rift setting

Journal Pre-proofs Mesoarchean bimodal volcanic rocks of the Onot greenstone belts, southwest‐ ern Siberian craton: implications for magmatism in an extension/rift setting O.M. Turkina, V.P. Sukhorukov, S.A. Sergeev PII: DOI: Reference:

S0301-9268(19)30655-2 https://doi.org/10.1016/j.precamres.2020.105731 PRECAM 105731

To appear in:

Precambrian Research

Received Date: Revised Date: Accepted Date:

22 November 2019 2 April 2020 2 April 2020

Please cite this article as: O.M. Turkina, V.P. Sukhorukov, S.A. Sergeev, Mesoarchean bimodal volcanic rocks of the Onot greenstone belts, southwestern Siberian craton: implications for magmatism in an extension/rift setting, Precambrian Research (2020), doi: https://doi.org/10.1016/j.precamres.2020.105731

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Mesoarchean bimodal volcanic rocks of the Onot greenstone belts, southwestern Siberian craton: implications for magmatism in an extension/rift setting

O.M. Turkinaa,b,*, V.P. Sukhorukova,b, S.A. Sergeevc a

V.S. Sobolev Institute of Geology and Mineralogy SB RAS, Koptyuga av. 3, Novosibirsk 630090, Russia

b

Novosibirsk State University, Pirogova st. 2, Novosibirsk 630090, Russia

c

Centre of Isotopic Research, A.P. Karpinsky Russian Geological Institute, Sredny av. 74, St. Peterburg

199106, Russia *Corresponding author: E-mail: [email protected] V.S. Sobolev Institute of Geology and Mineralogy SB RAS, Koptyuga av. 3, Novosibirsk 630090, Russia

Abstract This paper shows that Mesoarchean bimodal felsic and mafic metavolcanic rocks of the Onot greenstone belt (Sharyzhalgay uplift, the southwestern margin of the Siberian craton) formed as a result of rifting of Paleoarchean continental crust. Evidence for this comes from U-Pb zircon ages, whole-rock geochemical and Nd isotope data and Hf-in-zircon isotope data from a metarhyolite-basaltic unit. The over 80 km long Onot greenstone belt consists of tectonic sheets composed of metasedimentary-volcanogenic and tonalite-trondhjemite-granodiorite (TTG) rocks thrusted to the southwest onto high-grade rocks of the Kitoy terrane. The greenstone sequence includes a bimodal metavolcanic unit overlain by metavolcanogenic-sedimentary units. The felsic volcanic rocks formed at 2.88 Ga and were metamorphosed at 660-690°C and 6 kbar during the Late Paleoproterozoic. The high-Fe metarhyolites are enriched in the REE and HFSE and are compositionally similar to felsic volcanic rocks of Archean greenstone belts with lower La/Yb and Zr/Y (FIIIa and FIIIb types after Hart et al., 2004), and A-type granitoids. The highand low-Ti metabasalts (amphibolites) possess geochemical affinities to both subduction-related basalts and continental flood basalts contaminated by crustal material. The metarhyolites yielded

negative Nd(t) values (-3.8 to -0.8). Their zircons yielded negative to positive Hf(t) values (-8.5 to +1.1). The isotopic data suggest melting of heterogeneous crust, consisting of Paleoarchean plagiogneisses of the TTG basement complex and a juvenile source. The contribution of a mafic source resulted in enrichment of metarhyolites in FeO, MgO and TiO2. The crustal melting probably occurred at shallow depths at 2-4 kbar and at high temperatures higher than 900°C. The coeval formation of Onot basalts and low-P - high-T felsic melts could be triggered by decompressional mantle upwelling and subsequent extension and thinning of the subcontinental lithosphere. Worldwide, evidence for extension/rifting comes from Meso-Neoarchean mafic dikes and coeval intra-plate bimodal basalt-rhyolite volcanism. The events of extension/rifting and the formation of subcontinental lithospheric mantle suggest formation of stable and rheologically rigid continental plates. Keywords: Onot greenstone belt, U-Pb zircon age, geochemistry, Nd and Hf isotopes, stabilization of subcontinental lithospheric mantle

1. Introduction As recorded by numerous geological and geochemical data, the period from 3 Ga to 2 Ga was transitional due to the cooling of the Earth’s mantle and related changes in convective style (Condie, 2018 and references therein). The gradual cooling of the Earth led to the formation of stable, rheologically “rigid” continental lithospheric plates, isolated from the convecting mantle. Syntheses of data from mantle peridotites as well as sulfide and alloy inclusions record a sharp increase in the number of Re-depletion model ages (TRD) since 3 Ga and a distinct peak at 2.5-2.9 Ga (Carlson, 2005; Griffin et al., 2014), suggesting formation of stable subcontinental lithosphere in Late Archean. At upper crustal levels, the emplacement of mafic dikes and sills and the eruption of intraplate volcanic rocks, including bimodal basalt-rhyolite associations, typically accompany stretching and rifting of rheologically rigid lithospheric plates. Those processes can be best

illustrated by the formation of the Great Dyke of Zimbabwe at ca. 2.6 Ga (Oberthur et al., 2002). Several earlier events of extension/rifting happened in the beginning of the Mesoarchean, when a sequence of ca. 3.2 Ga sedimentary rocks, bimodal volcanic rocks, and layered mafic-ultramafic sills and dykes formed in the East Pilbara Terrane, Western Australia (Van Kranendonk et al., 2010) and when 3.2 Ga A-type granitoids intruded in the Kaapval craton of South Africa (Misra et al., 2017). For many years, the oceanic vs continental origin of Archaean greenstone belts has been a focus for many research teams (e.g., de Wit et al., 1987; 2011; Furnes et al., 2013; 2015; Kröner et al., 1996; 2013; Green et al., 2000; Tessalina et al., 2010; Sappin et al., 2018). In terms of geochemistry, the felsic volcanic rocks of greenstone belts can be classified into two major groups (Hart et al., 2004). The first dominant group (FI and FII types, after Hart et al., 2004) is compositionally similar to Archean TTGs (tonalite-trondhjemite-granodiorite) probably formed by the partial melting of mafic sources at high to intermediate pressures and emplaced in a compressional setting. The volcanic rocks of the second group (FIIIa and FIIIb types, after Hart et al., 2004) are characterized by lower (La/Yb)n and Zr/Y ratios and moderate to elevated HFSE content suggesting formation at shallower levels (< 10 km) in an extensional setting (Hart et al., 2004; Hollis et al., 2015). Group 2 volcanic rocks differ from typical Archean granitoids, such as TTG, sanukitoids and crust-derived granites (Laurent et al., 2014), but closer to modern A-type granitoids. Those potassium felsic volcanics are sparse in greenstone belts and other Archean provinces and are typically silica oversaturated and tholeiitic (ferroan). Their genesis, crustal sources, and geodynamic settings of emplacement remain debatable (Lesher et al., 1986; Morris, Witt, 1997; Hollis et al., 2015; Thurston, 2015; Zincone et al., 2016; Agangi et al., 2018; Savko et al., 2019). In the southwestern Siberian Craton, bimodal volcanic associations with potassic felsic volcanic rocks occur in the Onot granite-greenstone terrane of the Sharyzhalgay Uplift (superterrane) (Turkina and Nozhkin, 2008). The Onot granite-greenstone terrane comprises

plagiogneisses and plagiogranitoids of a Paleoarchean TTG complex and the metasedimentaryvolcanic rocks of the Onot greenstone belt (OGB). The geochronological data from the TTG complex suggest that it represents a basement to the bimodal volcanic greenstone sequence (Nozhkin et al., 2001). This paper presents new U-Pb zircon ages and Hf-in-zircon isotope data, and major and trace element bulk rock data from volcanic rocks of the Onot greenstone belt with foci on their petrogenesis and tectonic settings of formation. New geological and geochemical data provide evidence of an extension/rifting related origin of the Onot belt. Our new results contribute to the better understanding of the nature of Archean greenstone belts in general and the initiation of Archean intraplate felsic volcanism in particular.

2. Geological setting 2.1.

Regional geology

The Siberian Craton formed in the Late Paleoproterozoic (2.0–1.8 Ga) by the amalgamation of Archean and Paleoproterozoic terranes (Rosen et al., 1994, 2003). Gravity and magnetic survey data and limited lithological data from deep boreholes and from kimberlitehosted crustal xenoliths show that the basement of the Siberian Craton consists of five major provinces/domains: Tungus, Aldan, Anabar, Olenek and Stanovoy (Rosen et al., 1994; Smelov and Timofeev, 2007). A major part of the Siberian Craton is covered by typically 2-5 km thick Late Precambrian and Phanerozoic sedimentary rocks, but up to 10 km thick and more in rifts (Rosen et al., 1994). The basement of the Siberian Craton consists of two principal types of lithologic assemblages: granulite-gneiss and granite-greenstone (Rosen et al., 1994). The Archaean to Paleoproterozoic basement crops out in the Aldan–Stanovoy and Anabar shields, in the Sharyzhalgay and Olenek uplifts, and within the Akitkan and Angara orogenic belts. The Aldan–Stanovoy Shield includes five large terranes, the West Aldan granite–greenstone composite terrane, the Central Aldan granulite-gneiss terrane and the East Aldan Superterrane including the Uchur granulite-paragneiss and the Batomga granite–greenstone terranes, the

Tynda tonalite-trondhjemite gneiss terrane, and the Chogar granulite-orthogneiss terrane (Smelov and Timofeev, 2007). The Anabar Shield includes Magan tonalite-trondhjemite, Daldyn granulite-gneiss and Khapchan paragneiss terranes separated by the Kotuikan and Billyakh shear zones (Smelov and Timofeev, 2007). The largest Tungus province occupies the western part of the craton and crops out only in the Sharyzhalgay uplift (Fig. 1, A). The Sharyzhalgay uplift extends for a distance of 400 km north of Lake Baikal to the Oka River and consists of four units: the Bulun and Onot granitegreenstone terranes in the north-west and Kitoy and Irkut granulite-gneiss terranes in the southeast (Fig. 1, B). The terranes are bounded by NS- and NW-trending shear zones and faults. The Irkut and Kitoy terranes consist of Paleoarchean to Paleoproterozoic magmatic and sedimentary rocks metamorphosed in the amphibolite and granulite facies conditions (Poller et al., 2005; Sal’nikova et al., 2007; Turkina et al., 2012). The Bulun and Onot terranes consist of Paleoarchaean

tonalite–trondhjemite–granodiorite

(TTG)

complexes

and

greenstone

metasedimentary-volcanic units (Nozhkin et al., 2001; Bibikova et al., 2006; Turkina and Nozhkin, 2008; Turkina et al., 2009). The Sharyzhalgay superterrane formed by Late Paleoproterozoic events of accretion and collision. Evidence for those comes from the 1.88-1.84 Ga episodes of metamorphism and granitoid magmatism (Donskaya et al., 2002; Poller et al., 2004; Turkina and Kapitonov, 2019 and references therein). The Onot terrane is a 80 km long and 8-10 to 20 km wide graben-shaped structure extending from the Onot River in the northwest to the Kitoy River in the southeast (Fig. 1). The terrane borders the Kitoy terrane along the Alagna-Kholomkha fault in the southwest and the Irkut terrane in the east, along the Dabad fault. In the northeast, the Onot complexes are discordantly overlain by the Ediacaran-Cambrian sedimentary cover of the Siberian platform. The Onot terrane is dominated by Paleoarchean TTG rocks and metasedimentary-volcanic greenstone rocks (Fig. 1, 2) (Nozhkin et al., 2001; Turkina, Nozhkin 2008). The TTG and greenstone rocks are deformed to make a set of thrust sheets over the high-grade rocks of the

Kitoy terrane. The gneissic plagiogranites and biotite plagiogneisses (TTG complex) crops out mainly in the northwestern part of the terrane and along its northeast boundary (Nozhkin et al., 2001; Turkina and Nozhkin 2008). The Paleoarchean (3.4 Ga) TTG rocks contain numerous amphibolite boudins and are compositionally similar to Archean high-Al TTG suites (Bibikova et al., 2006; Rosen, Turkina, 2007). 2.2.

Onot greenstone sequence

The NW-striking metavolcanic-sedimentary sequence of the Onot greenstone belt (OGB), generally dips to the SW at high angles. According to the 1980-ties geological survey, the section comprises three units: Malaya Iret, Kamchadal and Sosnovy Baits (Fig. 2). The lithologically uniform ~3000 m thick Malaya Iret unit (lower) can be clearly traced along the southwestern boundary of the Onot terrane. It consists of light gray fine-grained biotiteamphibole, amphibole-biotite, or biotite microgneisses, and boudin-like bodies and rare thin (n∙10 cm) layers of fine-grained amphibolite (Fig. 3). There are also medium-grained marble and calciphyre in the upper part of the unit. The southeastern boundary of the Onot greenstone belt runs from the outcrops of Malaya Iret amphibole-biotite microgneiss on the left side of the Kitoy River, downstream the Kholomkha creek (Fig. 2B). Although most of the contacts between the microgneisses and the overlapping amphibolites of the Kamchadal unit are hidden by thrusting and related schistosity of amphibolites, they, in places, become visible due to layers of marbled dolomite. The 1250-2500 m thick Kamchadal unit (middle) is dominated by amphibolite and amphibole schists with subordinate marbled dolomite, and magnesite in its lower part, and talcchlorite and chlorite-sericite schists with thin (n∙10 cm) layers of limestone, dolomite, and quartzites in its upper parts. The Sosnovy Baits unit (upper) crops out along the northeastern boundary of the Onot terrane. The boundary between the Kamchadal and Sosnovy Baits units is a thrust separating tectonic sheets composed of TTG rocks. The lithology of the Sosnovy Baits unit changes laterally from mainly amphibolite with layers of chlorite-actinolite schists and

banded iron quatzites (BIF) to abundant biotite- and garnet-bearing schists in the central and southeastern parts of the Onot greenstone belt. The thickness of the Sosnovy Baits unit varies from 850-1150 m in the northwest to 3500 m in the southeast. The volcanic-sedimentary sequence and the TTG complex experienced epidoteamphibolite metamorphism at about 1.88 Ga and were intruded by post-collisional Paleoproterozoic granitoids (Turkina, Nozhkin 2008). The 1.86 Ga largest Shumikha pluton (1.86 Ga) consists of biotite granodiorites and A-type granites (Donskaya et al., 2002; Turkina, Kapitonov, 2017). 2.3.

Sample description

Most samples under study come from the Malaya Iret unit in the northwestern, central, and southeastern parts of the Onot greenstone belt. In the northwestern areas, on the southwestern slope of Mt. Kamchadal (5236, 10205), the unit is dominated by amphibole and biotite-amphibole microgneisses with scarce thin boudins of fine-grained amphibolites. Biotite and amphibole-biotite microgneisses are widespread in the central part of the belt (at the western slope of Mt. Arban) and host numerous amphibolite bodies. The microgneisses exposed in the southeastern part, on the left side of the Kitoy River (at 5211-5210 and 10242-10244) have similar compositions but look strongly deformed and contain rare boudins and thin layers of amphibolite (Fig. 3); they are cut by nearly concordant and branching granitic veins.

3. Analytical procedures Major elements were determined by X-ray fluorescence spectrometry (XRF) on a Thermo Scientific ARL 9900 IntelliPower™ Series ARL 9900 X-ray WorkStation at the Center for Multi-element and Isotope Research of the Russian Academy of Sciences, Siberian Branch (Novosibirsk). Trace element abundances were measured on a Finnigan Mat ELEMENT Inductively Coupled Plasma Mass Spectrometer equipped with a U-5000AT+ultrasonic spray, to

an average precision of 2-7 rel. %. The calibration was against the BHVO-1 external standard. The detection limits of REE and HFSE were 0.005 to 0.1 ppm. In situ U-Pb zircon dating was performed by sensitive high-resolution ion mass spectrometry (SHRIMP II) at the Center for Isotopic Research of the Russian Geological Research Institute (St. Petersburg). The selected zircon grains were mounted in epoxy resin together with chips of Temora and 91500 reference zircons. The choice of analytical sites was based on optical transmitted-light and cathodoluminescence (CL) images displaying the zircon structure and zonation. The U-Pb measurements followed the techniques described by Williams (1998), with a spot diameter of 25 m and a mass-filtered O2- primary beam of 4 nA at the 2 m pit depth. The results were processed in SQUID (Ludwig, 2000). The U/Pb ratios were normalized to that in the Temora standard zircon (0.0668) corresponding to the age of 416.75 Ma. The errors of single analyses (U/Pb isotope ratios and ages) and the errors of concordant ages and intercepts are quoted at the 1 and 2 levels, respectively. The concordia diagram was plotted in ISOPLOT/ET (Ludwig, 1999). The Lu–Hf isotope analyses of the zircons were carried out on a New Wave DUV 193 nm laser-ablation microprobe, attached to a ThermoFinnigan Neptunе multi-collеctor ICP-MS, at the Center of Isotopic Research of the Russian Geological Research Institute in St. Petersburg, following the method described in detail by Griffin et al. (2000). The measurements were applied to the same spot sites as the U-Pb SHRIMP analysis but the spot size and pit depth were, respectively, 50 μm and 20-40 μm. Laser repetition was 5-7 Hz at 130 mJ, and the ablation and pre-ablation time was 32 s and 10 s, respectively. Raw count rates for 176(Hf+Yb+Lu), 177Hf, 178Hf, 179Hf, 180Hf

interference corrections of

176Lu

and

176Yb

were collected; on

176

175Lu

172Yb, 174(Yb+Hf), 175Lu,

and

172Yb

Hf. The average values of

were used for 176Hf/177Hf

for

reference zircons were 0.28268023 (Temora; n=10), 0.28249716 (Mud Tank; n=6) and 0.28199420 (GJ-1; n=6). All errors are quoted at the 2 level. The

176Lu

decay constant of

1.867×10-11 yr-1 was taken from (Söderlund, 2004). The initial εHf(t) values were calculated with

reference to chondritic ratios of

176Hf/177Hf

of 0.282785 and

176Lu/177Hf

of 0.0336 (Bouvier et

al., 2008). Two-stage model Hf ages (TCHf(DM)) were calculated based on and

176Hf/177Hf

initial

176Lu/177Hf

=0.0384

=0.28325 ratios for the depleted mantle (Griffin et al., 2000), by projecting the

176Hf/177Hf

value of zircon onto the line of depleted mantle, with the use of the average

crustal value of 176Lu/177Hf = 0.015. The Sm-Nd isotope composition and the concentrations of Sm and Nd were analyzed at the Geological Institute of the Kola Science Centre, Russian Academy of Sciences (Apatity), on a Finnigan MAT 262 multi-collector mass spectrometer in a static mode, as in (Bayanova, 2004). The total laboratory blanks were 0.06 ng for Sm and 0.3 ng for Nd. The reproducibility was ~0.2% (2) for 147Sm/144Nd

147Sm/144Nd

and

146Nd/144Nd=0.7219

ratios and 0.003 (2) for Nd isotopic analyses (Table 2). All

143Nd/144Nd

ratios

were

normalized

to

the

standard

values

of

and 143Nd/144Nd=0.511860 or 0.512100 reported for the La Jolla and JNdi-1

Nd standards, respectively. During the period of measurement, the weighted mean

143Nd/144Nd

ratios for the La Jolla Nd standard were 0.511805±8 (n=13); values for the JNdi-1 standard were 0.512088±12 (n=17). Nd(t) values and model ages TNd(DM) were calculated using the currently accepted values of CHUR (143Nd/144Nd=0.512638 and

147Sm/144Nd=0.1967)

Wasserburg, 1984) and depleted mantle (DM) (143Nd/144Nd=0.513151 and

(Jacobsen,

147Sm/144Nd=0.2136)

reservoirs (Goldstein, Jacobsen, 1988).

4. Results 4.1.

U-Pb ages of zircons from felsic metavolcanics

Most samples of microgneisses (felsic metavolcanics) contain few zircons in spite of the high concentrations of Zr (360-600 ppm). Only microgneiss 36-13 from the southeastern OGB sampled in the left side of the Kitoy River carried an amount of zircons sufficient for isotopic dating. The zircons are cherry-brown to brown transluscent subhedral short-prismatic crystals, 120-320 μm long, with length to width proportions of 2 to 3.0. CL images show weakly zonal

central parts and dark rims (Fig. 4); large grains look strongly altered and contain numerous dark inclusions. Two zircon separates of different sizes were analyzed in two runs (Table 1). All zircons contain low to moderate amounts of U (71-494 ppm) and Th (33-409 ppm); the Th/U ratios (0.48-0.88) are in the range typical of magmatic zircons. The results show two age groups of zircons (Fig. 4). The older group includes fifteen grains with a concordant age of 2886 ± 8 Ma (MSWD = 0.015). Four grains yielded a concordant age of 2846 ± 13 Ma (MSWD = 0.24), and together discordant ages (D = 3-12) derived from other five zircons they form a discordia with an upper intercept age of 2838 ± 20 Ma (MSWD = 0.50). 4.2.

Petrography, mineral composition and PT-conditions of metamorphism

The Malaya Iret mircrogneisses are fine-grained rocks with thin banded texture. The mineral assemblage of most samples is Kfs+Pl+Qtz+Hbl+Bt, biotite and garnet-bearing biotite gneisses are less common. The total amount of biotite and amphibole is about 10 vol. %. Accessory minerals are titanite, apatite, ilmenite, allanite and zircon. The Malaya Iret amphibolites are fine-grained, weakly gneissic rocks with a mineral assemblage of Hbl+Pl+Qtz±Bt; the amount of quartz reaches 5 vol. %. The amphibolites sampled at boudin-like bodies contain no more than 2-3 vol. % biotite and are enriched in titanite. On the contrary, the amphibolites sampled in the thin interlayers are enriched in biotite (up to 8 vol. %) and ilmenite.

Mineral chemistry was analyzed in five biotite-amphibole microgneiss samples (Table 2). The extremely high values of Fe# in amphibole and biotite are consistent with the high ferroan whole-rock composition. According to the classification of (Leak et al., 1997) amphibole is hastingsite. Biotite shows extremely low XMg (0.01-0.02) and elevated TiO2 (3.2-3.7 wt. %). Plagioclase is albite with a minor portion of orthoclase. The temperature of metamorphism was estimated using an Amph-Pl geothermometer (Holland and Blundy, 1994) and edenite-tremolite equation, because the rocks are rich in quartz. The temperatures are 660 to 690оС for three samples and is as high as 750оС for sample 6-03

from the northwestern part of the Onot belt (Table 3). The pressures estimated using the Al-inAmph and amphibole-plagioclase geobarometers respectively range from 6.2 to 6.9 kbar (except for one estimate at 7.3 kbar) and from 5.4 to 6.2 kbar (one estimate at 8.5 kbar). The Malaya Iret microgneisses metamorphosed at temperature of 660-690oC and pressure of 6±0.5 kbar, i.e, under the PT conditions at the upper stability limit of the ferroan staurolite found in the high-Al garnet-staurolite-biotite schists sampled in the upper section of the Onot belt. 4.3.

Major- and trace-element chemistry of microgneiss and amphibolite

Compositionally, the microgneisses with SiO2 spanning 72-76 wt % correspond to medium to high-K (K2O/Na2O = 1.0-1.7) and metaluminous to slightly peraluminous (ASI=0.941.09) rhyolites (Table 4, Fig. 5a). ). According to the classification of Frost et al. (2001) they are alkali-calcic granitoids with extremely high Fe# (FeO*/(FeO*+MgO)=0.89-0.99) (Fig. 5b, c) Their TiO2 enrichment (0.31 to 0.55 wt.%) is consistent with the high percentage of titanite. As SiO2 increases, the contents of FeO, MgO, TiO2 and CaO decrease and K2O increases. These signatures coupled with the fine-grain textures suggest volcanic rather than sedimentary origin of the microgneisses. The samples from three geographically separated areas are generally similar in respect to the contents of both major oxides and trace elements. In spite of the high content of SiO2, the microgneisses have variable to high concentrations of Cr (8-106 ppm), Ni (6-56 ppm), and Sc (3.6-5.5 ppm). The K2O enrichment correlates with high Ba (640–936 ppm), medium Rb (83–199 ppm) and low Sr (49–130 ppm). The multi-element spectra (Fig. 6a, b) show a weak enrichment in LILE (Fig. 8). The REE patterns (Fig. 6 a, b) are weakly fractionated (La/Yb)n = 2.8–9.6) with enriched HREE and distinct negative Eu anomalies (Eu Eu* = 0.5–0.8). The multi-element spectra typically show deep Sr, P, and Ti troughs and shallow Nb trough. According to the high HFSE, in particular Zr (357-607 ppm) and the high Fe#, the metarhyolites are similar to A-type granitoids (Whalen et al., 1987; Frost and Frost, 2011). The temperatures of Zr saturation calculated as in (Watson and Harrison, 1983) vary from 824 to 888оC (Table 4).

The amphibolites have basaltic or leucobasaltic compositions with Mg# ranging from 0.59 to 0.27 (Table 4; Fig. 5a). Most samples of amphibolites from the boudin-like bodies are depleted in TiO2 (0.66–1.22 wt%) and P2O5 (0.045–0.11 wt%) and show lower concentrations of Cr (62-167 ppm) and Ni (23-89 ppm) compared to the biotite-enriched amphibolites from the thin layers having higher TiO2 (2.32-2.96 wt%) and P2O5 (0.23–0.34 wt%). The low-Ti amphibolites are characterized by low contents of incompatible elements and flat to weakly fractionated REE patterns (Fig. 6c). With increasing SiO2, (La/Sm)n ratios increase from 1.2–1.9 to 2.4-2.8 and Eu/Eu* ratios decrease from 0.9-1.0 to 0.7–0.8 (Fig. 9c). The growth of SiO2 contents is accompanied by the enrichment in incompatible elements (Rb, Th, Zr, Nb). The PMnormalized multi-element patterns of low-Ti amphibolites demonstrate LILE enrichment and shallow to deep troughs at Nb, Ti, and P (Fig. 6c). The high-Ti amphibolites are enriched in REE, Zr (203-270 ppm), Nb (10-13 ppm), and Y, but yield lower (La/Yb)n ratios of 1.7 to 4.1 and weakly fractionated multi-element patterns with deep troughs at Nb and Sr. 4.4.

Whole-rock Nd and Hf-in-zircon isotope composition

The metarhyolites are characterized by high

147Sm/144Nd

ratios of 0.115 to 0.134 and

Nd(t) values ranging from -0.8 to -3.8 (Table 5, Fig. 7a). Their 3.2-3.5 Ga TNd(DM) model ages suggest an origin from a source with long crustal prehistory. Two amphibolite samples from the Malaya Iret and Kamchadal units yielded positive Nd(t) values (+4.2 and +4.8) indicating a depleted mantle source. Twelve zircons from the metarhyolite show a large range of

176Hf/177Hf

(0.00872-

0.002967) and mainly negative to weakly positive Hf(t) (-8.5 to +1.1) (Table 6, Fig. 7b). The TCHf(DM) model age of zircon is 3.3 to 3.8 Ga, i.e. corresponds to a Paleoarchean crustal source.

5. Discussion 5.1.

Age of the Onot metavolcanic-sedimentary complex

The oscillatory zoning and Th/U ratios ranging from 0.48 to 0.88 imply that the zircon from metarhyolite is of magmatic origin, i.e. the age of 2886 ± 8 Ma is the correct age of the volcanic rocks in the lower part of the greenstone section. The younger ages obtained from several zircons may result from their postmagmatic alteration with participation of fluids. The previously analyzed zircons from the biotite-amphibole microgneiss (sample 2-03) sampled in the northwestern part of the Onot belt (Mt. Kamchadal) show strongly discordant U/Pb isotopic ratios. The age of zircon from sample 2-03 at the upper intercept of discordia and concordia is 2888 ± 25 Ma (Turkina, 2010), i.e. within the analytical error it fits the age of zircon from microgneisses in the southeastern Onot greenstone belt. The upper age boundary of sedimentation in the Onot belt is constrained by the U-Pb ages of detrital zircons of the staurolite-garnet-biotite schists from upper part of section. Most of the detrital zircons show the ages of 2.7, 2.78, 2.8, 2.94-2.97, and 3.35 Ga (Turkina et al., 2014). The age of the youngest cluster of zircons constrains the time of the deposition at ~ 2.7 Ga. Thus, the Onot greenstone sequence formed during period of 2.88 to 2.7 Ga. The

40Ar/39Ar

ages of amphibole from sample 2-03 record an episode of Late

Paleoproterozoic metamorphism at about 1.88 Ga (Turkina, Nozhkin, 2008), which was followed by post-orogenic emplacement of granites at 1.86 Ga (Turkina, Kapitonov, 2017). 5.2. Genesis of felsic metavolcanics The felsic metavolanics of the Onot greenstone belt are similar to A-type granites in major and trace element compositions: they possess extremely high FeO*/(FeO*+MgO), REE and HFSE, in particular Zr. A-type granites and their volcanic counterparts can form through fractional crystallization of mafic magma with possible crustal contamination, melting of a quartz-feldspathic crustal source or mixing of melts derived from mafic and felsic crustal sources (Turner et al., 1992; Patiño Douce, 1997; King et al., 2001; Vander Auwera et al., 2003; Yang et al., 2006; Shellnutt and Zhou, 2007; Frost and Frost, 2011; Turkina and Kapitonov, 2019). The composition of ferroan granitoids and the results of experimental melting suggest two main

scenarios of granite generation: (1) fractional crystallization of mafic magmas and crustal contamination to produce the calc-alkalic to alkalic varieties of A-type granitic rocks or (2) melting of quartz-feldspathic crustal sources to produce alkali-calcic to calc-alkalic leucogranites (Frost and Frost, 2011). The major-element compositions of the Onot metarhyolites, i.e. high SiO2 and extremely high Fe# suggest either extreme differentiation of a tholeiite-basaltic melt or melting of a crustal quartz-feldspathic source at low H2O activity. The decreasing of FeO+MgO, CaO, and TiO2 at increasing SiO2, and Sr, Ti, P, and Eu troughs in multi-element spectra (Fig. 6a, b), indicate fractionation of clinopyroxene, plagioclase, Fe-Ti oxides, and apatite. The metarhyolites show values of Fe# similar to those recorded in the Skaergaard granophyre (Wager and Brown, 1967) and in reduced А-granites (Dall'Agnol and Oliveira, 2007): >0.9 and ≥0.85, respectively. The origin of rhyolite through the differentiation of mafic magma is inconsistent with the bimodal composition of the Onot volcanic rocks. Most of SiO2-rich (>70%) alkali-calcic metaluminous and weakly peraluminous A-granites form by the melting of quartz-feldspar crustal material at shallow depths (Frost and Frost, 2011); that process was reproduced in dehydration melting experiments at 2 to 4 kbar (Patiño Douce, 1997; Bogaerts et al., 2006). Comparison with experimental data (Patiño Douce, 1997; Bogaerts et al., 2006) (Fig. 6b, c) shows that by key parameters, such as high Fe#, MALI index (Na2O+K2O-CaO) and low ASI index (0.94-1.08) the Onot rhyolites are similar to the melts derived at low pressures of 2-4 kbar. In a recent review, Frost and Frost (2011) also concluded that low-pressure conditions are required to obtain metaluminous to slightly peraluminous ferroan melts capable to produce Atype granite. On the contrary, the higher pressure experiments at 6-10 kbar produced strongly peraluminous melts (Skjerlie and Johnson, 1993). The Onot metarhyolites are show significantly higher FeO and MgO contents compared to the melts derived experimentally from tonalitic and granodioritic sources (Patiño Douce, 1997; Bogaerts et al., 2006) (Fig. 6d). The melts derived from quartz-feldspathic sources

typically have (FeO*+MgO) <5 wt.% at SiO2 = 64–80 wt.% (Singh and Johannes, 1996; Skjerlie and Johnston, 1996; Bogaerts et al., 2006). There are other localities of intra-continental rhyolites and granites characterized by higher FeO and MgO (Zincone et al., 2016; Savko et al., 2019) compared with experimental melts (Fig. 6). The enrichment in Fe and Mg can be explained by: (i) accumulation of cumulate phases, (ii) capture of residual phases, (iii) high temperature of melting. The Contendas rhyolites of the Gavião Block in Brazil contain inclusions of plagioclase, biotite and apatite which are cumulate minerals separated during the differentiation of a parental granitic melt (Zincone et al., 2016). Savko et al. (2019) suggest accumulation of magnetite in the Atamansky granites of the Kursk domain in the south of the East European Craton. According to the experiments of Patino Douce and Johnson (1991), the solubility of FeO and MgO in a silica-rich melt sharply increases at a temperature higher than 1000°C. Those parameters fit the derivation the Onot rhyolites and other intraplate rhyolites and granites at a temperature higher than 900°C. The high contents of FeO and MgO in granitoids can be often caused by selective entrainment of residual phases (Clemens et al., 2011). In our case, those residual phases are Ca-plagioclase, clinopyroxene and/or amphibole, ilmenite or magnetite. The content of FeO, MgO, Cr, and Sc in the Onot rhyolites are similar to those in the Kursk A-type granites (Savko et al., 2019), but higher than in the Kursk and Contendas rhyolites (Zincone et al., 2016). Therefore, we suggest that the enrichment of the Onot rhyolite in Fe, Mg, Cr, and Sc resulted from of smaller fractionation of their parental melts and/or from the capture of cumulate/residual minerals by felsic melts. The formation of the Onot rhyolites was previously attributed to the melting of Paleoarchean TTGs based on their Nd isotope composition (Turkina and Nozhkin, 2008). However, the plagiogneiss and plagiogranites of the Onot TTG basement have similar or lower concentrations of FeO, MgO, Cr and Sc compared to the rhyolites (Fig. 6). Therefore, we assume a contribution of a more mafic source to the genesis of the Onot rhyolites.

5.3.

Isotopic constraint on melt sources for felsic metavolcanics

The negative Nd(t) values (-3.8 to -0.8) and TNd(DM) of 3.3–3.6 Ga in the metarhyolites allow us to consider the Paleoarchean TTG basement as their most probable source, but the metavolcanics have a more radiogenic Nd isotope composition in comparison with that of the TTGs (Nd from -5.3 to -8.5 at 2.88 Ga), that implies a contribution from a juvenile source (Fig. 7a). The wide range of Hf(t) in zircons of metarhyolites (+1.1 to -8.5) may result from the melting of heterogeneous crustal material or from the presence of residual zircon in source. If zircon remains as a residual phase during low-degree melting of felsic crust, the resulting melts would become more radiogenic as unradiogenic Hf stays mostly in residual zircons (Tang et al., 2014; Chen et al., 2015). However, the metarhyolites under study formed at high temperatures of more than 890 estimated by the temperature of Zr saturation and do not contain inherited zircon cores. Therefore, we conclude that a protolithic zircon dissolved completely. Evidence for a heterogeneous crustal source present in the melting region comes from the wide range of Hf(t) which partly overlaps with the isotope evolution trend of the Onot Paleoarchean crust reconstructed from TTG zircons but shifts toward higher Hf values (Fig. 7b). Thus, the isotope data for the Onot metaryholite and their hosted zircons are in line with melting of Paleoarchean TTGs and a more juvenile mafic/sialic source. The contribution of a mafic source is consistent with the diagram of Nb/Yb – Th/Yb (Fig. 8) showing that the metaryolites plot between the Onot basalts and TTG basement rocks. The potential mafic source could be amphibolites from inclusions in the TTG basement rocks or mafic melt producing greenstone metabasalts. The amphibolites of both types have highly positive initial Nd(t) values ranging from +2.6 to +1.1 for the amphibolites in the TTGs and from +4.7 and +4.2 for greenstone amphibolites (Table 5, Fig. 7a). The Hf isotope compositions of most detrital zircons from the overlying garnet-staurolite-biotite schists (Hf(t) ranging from +3.7 to -11.8 at 2.7 Ga) was previously interpreted as a result of the contribution of juvenile crust (Turkina et al., 2014) (Fig. 7b).

5.4. Rifting or extension? Hartlaub with co-authors (2004) summarized characteristic features of greenstone belts formed in continental rifting settings. The strongest argument for continental rifting is the presence of old sialic basement unconformably overlain by clastic and/or carbonate sediments covered by mafic and ultramafic volcanic rocks. For example, the contact between the Archean Murmaс Bay Group of the Rae Province, Canada, with the basement is marked by a bed of polymictic conglomerates overlain by quartzites, marbles and basaltic flows (Hartlaub et al., 2004). There are no clastic terrigenous sediments, e.g., conglomerates with quartz pebbles or quartz-rich sandstones, at the base of the Onot greenstone belt, probably due to the later thrusting, which disturbed the initial geological relationships. The Onot felsic volcanics are overlain by marbles obviously deposited in shallow-water environment. The basalts overlying the felsic volcanics contain interbeds of carbonates, BIF and shales indicating a deepening basin due to continued extension and rifting. Another evidence for the rifting setting of the formation of the Onot belt is the bimodal and antidromic character of volcanic associations. Our new geochemical data from the Onot metavolcanic rocks provide more insights into their tectonic origin. The low-Ti metabasalts of the Malaya Iret unit show weak LILE enrichment and weak to prominent Nb depletion in the multi-element spectra (Fig. 6c), i.e. similar to subduction-related basalts. Note, that the increase of SiO2 and incompatible elements and the small negative Eu anomaly in the Onot low-Ti amphibolite may result from crustal contamination of mafic melts. Evidence for the crustal contamination comes from the oblique trend from the mantle array towards the rocks of the TTG complex in the Th/Yb - Nb/Yb diagram (Fig. 8) (Pearce, 2008). The amphibolites of the Onot upper units show lower Th/Yb ratios indicating a lesser degree of contamination. The high-Ti amphibolite occurring thin layers in the metarhyolites (Fig. 6c) are strongly enriched P, Zr, Th, and REE, i.e. similar to continental flood basalts (Hornig, 1993; Farmer, 2003). However, they have higher contents of HREE probably due to the melting at shallower depths above the garnet stability field. The coexistence

of high- and low-Ti basalts is typical of continental flood and riftogenic mafic volcanics that originated from sources with different contributions of asthenospheric mantle, subcontinental lithospheric mantle, and deep plume materials (Peate and Hawkesworth, 1996; Marzoli et al., 2000; Farmer, 2003). The formation of the Onot belt remains ambiguous because the amphibolites of the lower unit are compositionally similar to both subduction-related and continental within-plate/riftingrelated basalts. Therefore, to determine their tectonic origin we should also discuss compositions of associated felsic rocks and PT-conditions of their formation. Archean greenstone belts typically comprise several types of felsic volcanics (Lesher et al., 1986; Hart et al., 2004; Thurston, 2015; Hollis et al., 2015). Types FI and FII are calc-alkalic dacites and rhyolites with high to moderate (La/Yb)n and Zr/Y and low concentrations of HFSE interpreted by some as strong geochemical signature of subduction origin. They are similar to Archean TTG formed via high-pressure partial melting of mafic sources. Felsic volcanics of types FIIIa and FIIIb are ferroan (tholeiitic) rocks with low (La/Yb)n and Zr/Y, and moderate (FIIIa) to high (FIIIb) HFSE and show, respectively, either weak or zero geochemical subduction signatures. The last two types result from the low-pressure melting of mafic or felsic sources in extensional or rifting settings (Hart et al., 2004; Hollis et al., 2015). The Onot metarhyolites geochemically correspond to FIIIa or FIIIb felsic volcanics, i.e. similar to A-type granites. They formed at P = 2-4 kbar and T > 900°C as inferred from the compositions of the melts obtained in experiments (Patiño Douce, 1997; Bogaerts et al., 2006). The high temperatures of melting accord well with the crystallization of zircon at 820-890°C, which is the minimal melt temperature assumed in the absence of inherited zircon cores (Miller et al., 2003). Similar conditions were estimated for Atype granites formed in within-plate, rifting, and post-collisional extension settings (Frost and Frost, 2011). The coeval formation of Onot basalts and low-P - high-T felsic melts could be triggered by decompressional mantle upwelling and subsequent extension and thinning of the

subcontinental lithosphere. Those conditions can provide both basaltic volcanism and hightemperature melting of a heterogeneous (mafic and tonalitic) crustal source. In summary, the geochemical and isotope characteristics of the Onot felsic and associated mafic metavolcanics and the geological data allow us to consider the Onot GB as a Meso-Neoarchean extension/rifting structure formed on Paleoarchean continental basement. For example, continental rift origin of Mesoarchean felsic to ultramafic volcanic rocks has been proposed for the Lac Guyer area of the La Grande Subprovince, Canada (St. Seymour and Francis, 1988; Sappin et al., 2018). Similar to the Onot greenstone belt, the succession of magmatism recorded in the Guyer Group includes (older to younger): (i) emplacement of tonalite and granodiorite plutons, (ii) eruption and extrusion of felsic volcanic and subvolcanic rocks and (iii) successive eruptions of basalts to komatiites. Accordingly, we accept a similar model for the Onot belt. During the extension of the TTG basement, the basaltic melts derived from a deep mantle source initiated the partial melting of TTG crust to form both rhyolitic magmas (dominant) and TTG contaminated basaltic magmas (subordinate). Later, a crustal magma chamber was destroyed as the crust melted, allowing basaltic magmas to erupt without crustal contamination. 5.5.

Implications for evolution of the Sharyzhalgay superterrane

The previous and our new data allow reconstruction of the possible Early Precambrian tectonic history of the southwestern Siberian craton. The crustal growth within all terranes of the Sharyzhalgay Uplift (superterrane) started at ca. 3.4–3.3 Ga (Poller et al., 2005; Bibikova et al., 2006; Turkina et al., 2009; 2011), but their subsequent evolutions were different, probably due to the rifting resulting in the detachment of the Kitoy and Irkut terranes from the margin of the Tungus continental block at ca. 2.88 Ga. The ca. 2.7 Ga mafic to intermediate-felsic volcanic associations of the Kitoy and Irkut terranes formed in an active continental margin setting (Turkina et al., 2012). The 2.6-2.54 Ga high-grade metamorphic rocks and granites probably formed by the Neoarchean collision of the Kitoy and Irkut terranes (Poller et al., 2005; Sal’nikova et al., 2007; Turkina et al., 2012). The Onot terrane lacks signatures of Neoarchean

magmatic and metamorphic events. The accretions of the Kitoy and Irkut high-grade metamorphic terranes with the Onot granite-greenstone terrane occurred as late as in the Late Paleoproterozoic, and the greenstone and TTG sheets were thrust over the Kitoy terrane. The 1.88-1.84 Ga episodes of metamorphism and granitic magmatism are recorded in all terranes of the Sharyzhlagay uplift (Poller et al., 2004; Turkina, Kapitonov, 2019 and references therein). 5.6.

Meso-Neoarchean intraplate magmatism

Intrusions of mafic dike swarms and sills and intraplate volcanism associated with extension/rifting provide evidence for rheologically rigid lithospheric plates. The oldest extension/rifting events have been recorded in the Kaapvaal and Pilbara Cratons and are dated as early Mesoarchean (3.2 Ga) (Misra et al., 2017; Van Kranendok et al., 2010). However, worldwide, most intracontinental extensional or rifting structures are Late Mesoarchean to Neoarchean. The oldest dike swarms in the Kaapvaal craton formed at 2.97–2.98 Ga concurrently with the Nsuze Group lavas (2967–2985 Ma) in the world’s oldest preserved rift basin (Olson et al., 2010). Three cycles of bimodal volcanism in the eastern Pilbara craton (flood basalts and subordinate felsic tuffs of the Fortescue Group) occurred in extension/rifting setting at 2772 to 2715 Ma (Blake et al., 2004). The conglomerates, quartzites, tholeiitic basalt and overlying sedimentary rocks of the Murmac Bay Group of the Rae Province (Canadian Shield) deposited or erupted between 2.77 and 2.63 Ga during the rifting of the 3.06-2.99 Ga granitoid basement (Hartlaub et al., 2004). The Murmac Bay tholeiitic basalts have fractionated multi-element patterns with clear Nb, P and Ti troughs similar to those of the Onot low-Ti metabasalts. Moreover, they follow the same crustal contamination trend from the mantle array in the Th/Yb vesus Nb/Yb diagram (Fig. 8). The plume-related events of rifting at ca. 2.88-2.82 Ga have been recorded in the La Grande Subprovince of the northeast Canadian Shield (Sappin et al., 2018). The Mesoarchean tholeiites from that area show signature of crustal contamination, while the composition of intermediate and felsic rock apparently resulted from the mixing between basaltic and crustal anatectic magmas. The Keivy terrane of the Kola Peninsula in the

Fennoscandian Shield hosts 2.86–2.68 Ga intraplate felsic metavolcanic rocks and subalkaline and alkaline granites (Mitrofanov et al., 2000; Belyaev et al., 2001; Vetrin, Rodionov, 2009). The Meso-Neoarchean events of within-plate extensional and rifting magmatism fit the available Re-Os data that constrain the stabilization of subcontinental lithosphere mantle (SCLM). The time when stable SCLM formed separately from convective mantle can be inferred from the age of Re depletion (TRD), which record the time of melting and removal of mafic melts that left a complementary buoyant depleted residue. The available data from xenoliths of mantle peridotite and sulfides and alloys from mantle xenoliths show TRD younger than 3 Ga and a pronounced peak at 2.5–2.75 Ga (Carlson, 2005; Griffin et al., 2014), indicating the formation of stable SCLM in Meso-Neoarchean time. The TRD values from the SCLM of the Siberian Craton show bimodal age distributions with both Late Archean and Paleoproterozoic peaks suggesting two stages of its formation: at 2.6-2.9 Ga and 2 Ga (Pearson et al., 1995; Ionov et al., 2015). The Archean stage of SCLM formation was coeval with the Onot bimodal volcanism. Thus, the intraplate and rift-related magmatism was coeval to SCLM stabilization and preserved a record of Meso-Neoarchean thermal and tectonic changes on the Earth (Condie, 2018). 5.7.

Typical geochemical features of Archean potassic felsic volcanics

The felsic volcanic rocks hosted by Archean greenstone belts (Lesher et al., 1986; Morris, Witt, 1997; Hollis et al., 2015; Thurston, 2015) and other Archean provinces (Zincone et al., 2016; Savko et al., 2019) are characterized by high contents of SiO2 (70-78 wt.%), K2O (≥3.5 wt.%) and Fe-index and flater REE patterns ((La/Yb)n = 3-16) with clear Eu minimums (Eu/Eu* = 0.7-0.1). In addition, those felsic volcanic rocks are strongly enriched in HFSE, Y and HREE, resulting in low Zr/Y and Sr/Y. In general, those geochemical features are typical of Phanerozoic A-type granites (Whalen et al., 1987; Eby, 1992; Frost and Frost, 2011), suggesting similar magma sources. Unlike the Archean felsic volcanics with TTG geochemical affinities, those Atype-like rocks have subchondrite Nd and Hf values, implying derivation from older crustal sources or crustal contamination of parental magma. The negative Nd and Hf values of those

rhyolites indicate reworking of older crustal components and therefore the continental setting of their formation. The highly potassium composition as well as the enrichment in Th also confirm dominantly crustal sources, e.g. tonalites or granodiorites. Experiments with those rocks produced melts with similar major element composition at low pressures (2-4 kbars) and high temperature (> 900°C) (Patiño Douce, 1997; Bogaerts et al., 2006). Experimental data are consistent with the high contents of HREE, Y, and HFSE in the volcanic rocks, whereas their ferroan affinities imply melting in dry conditions.

6.

Conclusions

The more than 80 km long Onot greenstone belt is a collage of metamorphosed volcanicsedimentary and TTG sheets thrust southwestwards onto the high-grade rocks of the Kitoy terrane. The lower greenstone unit is dominated by biotite-amphibole microgneisses with boudins and thin layers of amphibolites. The rhyolitic protoliths of the microgneiss formed at ca. 2.88 Ga and were metamorphosed at 660-690°C and 6 kbar during the Late Paleoproterozoic. The high-Fe metarhyolites have high REE and HFSE contents and are comparable with felsic volcanics of types FIIIa and FIIIb hosted by Archean greenstone belts and with A-type granites. The high- and low-Ti metabasalts (amphibolites) show geochemical signatures of subduction-related or continental flood basalts contaminated with crustal material. The Nd isotope composition of the Onot metarhyolite (Nd(t) from -3.8 to -0.8) and the wide range of Hf(t) values of zircons (from -8.5 to +1.1) agree with the melting of heterogeneous crust

consisting of Paleoarchean TTG basement rocks and juvenile material. The crustal sources melted at shallow depths (near 2-4 kbar) and temperatures higher than 900оС. The coeval formation of Onot basalts and low-P - high-T felsic melts could be triggered were probably triggered by mantle upwelling and decompressional melting in a setting of extension. Thus, the volcanism of the Onot greenstone belt was related to extension/rifting of

Paleoarchean continental crust. The extension/rifting and related bimodal volcanism suggest that the lithosphere of the Siberian Craton was already rigid in Mesoarchean time.

Acknowledgements. We thank Nikolai Rodionov (SHRIMP) and Igor Kapitonov (LA ICP MS) from the Centre of Isotopic Research of VSEGEI for analytical assistance. We would like to thank Elis Hoffmann and an anonymous reviewer for their constructive comments and highprofessional expertise, which helped us to improve the manuscript to a significant degree. We thank Inna Safonova for hot discussions and useful comments. The editorial comments from Wilson Teixeira and Elson Oliveira are greatly appreciated. This research was partially funded by the Scientific Project of IGM SB RAS and by the Russian Foundation for Basic Research (project # 20-05-00265). Contribution to IGCP 662.

References Agangi, A., Hofmann, A., Elburg, M.A., 2018. A review of Paleoarchaean felsic volcanism in the eastern Kaapvaal craton: linking plutonic and volcanic records, Geoscience Frontiers 9, 667-688. Bayanova, T.B., 2004. The Age of Reference Geologic Complexes in the Kola Region and Duration of Magmatism Processes [in Russian]. Nauka, St. Petersburg. Belyaev, O.A., Mitrofanov, F.P., Bayanova, T.B., Petrov, V.P., Levkovitch, N.V., 2001. Late Archean age of felsic metavolcanics of the Malaya Keivy region (Kola Peninsula). Doklady Earth Sciences 379, 651-645. Bibikova, E.V., Turkina, O.M., Kirnozova, T.I., Fugzan, M.M., 2006. Ancient plagiogneisses of the Onot block of the Sharyzhalgai metamorphic massif: isotopic geochronology. Geochemistry International 44, 310–314. Blake, T.S., Buick, R., Brown, S.J.A., Barley, M.E., 2004. Geochronology of a Late Archaean flood basalt province in the Pilbara Craton, Australia: constraints on basin evolution, volcanic and sedimentary accumulation, and continental drift rates. Precambrian Research 133, 143-173. Bogaerts M., Scaillet, B., Vander Auwera, J., 2006. Phase equilibria of the Lyngdal granodiorite (Norway): implications for origin of metaluminous ferroan granotoids.

Journal of

Petrology 47, 2405-2431. Bouvier, A., Vervoort, J. D., and Patchett, P. J., 2008, The Lu–Hf and Sm–Nd isotopic composition of CHUR: Constraints from unequilibrated chondrites and implications for the bulk composition of terrestrial planets. Earth and Planetary Science Letters 273, 4857. Carlson R.W., 2005. Application of the Pt-Re-Os isotopic systems to mantle geochemistry and geochronology. Lithos 82, 249-272.

Chen, Y.-X., Gao, P., Zheng, Y-F., 2015. The anatectic effect on the zircon Hf isotope composition of migmatite and associated granites. Lithos 238, 174-184. Clemens, J.D., Stevens, G., Farina, F., 2011.The enigmatic sources of I-type granite: the peritectic connexion. Lithos 126, 174-181. Condie, K.C., 2018. A planet in transition: the onset of plate tectonics on Earth between 3 and 2 Ga? Geoscience Frontiers 9, 51-60. Dall'Agnol, R., Oliveira, D.C., 2007. Oxidized, magnetite-series, rapakivi-type granites of Carajás, Brazil: implications for classification and petrogenesis of A-type granites. Lithos 93, 215-233. De Wit, M.J., Armstrong, R.A., Hart, R.J., Wilson, A.H., 1987. Felsic igneous rocks within the 3.3-3.5 Ga Barberton Greenstone Belt: high crustal level equivalents of the surrounding Tonalite-Trondhjemite terrain, emplaced during thrusting. Tectonics 6, 529–549. De Wit, M.J., Furnes, H., Robins, B., 2011. Geology and tectonostratigraphy of the Onerwacht Suite, Barberton Greenstone Belt, South Africa. Precambrian Research 186, 1–27. Donskaya, T.V., Sal'nikova, E.B., Sklyarov, E.V., Gladkochub, D.P., Mazukabzov, A.M., Kovach, V.P., Yakovleva, S.Z., Berezhnaya, N.G., 2002. Early Proterozoic postcollisional magmatism at the southern flank of the Siberian craton: new geochronological data and geodynamic implications. Doklady Earth Sciences 383, 125– 128. Eby, G.N., 1992. Chemical subdivision of the A-type granitoids: petrogenetic and tectonic implications. Geology 20, 641–644. Farmer, G.L., 2003. Continental Basaltic Rocks. in Treasure of geochemistry (Eds. Turekian K.K., Holland H.D.). Oxsford: Elsevier Ltd. 3, 85-121. Frost B.R., Barnes, C.G., Collins, W.J., Arculus, R.J., Ellis, D.J., Frost, C.D., 2001. A geochemical classification for granitic rocks. Journal of Petrology 42, 2033-2048.

Frost, C.D., Frost, B.R., 2011. On ferroan (A-type) granitoids: their compositional variability and modes of origin. Journal of Petrology 52, 39-53. Furnes, H., de Wit, M.J., Robins, B., 2013. A review of new interpretations of the tectonostratigraphy, geochemistry and evolution of the Onverwacht Suite, Barberton Greenstone Belt, South Africa. Gondwana Research 23, 403–428. Furnes, H., Dilek, Y., De Wit, M., 2015. Precambrian greenstone sequences represent different ophiolite types. Gondwana Research 27, 649–685. Goldstein, S.J., Jacobsen, S.B., 1988. Nd and Smisotopic systematics of rivers water suspended material: implications for crustal evolution. Earth and Planetary Science Letters 87, 249265. Green, M.G., Sylvester, P.J., Buick, R., 2000. Growth and recycling of early Archaean continental crust: geochemical evidence from the Coonterunah and Warrawoona Groups, Pilbara Craton, Australia. Tectonophysics 322, 69–88. Griffin, W.L., Pearson, N.J., Belousova, E., Jackson, S.E., van Achterbergh, E., O’Reilly, S.Y., Shee, S.R., 2000. The Hf isotope composition of cratonic mantle: LAM–MC–ICPMS analysis of zircon megacrysts in kimberlites. Geochemica et Cosmochemica Acta 64, 133-147. Griffin, W.L., Belousova, E.A., O’Neil, C., O’Reilly, Y., Malkovets, V., Pearson, N.J., Spetsius, S., Wilde, S.A., 2014. The world turns over: Hadean-Archean crust-mantle evolution. Lithos 189, 2-15. Guitreau, M., Blichert-Toft, J., Mattin, H., Mojzsis, S.J., Albarede, F., 2012. Hafnium isotope evidence from Archean granitic rocks for deep-mantle origin of continental crust. Earth Planetary Science Letters 337-338, 211-223. Hart, T.R., Gibson, H.L., and Lesher, C.M., 2004. Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu–Zn–Pb sulfide deposits. Economic Geology 99, 1003–1013.

Hartlaub, R.P, Heaman, L.M, Ashton, K.E, Chacko, T., 2004. The Archean Murmac Bay Group: evidence for a giant Archean rift in the Rae Province, Canada. Precambrian Research 131, 345-372. Holland, T., Blundy, J., 1994. Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry. Contributions to Mineralogy and Petrology 116, 433-447. Hollis, S.P., Yeats, C.J., Wyche, S., Barnes, S.J., Ivanic, T.J., Belford, S.M., Davidson, G.J., Roache, A.J., Wingate, M.T.D., 2015. A review of volcanic-hosted massive sulfide (VHMS) mineralization in the Archaean Yilgarn Craton, Western Australia: Tectonic, stratigraphic and geochemical associations. Precambrian Research 260, 113–135. Hornig, I. 1993. High-Ti and low-Ti tholeiites in the Jurassic Ferrar Group, Antarctica. Geologisches Jahrbuch. Reihe E: Geophysics 47, 335–369. Ionov, D.A., Carlson, R.W., Doucet, L.S., Golovin, A.V., Oleinikov, O.B., 2015. The age and history of the lithospheric mantle of the Siberian craton: Re–Os and PGE study of peridotite xenoliths from the Obnazhennaya kimberlite. Earth and Planetary Science Letters 428, 108-119. Jacobsen, S.B., Wasserburg, G.J., 1984. Sm–Nd evolution of chondrites and achondrites. Earth and Planetary Science Letters 67, 137-150. King, P.L., Chappell, B.W., Allen, C.M., White, A.J.R., 2001. Are A-type granites the hightemperature felsic granites? Evidence from fractionated granites of the Wangrah Suite. Australian Journal of Earth Sciences 48, 501–514. Kröner, A., Hegner, E., Wendt, J.I., Byerly, G.R., 1996. The oldest part of the Barerton granitoid-greenstone terrain, South Africa: evidence for crust formation between 3.5 and 3.7 Ga. Precambrian Research 78, 105–124.

Kröner, A., Hoffmann J.E., Xia H. Wu F., C. Münkerd, Hegner E., Wong J. , Wan Y., Dunyi Liu D., 2013. Generation of early Archaean felsic greenston evolcanic rocks through crustal melting in the Kaapvaal, craton, southern Africa Earth and Planetary Science Letters 381, 188–197. Laurent, O., Martin, H., Moyen, J.-F., Doucelance, R., 2014. The diversity and evolution of lateArchaean granitoids: evidence for the onset of “modern-style” plate tectonics between 3.0 and 2.5 Ga. Lithos 205, 208–235. Leak, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice, J.D., Hawthorne,F.C., Kato, A., Kisch, H.J., Krivovichev, V.G., Linthout, K., Laird, J. Mandarino, J.A., Maresch, W.V., Nickel, E.H., Rock, N.M.S., Schumacher, J.C., Smith, D.C.,

Stephenson,

N.C.N.,

Ungaretti,

L.,

Whittaker,

E.J.W.,

Youzhi,

G.

(1997).Nomenclature of Amphiboles: Report of the sub-committee on amphiboles of the international mineralogical association. Commission on new minerals and mineral names. American Mineralogist 82, 1019-1037. Lesher, C.M., Goodwin, A.M., Campbell I.H., and Gorton, M.P., 1986, Trace-element geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior Province, Canada. Canadian Journal of Earth Sciences, 23, 222–237. Ludwig, K.R., 1999. User’s manual for Isoplot/Ex, Version 2.10, A geochronological toolkit for Microsoft Excel. Berkeley Geochronology Center Special Publication. 46 p. Ludwig, K.R., 2000. SQUID 1.00, A User's Manual; Berkeley Geochronology Center Special Publication. No.2, 2455 Ridge Road, Berkeley, CA 94709, USA. Marzoli, A., Piccirillo, E. M., Renne, P. R., Bellieni, G., Iacumin, M., Nyobe, J. B., Tongwa, A. T., 2000. The Cameroon Volcanic Line revisited: petrogenesis of continental basaltic magmas from lithospheric and asthenospheric mantle sources. Journal of Petrology 41, 87–109.

Miller, C.F., McDowell, S.M., Mapes, R.W., 2003. Hot and cold granite? Implications of zircon saturation temperatures and preservation of inheritance. Geology 31, 529-532. Misra, S., Reinhardt, J., Wilson, A. H., 2017. Petrochemical evolution of the White Mfolozi Granite pluton: Evidence for a late Palaeoarchaean A-type granite from the SE Kaapvaal Craton, South Africa. Lithos 286, 480-501. Mitrofanov, F.P., Zozulya, D.P., Bayanova, T.B., Levkovitch, N.V., 2000. The world’s oldest anorogenic peralkaline granite magmatism in the Keivy structure of the Baltic Shield. Doklady Earth Sciences 374, 238-241. Molina, J.F., Moreno, J.A., Castro, A., Rodríguez, C., Fershtater, G.B., 2015. Calcic amphibole thermobarometry in metamorphic and igneous rocks: New calibrations based on plagioclase/amphibole Al-Si partitioning and amphibole/liquid Mg partitioning. Lithos 232, 286–305. Morris P.A., Witt W.K., 1997. Geochemistry and tectonic setting of two contrasting Archaean felsic volcanic associations in the Eastern Goldfields, Western Australia. Precambrian Research 83, 83-107. Nozhkin, A. D., Turkina, O. M., Mel’gunov, M. S., 2001. Geochemistry of the metavolcanosedimentary and granitoid rocks of the Onot greenstone belt. Geochemistry International 39, 27-44. Oberthür, T., Davis, D.W., Blenkinsop, T.G., and Höhndorf, A., 2002, Precise U-Pb mineral ages, Rb-Sr and Sm-Nd systematics for the Great Dyke, Zimbabwe - constraints on late Archean events in the Zimbabwe craton and Limpopo belt. Precambrian Research 113, 293−305. Olsson, J.R., Söderlund, U., Klausen, M.B., Ernst, R.E., 2010. U–Pb baddeleyite ages linking major Archean dyke swarms to volcanic-rift forming events in the Kaapvaal craton (South Africa), and a precise age for the Bushveld Complex. Precambrian Research 183, 490–500

Patiño Douce, A.E., 1997. Generation of metaluminous A-type granites by low-pressure melting of calc-alkaline granitoids. Geology 25, 743–746. Patino Douce, A.E., Johnson, A.D., 1991. Phase equilibria and melt productivity in the pelitic system: implication for the origin of peraluminous granitoids and aluminous granulites. Contributions to Mineralogy and Petrology 107, 202–218. Pearce, J.A., 2008. Geochemical fingerprinting of oceanic basalts with implications for the classification of ophiolites and search for Archean oceanic crust. Lithos 100, 14-48. Pearson, D.G., Snyder, G.A., Shirey, S.B., Taylor, L.A., Carlson, R.W., Sobolev, N.V., 1995. Archaean Re–Os age for Siberian eclogites and constraints on Archaean tectonics. Nature 374, 711–713. Peate, D. W., Hawkesworth, C. J., 1996. Lithospheric to asthenospheric transition in low-Ti flood basalts from southern Parana, Brazil. Chemical Geology 127, 1–24. Poller, U., Gladkochub, D., Donskaya, T., Mazukabzov, A., Sklyarov, E., Todt, W., 2005. Multistage magmatic and metamorphic evolution in the Southern Siberian craton: Archean and Paleoproterozoic zircon ages revealed by SHRIMP and TIMS. Precambrian Research 136, 353–368. Poller, U., Gladkochub, D., Donskaya, T., Mazukabzov, A., Sklyarov, E., Todt, W., 2004. Timing of Early Proterozoic magmatism along the Southern margin of the Siberian craton (Kitoy area). Transactions of Royal Society of Edinburgh. Earth Science 95, 215-225. Rosen, O.M., 2003. The Siberian Craton: tectonic zonation and stages of evolution. Geotectonics 37, 175–192. Rosen, O.M., Condie, K.C., Natapov, L.M., Nozhkin, A.D., 1994. Archean and Early Proterozoic evolution of the Siberian craton: a preliminary assessment. In: Condie, K. (Ed.), Archean Crustal Evolution. Elsevier, Amsterdam, pp. 411–459.

Rosen, O.M., Turkina, O.M., 2007. The oldest rocks assemblages of the Siberian craton In: M.J. Van Kranendonk, R.H. Smithies, V.C. Bennett. (Ed.), Earth’s Oldest Rocks. Amsterdam: Elsevier, pp. 793-838. Sal'nikova, E.B., Kotov, A.B., Levitskii, V.I., Reznitskii, L.Z., Mel'nikov, A.I., Kozakov, I.K., Kovach, V.P., Barash, I.G., Yakovleva, S.Z., 2007. Age constraints of high-temperature metamorphic events in crystalline complexes of the Irkut block, the Sharyzhalgai Ledge of the Siberian platform basement: results of the U–Pb single zircon dating. Stratigraphy and Geological Correlation 15, 343–358. Sappin, A.-A., Guilmerre, C., Goutier J., Beaudon, G., 2018. Geochemistry of Mesoarchean felsic to ultramafic volcanic rocks of the Lac Guyer area, La Grande Subprovince (Canada): evidence for plume-related magmatism in a rift setting. Precambrian Research 316, 83-102. Savko, K.A., Alexander V. Samsonov, A.V., Kholina, N.V., Alexander N. Larionov, A.N., Zaitseva, M.V., Korish, E.H., Bazikov, N.S., Terentiev, R.A., 2019. 2.6 Ga high-Si rhyolites and granites in the Kursk Domain, Eastern Sarmatia: petrology and application for the Archaean palaeocontinental correlations. Precambrian Research 322, 170–192. Schmidt, M.W., 1992. Amphibole composition in tonalite as a function of pressure: an experimental calibration of the Al-in-hornblende barometer. Contributions to Mineralogy and Petrology 110, 304-310. Shellnutt J.G., Zhou, M.F., 2007. Permian peralkaline, peraluminous and metaluminous A-type granites in the Panxi district, SW China: their relationship to the Emeishan mantle plume. Chemical Geology 243, 286–316. Singh, J., Johannes, W., 1996. Dehydration melting of tonalites. 2. Compositions of melts and solids. Contributions to Mineralogy and Petrology 125, 26–44.

Skjerlie, K. P. & Johnston, A. D., 1993. Fluid-absent melting behavior of an F-rich tonalitic gneiss at mid-crustal pressures: Implications for the generation of anorogenic granites. Journal of Petrology 34, 785-815. Skjerlie, K.P., Johnston, A.D., 1996. Vapour-absent melting from 10 to 20 kbar of crustal rocks that contain multiple hydrous phases: implications for anatexis in the deep to very deep continental crust and active continental margins. Journal of Petrology 37, 661–691. Smelov, A.P., Timofeev, V.F., 2007. The age of the North Asian Cratonic basement: an overview. Gondwana Research 12, 279–288. Söderlund, U., Patchett, P.J., Vervoort, J.D., Isachsen, C.E., 2004. The 176Lu decay constant determined by Lu–Hf and U–Pb isotope systematics of Precambrian mafic intrusions. Earth and Planetary Science Letters 219, 311–324. St. Seymour, K., Francis, D., 1988. Magmatic interaction between mantle and crust during the evolution of the Archean Lac Guyer greenstone belt, New Quebec. Canadian Journal of Earth Science 25, 691-700. Tang, M., Wang, X.L., Shu, X.J., Wang, D., Yang, T., Gopon, P., 2014. Hafnium isotopic heterogeneity in zircons from granitic rocks: geochemical evaluation and modeling of “zircon effect” in crustal anatexis. Earth and Planetary Science Letters 389, 188–199. Tessalina, S.G., Bourdon, B., Van Kranendonk, M., Birck, J.-L., Philippot, P., 2010. Influence of Hadean crust evident in basalts and cherts from the Pilbara craton. Nature Geoscience 3, 214-217. Thurston, P.C., 2015. Igneous rock associations 19. Greenstone belts and granite-greenstone terranes: constraints on the nature of the Archean World. Geoscience Canada 42, 437484. Turkina, O.M., 2010. Formation stages of the Early Precambrian crust in the Sharyzhalgai basement uplift, southwestern Siberian craton: synthesis of Sm–Nd and U–Pb data. Petrology 18, 158–176.

Turkina, O.M., Nozhkin, A.D., 2008. Oceanic and riftogenic metavolcanic associations of greenstone belts in the northwestern part of the Sharyzhalgai uplift, Baikal region. Petrology 16, 68–491. Turkina, O.M., Berezhnaya, N.G., Larionov, A.N., Lepekhina, E.N., Presnyakov, S.L., Saltykova, T.E., 2009. Paleoarchean tonalite-trondhjemite complex in the northwestern part of the Sharyzhalgai uplift (southwestern Siberian craton): results of U–Pb and Sm– Nd study. Russian Geology and Geophysics 50, 15–28. Turkina, O.M., Urmantseva, L.N., Berezhnaya, N.G., Skublov, S.G., 2011. Formation and Mezoarchean metamorphism of hypersthene gneiss from the Irkut granulite-gneiss block (Sharyzhalgay uplift in the southwestern Siberian craton). Russian Geology and Geophysics 52, 97-108. Turkina, O.M., Berezhnaya, N.G., Lepekhina, E.N., Kapitonov, I.N., 2012. U–Pb (SHRIMP II), Lu–Hf isotope and trace element geochemistry of zircons from high-grade metamorphic rocks of the Irkut terrane, Sharyzhalgay Uplift: implications for the Neoarchaean evolution of the Siberian Craton. Gondwana Research 21, 801–817. Turkina, O.M., Kapitonov, I.N., Sergeev, S.A., 2013. The isotope composition of Hf in zircon from Paleoarchean plagiogneisses and plagiogranitoids of the Sharyzhalgai uplift (southern Siberian craton): implications for the continental crust growth. Russian Geology and Geophysics 54, 272–282. Turkina, O.M., Sergeev, S.A., Kapitonov, I.N., 2014. The U–Pb age and Lu–Hf isotope composition of detrital zircon from metasedimentary rocks of the Onot greenstone belt (Sharyzhalgay uplift, southern Siberian craton). Russian Geology and Geophysics 55, 1249–1263. Turkina, O.M., Kapitonov, I.N., 2017. Lu–Hf isotope composition of zircon as an indicator of the sources for Paleoproterozoic collisional granites (Sharyzhalgai uplift, Siberian craton). Russian Geology and Geophysics 58, 149–164

Turkina, O.M., Kapitonov, I.N., 2019. The sources of Paleoproterozoic collisional granitoids (Sharyzhalgai Uplift, southwestern Siberian Craton): from lithospheric mantle to upper crust. Russian Geology and Geophysics 60, 414-434. Turner, S.P., Foden, J.D., Morrison, R.S., 1992. Derivation of some A-type magmas by fractionation of basaltic magma: an example from the Padthaway Ridge, South Australia. Lithos 28, 151–179. Van Kranendonk, M.J., Smithies, R.H., Hickman, A.H., Wingate, M.T.D., Bodorkos, S., 2010. Evidence for Mesoarchean (3.2 Ga) rifting of the Pilbara Craton: the missing link in an early Precambrian Wilson cycle. Precambrian Research 177, 145-161. Vander Auwera, J., Bogaerts, M., Liégeois, J.P., Demaiffe, D., Wilmart, E., Bolle, O., Duchesne, J.C., 2003. Derivation of the 1.0–0.9 Ga ferro-potassic A-type granitoids of southern Norway by extreme differentiation from basic magmas. Precambrian Research 124, 107– 148. Vetrin, V.R., Rodionov, N.V., 2009. Geology and geochronology of Neoarchean anorogenic magmatism of the Keivy structure, Kola Peninsula. Petrology 17, 537-557. Wager, L.R., Brown, G.M., 1967. Layered igneous rocks. San Francisco. 588 p. Watson, E.B., Harrison, T.M., 1983. Zircon saturation revisited: temperature and composition effects in a variety of crustal magma types. Earth and Planetary Science Letters 6, 295304. Whalen, J.B., Currie, K.L., Chappell, B.W., 1987. A-type granites: geochemical characteristics, discrimination and petrogenesis. Contribution to Mineralogy and Petrology 95, 407-419. Williams, I.S., 1998. U-Th-Pb geochronology by ion-microprobe. Reviews in Economic Geology / Eds. M.A. McKibben, III, W.C. Shanks, W.I. Ridley 7, pp. 1-35.

Yang, J.H., Wu, F.Y., Chung, S.L., Wilde, S.A., Chu, M.F., 2006. A hybrid origin of the Qianshan A-type granite, northeast China: geochemical and Sr-Nd-Hf isotopic evidence. Lithos 89, 89–106. Zincone, S.A. Oliveira, E.P., Laurent, O., Zhang, H., Mingguo Zhai, M., 2016. 3.30 Ga highsilica intraplate volcanic–plutonic system of the Gavião Block, São Francisco Craton, Brazil: Evidence of an intracontinental rift following the creation of insulating continental crust. Lithos 266–267, 414-434.

Figure captions Fig. 1. Geological sketch map of the central and eastern parts of the Sharyzhalgay uplift. Dots show sampling sites: 1 = Mt. Kamchadal, 2 = Mt. Arban, 3 = Kitoy R. Inset A: Major tectonic units of the Siberian craton, modified after (Rosen et al., 1994). Inset B: Location of terranes in the Sharyzhalgay uplift. Boxes frame areas enlarged in Fig. 2.

Fig. 2. Simplified geological map of the northwestern (A) and southeastern (B) parts of the Onot greenstone belt. Fig. 3. Field photographs of amphibole-biotite mcrogneiss with thin granite veins (a) and a thin layer of amphibolite in microgneisses (b) from outcrops on the left bank of the Kitoy River.

Fig. 4. Concordia diagram for zircon of metarhyolite and cathodoluminescence images of zircon grains and their 206Pb/207Pb ages.

Fig. 5. SiO2 vs Na2O+K2O, SiO2 vs FeO*/ (FeO*+MgO), SiO2 vs Na2O+K2O-CaO, SiO2 vs FeO*+MgO, FeO*+MgO vs Cr, and FeO*+MgO vs Sc diagrams for metavolcanic rocks of the Malaya Iret unit, Onot greenstone belt. Rhyolite1, 2, 3 are metarhyolites (microgneisses) from three areas of the Onot belt. Basalt – amphibolites of the Malaya Iret unit. Fields of ferroan (Fe) and magnesian (Mg) granites; calcic (c), calc-alkalic (c-a), alkali-calcic (a-c), and alkalic (a) granites are after (Frost et al., 2001). Experimental melts from granodiorite (GRD) (Bogaerts et al., 2006), and tonalite (To) sources (Patino Douce, 1997) are shown for comparison. The data from the Contendas rhyolites are after

(Zincone et al., 2016), the data from the Kursk rhyolites and granites are from (Savko et al., 2019).

Fig. 6. Chondrite-normalized REE patterns and PM-normalized multi-element patterns of metarhyolites and metabasalts (sample numbers as in Table 4). Fig. 7. T (Ma) - Nd(t) (a) and T (Ma) vs Hf(t) (b) diagrams for the Onot rocks. The data from rhyolite and basalt are original (Tables 5 and 6). The TTG data are from (Turkina et. al., 2013; Guitreau et al., 2012), the data from amphibolite in TTG are after (Turkina, 2010), and the data from Grt-St-Bt schist are from (Turkina et al., 2014).

Fig. 8. Th/Yb – Nb/Yb diagram (Pearce, 2008) for metabasalts of the Malaya Iret and Kamchadal units of the Onot GB. Low-Ti and high-Ti basalts from the Malaya Iret unit, upper basalts from the Kamchadal unit. The data from basalts of the Murmac Bay Group, Rae Province (Canadian Shield) are after (Hartlaub et al., 2004). The data from the Onot TTG are after (Rosen and Turkina, 2007). Arrow shows a crustal contamination trend.

Table captions

Table 1. U-Pb isotope data on zircons from metarhyolite of the Onot greenstone belt. Table 2. Composition of amphibole, biotite and plagioclase from metarhyolites. Table 3. PT estimates of metarhyolites of the Onot greenstone belt. Table 4. Whole rock major (wt. %) and trace elements (ppm) for metarhyolites and metabasalts of the Onot greenstone belt.

Table 5. Whole-rock Nd isotope data from metavolcanic and metasedimentary rocks of the Onot greenstone belt. Table 6. Lu-Hf isotope compositions of zircons from metarhyolite of the Onot greenstone belt. Declaration of interests

☒ The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

☐The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:

Bimodal volcanics of the Onot greenstone belt formed at ca. 2.9 Ga. Felsic melts were derived from Paleoarchean TTG basement and mafic sources. Mesoarchean Onot greenstone belt formed in an intraplate setting.

Table 1 U-Pb isotope data for zircon from metarhyolite of the Onot greenstone belt U, Spot

% 206Pb

c

ppm 3

Th, ppm

1

2

4

1 2

0.33 98 69 0.14 203 116

232Th 238U

5 0.72 0.59

Age, Ma

206Pb*,

ppm 6

206Pb

207Pb

238U

206Pb

7

8

D, %

238U 206Pb*

±%

9 10 11 Session 2 48.6 2927±27 2885±16 -1 1.74 1.2 99.2 2894±19 2886±10 0 1.765 0.81

207Pb* 206Pb*

12

±% 13

207Pb

235U

14

0.2074 0.99 16.4 0.2075 0.6 16.2

3 4 5 6 7 8 9 10 11 12 13 14

0.33 0.52 0.07 0.18 0.21 0.13 0.07 0.15 0.04 0.13 0.05 0.17

71 125 272 102 78 86 407 134 338 157 326 255

33 82 203 49 42 50 346 72 256 111 226 164

0.48 0.68 0.77 0.50 0.56 0.60 0.88 0.56 0.78 0.73 0.72 0.66

33.3 2789±26 55.8 2691±19 129 2838±14 45 2663±21 37 2832±28 41.1 2838±23 200 2917±14 65.8 2915±23 167 2930±15 76.1 2875±18 158 2884±15 124 2893±16

1 2 3 4 5 6 7 8 9 10 11 12

0.07 0.08 0.02 0.11 0 0 0.08 0.04 0.11 0.15 6.90 0.04

491 311 282 261 494 260 193 343 138 91 198 298

348 198 213 160 409 165 150 220 115 51 128 195

0.73 0.66 0.78 0.63 0.86 0.65 0.80 0.66 0.86 0.57 0.67 0.68

204 2543±16 151 2883±20 115 2506±18 125 2847±20 240 2890±19 123 2825±23 93.7 2878±21 166 2876±19 66 2855±23 43.5 2543±16 99.7 2883±20 142 2506±18

2796±18 0 1.847 2781±15 3 1.93 2892±8 2 1.808 2748±15 3 1.955 2886±16 2 1.813 2884±15 2 1.808 2880±7 -1 1.748 2878±12 -1 1.749 2882±8 -2 1.738 2900±11 1 1.779 2897±8 0 1.772 2877±9 -1 1.766 Session 1 2856±7 12 2.067 2886±8 0 1.772 2669±10 7 2.105 2843±11 0 1.8 2853±7 -1 1.768 2843±9 1 1.818 2885±10 0 1.776 2892±8 1 1.778 2892±11 1 1.794 2855±15 0 1.804 2792±79 1 1.831 2911±8 3 1.808

1.1 0.1963 1.1 14.6 0.87 0.1946 0.93 13 0.6 0.2082 0.51 15.8 0.96 0.1907 0.93 13.4 1.2 0.2075 1 15.7 1 0.2072 0.91 15 0.58 0.2067 0.43 16.3 0.97 0.2065 0.75 16.2 0.64 0.20691 0.46 16.4 0.78 0.2093 0.69 16.2 0.64 0.20892 0.47 16.2 0.7 0.2064 0.56 16.1

0.78 0.20374 0.4 13.5 0.85 0.2075 0.51 16.1 0.88 0.1818 0.57 11.9 0.89 0.2021 0.65 15.4 0.8 0.20325 0.42 15.8 1 0.2021 0.54 15.3 0.92 0.2073 0.61 16.0 0.83 0.20821 0.47 16.1 1 0.2082 0.71 15.9 1.1 0.2036 0.91 15.5 1.3 0.1961 4.8 14.3 0.88 0.2107 0.52 16.0

Note. Uncertainties are given at one sigma level. Pbc and Pb* indicate the common and radiogenic portions, respectively. Errors in TEMORA standard Pb calibration were 0.33% for session 2 and 0.40% for session 1. Correction for common Pb made using the measured 206Pb. Rho – error correlation between 207Pb*/235U and 206Pb*/238U.

Table 2 Composition of amphibole, biotite and plagioclase from metarhyolite of the Onot GB

SiO 2

TiO 2

Al2 O3 Fe2 O3 Fe

603 38. 90 0.9 7 10. 64 7.9 6 23.

Amphibole 2- 36- 4903 13 16 38. 40. 40. 02 77 01 1.1 1.1 0.7 0 4 0 11. 10. 10. 21 42 26 9.0 2.1 2.2 4 3 5 23. 26. 29.

4616 38. 38 0.9 5 11. 07 5.6 8 26.

603 67. 21

Plagioclase 2- 36- 4903 13 16 66.5 63. 65. 0 10 66

22.2 2

603 34. 98 3.5 5 14. 22

203 33. 61 3.7 4 14. 03

Biotite 3613 33. 96 3.2 2 15. 19

4616 64.0 2

19. 83

20.7 2

22. 25

20. 76

0.1

0.15

0.1

0.0

0.03

31.

32.

29.

4916 33. 04 2.6 2 15. 78

4616 32. 90 3.3 8 14. 72

33.

33.

O Mg O Mn O Ca O Na2 O K2 O Tot al

xMg

15 2.1 8 0.7 2 10. 88 1.6 9 1.8 7 98. 97 5.9 11 0.1 11 1.9 06 0.9 11 2.9 42 0.4 93 0.0 92 1.7 71 0.4 99 0.3 63 0.1 4

02 1.1 2 0.7 4 10. 62 1.9 3 1.9 5 98. 74 5.8 19 0.1 26 2.0 21 1.0 42 2.9 46 0.2 55 0.0 96 1.7 41 0.5 73 0.3 81 0.0 8

50 3.0 4 0.6 2 11. 10 1.4 3 1.4 1 98. 54 6.1 66 0.1 29 1.8 57 0.2 42 3.3 53 0.6 85 0.0 79 1.7 98 0.4 19 0.2 71 0.1 7

90 0.2 2 0.8 6 10. 63 1.5 3 1.4 0 97. 77 6.2 16 0.0 81 1.8 79 0.2 64 3.8 85 0.0 52 0.1 14 1.7 69 0.4 62 0.2 77 0.0 1

29 0.7 0 0.7 8 10. 85 1.5 2 1.7 1 97. 94 5.9 44 0.1 11 2.0 20 0.6 62 3.4 04 0.1 63 0.1 02 1.8 01 0.4 55 0.3 38 0.0 5

xAn

-

-

-

-

-

xAb

-

-

-

-

-

xOrt

-

-

-

-

-

Si Ti Al Fe3 +

Fe2 +

Mg Mn Ca Na K

1

0

5

55 0.4 0 3.0 3

87 0.5 6 1.6 4

27 0.3 9 3.5 5

70 0.5 2 0.5 6

86 0.5 1 0.9 9

0.0 1 9.8 1 97. 54 2.7 89 0.2 13 1.3 36

0.0 0 9.3 6 95. 87 2.7 57 0.2 31 1.3 57

0.0 0 9.9 1 95. 50 2.7 46 0.1 96 1.4 48

0.1 1 9.3 8 95. 82 2.7 20 0.1 62 1.5 31

0.0 7 9.5 1 95. 94 2.7 15 0.2 10 1.4 32

2.1 04 0.0 27 0.3 60

2.2 55 0.0 39 0.2 00

1.9 79 0.0 27 0.4 27

2.3 21 0.0 36 0.0 69

2.3 38 0.0 35 0.1 21

0.0 01 0.9 99 0.0 1

0.0 00 0.9 80 0.0 2

0.0 00 1.0 22 0.0 1

0.0 17 0.9 85 0.0 2

0.0 11 1.0 02 0.0 1

0.18

-

-

-

-

-

0.82

-

-

-

-

-

0.01

-

-

-

-

-

-

-

-

-

-

0.9 8 11. 72 0.0 7 99. 94 2.9 53

-

2.1 2 10. 81 0.1 1 99. 53 2.9 03

-

0.14 100. 45 2.91 4

3.9 1 9.9 5 0.1 0 99. 44 2.8 11

1.0 27

1.07 0

1.1 68

1.0 82

1.15 7

0.0 04

0.00 5

0.0 04

0.0 02

0.00 1

-

-

-

-

-

0.0 46 0.9 98 0.0 04

0.08 6 0.94 4 0.00 8

0.1 87 0.8 60 0.0 06

0.1 00 0.9 26 0.0 06

0.18 0 0.84 1 0.00 6

0.0 4 0.9 5 0.0 0

-

0.1 0.08 8 0.8 0.91 2 0.0 0.01 1

0.1 0 0.9 0 0.0 1

-

1.83 11.1 2

3.80 9.81 0.11 100. 02 2.82 9

Table 3. РТ estimates of biotite-amphibole microgneisses of the Onot GB Sample 6-03

Temperature (HB 94) 750

Pressure (Schm 92) 6.4

Pressure (M 2015) 5.5

2-03 7.3 5.8 36-13 685 6.2 6.1 49-16 663 6.4 8.5 46-16 690 6.9 6.2 Note. HB94 - Holland. Blundy, 1994; Schm 92 – Schmidt, 1992; M 2015 - Molina et al., 2015. Sample 6-03 and 2-03 are from north-western and other samples are from south-eastern part of Onot greenstone belt.

Table 4 Whole rock major (wt. %) and trace elements (ppm) for metarhyolites and metabasalts of the Onot greenstone belt. Sample SiO2 TiO2 Al2O3 Fe2O3t MnO MgO CaO Na2O K2O P2O5 LOI Total Mg# Th U Rb Ba Sr La Ce Pr Nd Sm Eu Gd Tb Dy Ho

2-03 6-03 1 2 71.76 70.47 0.354 0.549 11.89 11.99 6.42 7.52 0.129 0.109 0.43 0.52 1.11 1.57 3.03 2.68 3.72 4.14 0.032 0.109 0.69 0.3 99.66 100.06 12.2 13.6 2.5 2.3 199 108 712 640 130 88 73 38 125 81 15.6 9.9 59 39 10.7 8.3 2.06 1.85 11.3 8.8 1.82 1.49 11.3 9.1 2.3 2.1

1716 3 74.02 0.33 11.58 4.52 0.07 0.22 1.15 3.37 4.39 0.05 0.14 99.96 14.5 3.0 114 711 83 38 73 6.9 28 6.2 1.54 7.2 1.32 8.9 2.0

1816 4 76.42 0.31 11.51 2.16 0.03 0.97 0.59 2.76 4.79 0.03 0.23 99.94 15.4 3.4 95 840 49 44 99 10.8 40 7.9 1.90 7.9 1.41 9.3 2.1

1895 5 74.91 0.36 11.47 4.63 0.078 0.21 0.99 3.1 4.24 0.03 0.18 100.2 13.3 1.83 133 720 91 25.0 60 3.8 15.5 4.4 1.22 6.0 1.21 8.3 1.82

1995 6 74.68 0.33 11.18 4.71 0.074 0.1 0.94 3.48 4.24 0.03 0.12 99.9 13.4 2.0 162 826 90 41 113 9.7 38 8.4 1.72 8.6 1.49 9.8 2.1

2695 34-13 7 8 75.53 72.55 0.35 0.42 11.59 11.39 4.21 5.94 0.077 0.07 0.19 0.67 0.93 1.68 2.78 2.51 4.08 4.13 0.03 0.07 0.5 0.50 100.3 100.09 13.5 13.4 2.2 1.58 150 83 731 936 90 112 39 69 58 136 8.4 15.7 32 61 6.0 11.1 1.35 2.7 5.5 11.0 0.88 1.66 5.9 9.5 1.41 1.93

3613 37-13 9 10 72.92 74.75 0.52 0.34 11.74 11.09 4.56 5.33 0.06 0.08 0.40 0.04 1.46 1.01 2.85 3.13 3.87 4.04 0.10 0.03 0.61 0.24 99.21 100.18 16.6 12 1.67 1.9 126 137 751 758 90 90 50 63 101 123 12.1 15 46 52 8.4 9.6 1.83 2.1 7.6 9.8 1.19 1.6 7.4 9.9 1.51 2.0

3813 11 75.90 0.33 10.81 4.06 0.06 0.08 0.97 2.98 4.23 0.03 0.31 99.89 13 1.9 167 756 82 22 70 5.7 22 5.6 1.4 7.1 1.3 8.4 1.9

Er 7.1 6.0 6.2 6.2 5.6 6.1 4.9 5.1 4.5 5.8 5.7 Tm 0.99 0.85 0.93 0.93 0.87 0.95 0.79 0.75 0.70 0.94 0.88 Yb 6.8 5.6 6.2 6.1 5.4 6.3 5.3 4.8 4.5 5.9 5.4 Lu 1.12 0.85 0.88 0.91 0.82 0.90 0.79 0.69 0.70 0.85 0.83 Zr 607 357 397 417 403 416 388 468 525 383 396 Hf 10.6 9.1 10.3 10.6 10.0 9.7 9.3 11.7 13.2 9.7 10.0 Ta 1.46 1.32 1.41 1.44 1.06 1.32 1.35 0.6 1.1 1.2 1.1 Nb 26 17.4 18.8 19.8 17.9 19.7 18.7 15.8 18.0 18.5 18.9 Y 89 55 55 58 50 59 39 52 42 56 51 Cr 106 7.9 38 41 90 45 Ni 30 6.2 12 15 56 48 Sc 5.5 5.2 4.3 4.2 (La/Yb)n 7.3 4.6 4.2 4.9 3.1 4.4 5.0 9.8 7.5 7.2 2.8 * Eu/Eu 0.6 0.7 0.7 0.7 0.7 0.6 0.7 0.7 0.7 0.6 0.7 Т, оС 888 824 828 852 849 833 847 849 868 832 835 Note. Fe2O3t – total Fe; Eu/Eu* = Eun/((Smn+Gdn)0.5); n - chondrite-normalized ratio. Т,оС – temperature of zirconium saturation after (Watson and Harrison, 1983). 1-15 – metarhyolites: 14 –north-western. 5-7 - central and 8-15 – south-eastern parts of GB; 16-22 – metabasalts.

Table 4 (continued). Sample SiO2 TiO2 Al2O3 Fe2O3t MnO MgO CaO Na2O K2O P2O5 LOI Total Mg# Th U Rb Ba Sr La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Zr Hf Ta Nb Y Cr Ni Sc (La/Yb)n Eu/Eu* Т, оС

3913 12 74.92 0.34 11.08 4.49 0.08 0.11 0.71 2.90 4.00 0.03 0.53 99.33 14.6 2.1 147 713 71 40 102 9.5 36 7.3 1.79 7.0 1.13 7.5 1.70 5.5 0.86 5.6 0.84 404 10.5 1.2 18.0 48 41 10 3.6 4.8 0.8 849

5216 13 75.23 0.36 11.45 4.43 0.06 0.20 0.98 3.12 3.82 0.04 0.08 99.88 13.5 1.7 90 725 93 33 90 8.2 30 6.3 1.58 6.7 1.13 7.6 1.71 4.9 0.73 4.4 0.66 386 10.0 0.8 15.7 44 5.0 0.7 840

4916 14 74.77 0.35 11.34 4.24 0.07 0.10 1.08 2.86 4.42 0.04 0.13 99.51 14.2 1.52 105 745 88 41 93 10.2 38 8.3 1.70 8.5 1.38 9.2 1.91 5.6 0.81 5.2 0.76 392 9.7 0.9 16.8 52 5.3 0.6 836

4616 4-03 15 1 74.91 54.03 0.30 0.66 10.99 12.23 3.86 12.81 0.06 0.219 0.16 7.59 1.32 9.21 2.58 1.85 4.65 1 0.04 0.067 0.04 0.4 99.02 100.07 54 15.0 3.3 1.8 0.7 113 45 663 330 68 222 73 15 133 29 16.8 3.9 63 14 11.8 3.5 2.0 0.90 11.6 4.2 1.85 0.65 11.4 4.5 2.4 0.98 6.6 2.7 1.01 0.41 6.0 2.8 0.88 0.41 472 105 12.0 3.3 1.4 0.41 19.4 5.5 65 24 8.2 3.7 0.5 0.7 856 -

5-03 2 55.4 0.65 13.04 11.98 0.184 4.98 8.23 2.15 2.17 0.07 0.36 99.22 45 4.8 1.40 85 494 157 14.1 29 3.7 15.0 3.7 1.00 3.9 0.69 4.5 1.02 3.0 0.48 3.0 0.46 158 3.8 0.55 7.7 28 167 89 3.2 0.8 -

2195 3 49.99 0.73 13.23 11.41 0.207 8.18 12.24 1.72 0.34 0.045 0.2 98.29 59 0.98 0.41 19 106 127 4.2 10.5 1.5 6.4 2.0 0.73 2.8 0.49 3.3 0.73 2.2 0.33 2.1 0.33 43 1.50 0.2 2.8 19 167 81 1.3 0.9 -

2595 4 48.92 1.22 17.74 13.03 0.18 3.97 10.68 2.85 0.55 0.11 0.52 99.87 38 1.72 0.41 14.8 94 162 5.7 14.4 2.1 9.9 3.0 1.10 3.6 0.66 4.3 0.96 2.9 0.42 2.7 0.43 89 2.2 0.26 4.4 27 62 23 1.4 1.0 -

4516 51-16 5 6 50.94 47.85 0.83 2.96 14.50 12.40 13.39 19.31 0.18 0.26 6.19 4.71 9.09 8.62 2.83 2.43 0.90 0.86 0.10 0.34 0.44 0.17 99.50 100.05 48 33 1.5 4.6 1.2 1.4 30 29 141 156 166 120 11.1 14.3 24 37 3.5 5.5 14.9 27 3.6 7.3 0.97 2.3 4.2 9.1 0.77 1.60 5.2 10.1 1.11 2.2 3.2 6.2 0.48 0.90 3.2 5.7 0.46 0.85 97 270 2.6 6.9 0.60 0.69 8.3 10.0 31 59 3.7 1.7 0.7 0.8 -

4716 7 52.73 2.32 12.31 15.86 0.15 2.97 9.62 2.12 0.92 0.23 0.36 99.72 27 7.5 2.1 22 65 170 31 55 7.9 33 7.6 2.1 8.6 1.35 8.6 1.82 5.5 0.81 5.1 0.79 203 5.3 0.83 12.9 51 77 11.0 4.1 0.8 -

Table 5. Whole-rock Nd isotope data from metavolcanic rocks of the Onot greenstone belt. №

Sample

t, Ma

1 2 3 4 5 6 7 Note.

Nd

Sm

147Sm/144Nd

143Nd/144Nd

TNd(DM),

Nd(t)

Ma

ppm

2-03 2880 68.14 12.97 6-03 2880 36.24 7.95 19-95 2880 33.02 7.33 34-13 2880 58.43 11.4 36-13 2880 43.37 8.42 20-95 2880 5.26 1.63 80-95 2880 6.85 2.13 1-5 – metarhyolites. 6-7 – metabasalts.

0.11536 0.13264 0.13415 0.1179 0.1174 0.18718 0.18749

0.511050±9 0.500227±25 0.511408±20 0.511056±5 0.511033±6 0.512670±19 0.512715±16

3223 3589 3316 3309 3328 -

-1.3 -3.8 -0.8 -1.6 -1.9 4.2 4.7

Table 6. Lu-Hf isotope composition of zircons from metarhyolite of the Onot GB.

Spot

36-13_2 36-13_4 36-13_5 36-13_6 36-13_7 36-13_8 36-13_9 36-13_10 36-13_11 36-13_12 36-13_13 36-13_14

t,

176Yb

176Lu

176Hf

Ma

177Hf

177Hf

177Hf

2886

0.01706

0.001589

2781

0.01808

0.000872

2892

0.03866

0.001834

2748

0.01277

0.000894

2886

0.01218

0.000979

2884

0.02052

0.001599

2880

0.03422

0.002530

2878

0.01039

0.000728

2882

0.02508

0.002282

2900

0.02223

0.001724

2897

0.02158

0.001722

2877

0.02723

0.002967

±

0.280968 0.000056 0.280931 0.000084 0.281054 0.000070 0.280868 0.000044 0.280926 0.000036 0.280909 0.000031 0.280882 0.000040 0.280873 0.000033 0.280895 0.000039 0.280838 0.000034 0.280953 0.000040 0.280856 0.000042

εHf (t)

±

TCHf (DM), Ma

-1.6

2.0

3413

-3.9

3.0

3470

1.1

2.5

3255

-7.0

1.6

3627

-1.9

1.3

3430

-3.8

1.1

3543

-6.6

1.4

3711

-3.5

1.2

3519

-5.7

1.4

3654

-6.2

1.2

3699

-2.1

1.4

3455

-8.5

1.5

3821