandstone
n = 89 zircons
Youngest zircon: 294 ± 6 Ma
Oldest zircon: 2493 ± 45 Ma

Sample T-239
Jurassic sandstone
n = 116 zircons
Youngest zircon: 195 ± 4 Ma
Oldest zircon: 2698 ± 48 Ma

Sample T-267
Jurassic sandstone
n = 116 zircons
Youngest zircon: 235 ± 5 Ma
Oldest zircon: 2590 ± 45 Ma

Sample T-295
Jurassic sandstone
n = 105 zircons
Youngest zircon: 210 ± 4 Ma
Oldest zircon: 2049 ± 41 Ma
<figure><img src="image_7.png" /><figcaption>Figure 6</figcaption></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>

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

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

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

<figure><img src="image_16.png" /></figure>
a Late Carboniferous-Early Permian

<figure><img src="image_17.png" /><figcaption>b 215-198 Ma (latest Triassic-earliest Jurassic)</figcaption></figure>

<figure><img src="image_18.png" /><figcaption>c 250-230 Ma (Triassic)</figcaption></figure>

<figure><img src="image_19.png" /><figcaption>d 326-310 Ma (Late Carboniferous)</figcaption></figure>

<figure><img src="image_20.png" /><figcaption>e ca.345 Ma (Early Carboniferous)</figcaption></figure>

<figure><img src="image_21.png" /><figcaption>f 400-380 Ma (Middle Devonian)</figcaption></figure>

<figure><img src="image_22.png" /><figcaption>g 460-430 Ma (Late Ordovician - Silurian)</figcaption></figure>

<figure><img src="image_23.png" /></figure>
<figure><img src="image_24.png" /></figure>

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

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