Petrogenesis of siliceous high-Mg series: Evidence from Early Paleoproterozoic mafic volcanic rocks of the Vodlozero Domain, Fennoscandian Shield

Accepted Manuscript Petrogenesis of siliceous high-Mg series: Evidence from Early Paleoproterozoic mafic volcanic rocks of the Vodlozero Domain, Fennoscandian Shield M. Bogina, V. Zlobin, S. Svetov, E. Sharkov, A. Chistyakov PII:

S1674-9871(17)30047-6

DOI:

10.1016/j.gsf.2017.02.009

Reference:

GSF 545

To appear in:

Geoscience Frontiers

Received Date: 19 October 2016 Revised Date:

10 February 2017

Accepted Date: 21 February 2017

Please cite this article as: Bogina, M., Zlobin, V., Svetov, S., Sharkov, E., Chistyakov, A., Petrogenesis of siliceous high-Mg series: Evidence from Early Paleoproterozoic mafic volcanic rocks of the Vodlozero Domain, Fennoscandian Shield, Geoscience Frontiers (2017), doi: 10.1016/j.gsf.2017.02.009. This is a PDF file of an unedited manuscript that has been accepted for publication. As a service to our customers we are providing this early version of the manuscript. The manuscript will undergo copyediting, typesetting, and review of the resulting proof before it is published in its final form. Please note that during the production process errors may be discovered which could affect the content, and all legal disclaimers that apply to the journal pertain.

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Petrogenesis of siliceous high-Mg series: Evidence from Early Paleoproterozoic mafic

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volcanic rocks of the Vodlozero Domain, Fennoscandian Shield

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M. Bogina a,*, V. Zlobin b, S. Svetov c, E. Sharkov a, A. Chistyakov a

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a

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Academy of Sciences, Staromonetny per. 35, Moscow, 119017 Russia

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Institute of Geology of Ore Deposits, Petrography, Mineralogy, and Geochemistry, Russian

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b

Geological Institute, Russian Academy of Sciences, Pyzhevskii per. 7, Moscow, 109017 Russia

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Institute of Geology, Karelian Scientific Center, Russian Academy of Sciences, Pushkinskaya

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ul. 11, Petrozavodsk, 185610 Russia

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* Corresponding author email address: [email protected];[email protected]

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Abstract

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Compositional peculiarities of the siliceous high-Mg series (SHMS) rocks formed at the

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Archean–Paleoproterozoic boundary as a function of plume activity are discussed using example

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of Early Paleoproterozoic mafic volcanic rocks of the Vodlozero Domain, Fennoscandian Shield.

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These rocks are characterized by wide variations in Mg# (from 33 to 67) and Cr contents (from

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25 to 1123 ppm), LREE enrichment, and weakly negative εNd (from –0.7 to –2.9). The high

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Gd/Yb ratio in the primitive high-Mg rocks of the Vodlozero Domain suggests their generation

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from a garnet-bearing source. At the same time, their negative εNd in combination with LREE

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enrichment points to the crustal contamination. A new model was proposed to explain the

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remarkable global-scale similarity of SHMS. Such rocks can be generated by the contamination

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of a high-degree (30%) partial melt derived from a depleted mantle. The lower crustal

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sanukitoid-type rocks can be considered as a universal crustal contaminant. Modeling showed

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that such mixing can provide the observed narrow εNd variations in Early Paleoproterozoic

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volcanics. The Neoarchean sanukitoid suites which are widespread on all cratons, presumably

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composed the lower crust at the beginning of the Paleoproterozoic. Therefore, this mechanism

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can be considered universal for the genesis of the SHMS rocks. The high- to low-Cr rock series

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can be produced by the fractionation of the mafic melt coupled with an insignificant crustal

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assimilation of felsic end members of the sanukitoid suite of the Vodlozero Domain en route to

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the surface, as suggested by the positive correlation of εNd with Cr and Mg#, negative correlation

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with Th, and slight decrease of εNd in the more evolved varieties.

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Keywords: Early Palaeoproterozoic, siliceous high-magnesian series (SHMS) rocks, sanukitoids,

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mantle plume, depleted mantle, crustal contamination

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1. Introduction The Early Paleoproterozoic stage in the Earth's evolution was marked by the initiation of

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the global rift systems caused by a superplume event and the formation of the large igneous

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provinces (LIPs) consisting of the mantle-derived dike swarms, mafic–ultramafic layered

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intrusions, and volcanic fields (e.g., Sharkov and Bogatikov, 2010; Ernst, 2014).

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These LIPs are made up of derivatives of the siliceous high-Mg series (SHMS) magmas

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formed in an intracontinental setting during the incipient rifting (Sharkov and Bogatikov, 2010;

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Bogina et al., 2015; Mints et al., 2015). These rocks are mainly developed at or near the

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Archean–Paleoproterozoic boundary (ca. 2.5–2.4 Ga). The best preserved components are dike

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swarms (likely volcanic pathways) and layered mafic–ultramafic intrusions that frequently host

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PGE, Cu, Ni, and Cr deposits (Alapieti et al., 1990; Cawthorn, 1996; Sharkov and Bogina, 2006;

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Mitrofanov et al., 2013), whereas volcanic fields are preserved only as scarce fragments. The

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SHMS rocks are characterized by peculiar chemical signatures, which are atypical of the

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younger LIPs. They resemble boninites in their high MgO, Cr, and low TiO2 contents at

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relatively high silica content. At the same time, unlike Phanerozoic boninites, the SHMS rocks

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lack the U-shaped REE pattern and positive Zr anomalies. Instead, they carry the typical

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“crustal” (or slab-related) trace-element and isotopic signatures, such as LILE and LREE

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enrichment, negative Nb and Ti anomalies, and usually weakly negative εNd values against the

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background of elevated Cr, Cu, Ni, and V contents (Sun et al., 1989; Vrevsky, 2011; Hanski et

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al., 2012; Ketchum et al., 2013; Bogina et al., 2015). Therefore, they are often defined as the

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boninite-like rocks (Sharkov et al., 2005).

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One of the Paleoproterozoic LIPs was formed about 2.45 Ga in the Fennoscandian and

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Canadian shields assembled in the Kenorland continent at that time (Vogel et al., 1998; Bleeker

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et al., 2008; Eriksson and Condie, 2014). In the Fennoscandian Shield, these rocks are developed

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in the Kola and Karelian cratons. As other LIPs, their main components are observed as

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numerous dike swarms (Vuollo and Huhma, 2005; Arzamastsev et al., 2009), small coronitic

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mafic-ultramafic bodies (drusites) scattered in the Belomorian mobile belt (Sharkov et al.,

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2004), layered intrusions (Alapieti et al., 1990; Amelin et al., 1995), and volcanic piles (Svetov

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et al., 2004; Lauri et al., 2012). Layered intrusions of the Fennoscandian Shield are subdivided

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into two age groups. The first group includes the Monchegorsk Complex, Fedorova-Pana

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Tundra, and Mt. General’skaya massifs (Fig. 1) with ages varying within 2520–2490 Ma

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(Amelin et al., 1995; Bayanova, 2004; Bayanova et al., 2010). The second group with an age of

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about 2450–2430 Ma includes the Portimo, Kemi, Burakovsky, Oulanka Group, and other

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massifs (Fig. 1; Amelin et al., 1995). Paleoproterozoic SHMS rocks in the Canadian Shield are as follows: the Thessalon

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bimodal volcanic association of the Huronian Supergroup (Jolly, 1987, 1992; Ketchum et al.,

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2013); the Matachewan dike swarm with age within 2473–2446 Ma (Heaman, 1997), the Spi

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Group volcanic rocks and Kaminak dikes in the West Churchill Province, Hearne Domain

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(Sandemann et al., 2008, 2013) and others. The SHMS rocks of similar age also compose LIPs in

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Western Australia (Woongarra event at 2452–2446 Ma in the Hamersley Basin, Barley et al.,

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1997), Antarctica (ca. 2.4 Ga old Westfold Hills, e.g., Collerson and Sheraton, 1986) and India

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(up to 2.5 Ga volcanic rocks of the Dongargarh Supergroup; Sensarma et al., 2002).

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The origin of the SHMS rocks is hotly debatable. Some researchers believe that their

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genesis is related to melting of metasomatized subcontinental lithospheric mantle (Cadman et al.,

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1997, 2008; Lauri et al., 2012; Ketchum et al., 2013). Others suggest their derivation from

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depleted mantle plume material (high-Mg komatiitic or picritic melts) with the subsequent

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crustal contamination and fractionation (Amelin et al., 1995; Puchtel et al., 1996, 2016; Sharkov

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et al., 1997; Yang et al., 2016) or the involvement of both these mechanisms (Amelin et al.,

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1995). At the same time, in spite of the obvious geochemical and isotope-geochemical similarity

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of the SHMS rocks, variable proportions of the crustal contaminant of different composition

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have been proposed to explain the composition of these rocks. In particular, Puchtel and co-

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authors suggested that the Early Paleoproterozoic volcanic rocks of the Vetreny Belt were

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formed when the strongly LREE-depleted parental komatiitic melt assimilated 4% tonalite of the

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host ca. 3.2 Ga Vodlozero Domain (Puchtel et al., 1997, 2016). A crustal contamination range

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between 4 and 15% was assumed in (Amelin et al., 1996) to explain the Nd isotope composition

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of the Burakovka Massif also located in the Vodlozero Domain. Minor crustal contamination

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with occasional involvement of the asthenospheric material was proposed for other Early

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Paleoproterozoic layered intrusions of the Fennoscandian Shield (Amelin et al., 1996).

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Insignificant crustal contamination (ca. 2%) of komatiitic melt derived by melting of the

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enriched mantle was assumed as possible mechanism for the formation of Early Paleoproterozoic

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supracrustal mafic rocks of the Arvarench structure in the Kola Craton (Fig. 1; Vrevsky, 2011).

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At the same time, it should be taken into account that the Archean basement of both

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Fennoscandian and Canadian shields is very heterogeneous in composition and age and consists

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of domains with different crustal histories, which are not generally consistent with the

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manifestation of such similar global-scale geochemical signatures. We propose another universal

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mechanism to explain the formation of Early Paleoproterozoic SHMS rocks, with volcanic rocks

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of this age in the eastern Fennoscandian Shield as example. As compared to Early

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Paleoproterozoic dike swarms and layered intrusions, the volcanic rocks under consideration are

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preserved only locally and less studied. However, they bear the important parental melt-related

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information, which is difficult to obtain by studying, for instance, highly evolved dikes or

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layered intrusions, because their long-term evolution in an intermediate chamber can

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significantly modify and obliterate the parental magma composition. In this paper, we attempt to

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resolve this problem by detailed geochemical and isotope-geochemical study of Early

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Paleoproterozoic volcanic rocks developed in the Vodlozero Domain (Karelian Craton).

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2. Geology and age of the Early Paleoproterozoic SHMS volcanic rocks of the

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Fennoscandian Shield

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Early Paleoproterozoic volcanic rocks are developed in the eastern Fennoscandian Shield

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(Kola and Karelian cratons). The Kola Craton is a mosaic of Mesoarchean and Neoarchean units.

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The Archean growth of this province occurred from 2.92 to 2.68 Ga (Vetrin et al., 2013). The

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Karelian Craton is commonly subdivided into the Western Karelian, Central Karelian, and

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Vodlozero domains with different ages of the crust formation (Lobach-Zhuchenko et al., 2000).

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The Vodlozero and Western Karelian domains are made up of the Mesoarchean (3.5–3.0 Ga)

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tonalite–trondjemite–granodiorite (TTG) gneiss cores surrounded by the 3.0–2.8 Ga tonalite–

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greenstone belts. They are separated by the Central Karelian Domain dominated by the juvenile

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Neoarchean granitoids, which compose the giant Pyaozero batholiths surrounded by relicts of the

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tonalite–greenstone belts (e.g., Lobach-Zhuchenko et al., 2005; Larionova et al., 2007). The

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Neoarchean granitoids are ascribed to the diorite–monzodiorite–granodiorite series collectively

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referred to as sanukitoid suites and emplaced at the end of the Archean (ca. 2.7–2.6 Ga)

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worldwide (Shirey and Hanson, 1984; Stern et al., 1989; Stern and Hanson, 1991; Larionova et

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al., 2007; Lahtinen, 2008; Heilimo et al., 2011). The sanukitoids were also recognized in the

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Western Domain (plutons in the Suomussalmi and Kostomuksha areas). In the Vodlozero

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Domain, the rocks of the sanukitoid series are widespread in the Bergaul area, where they

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compose differentiated diorite–granodiorite (± granite) massifs that intrude the

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volcanosedimentary sequences of greenstone belts (Larionova et al., 2007).

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Early Paleoproterozoic volcanic rocks compose thick volcanic sequences (total thickness

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up to 5.5 km) in the lower portions of the Pechenga–Varzuga structure, Kola Craton; the Shomba

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and Lekhta structures, East Karelian Belt, Central Domain, Karelian Craton (Fig. 1); numerous

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small structures (Elmus, Koikary, Krasnaya Rechka and Semch) and the Vetreny Belt in the

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oldest Vodlozero Domain (Karelian Craton); and other structures, such as the Salla, Perapohja,

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Lapland belts, and Paanajarvi in the Western and Central domains of the Karelian Craton (Fig.

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1). According to the regional stratigraphic scheme, Early Paleoproterozoic SHMS rocks are

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ascribed to the Sumian (2505–2430 Ma) and Sariolian (2430–2300 Ma) systems (Hanski, 2012). Minor relicts of the primary magmatic minerals survived intense secondary alteration and

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low-grade metamorphism only in the Vetreny Belt. Rocks from other structures contain only

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metamorphic minerals. However, volcanic textures have been retained elsewhere, for instance, as

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spinifex texture in the lavas of the Vetreny Belt, as pillowed, amygdaloidal, and variolitic

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textures in other localities (Puchtel et al., 1997; Svetov et al., 2004; Golubev et al., 2011; Bogina

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et al., 2015). The most complete Early Paleoproterozoic sequence was identified in the Shomba

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and Lekhta structures in the Central Domain (Karelian Craton) and was considered in detail in

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(Bogina et al., 2015). Early Paleoproterozoic mafic rocks are frequently associated with the

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minor felsic “quartz porphyry” rocks, except for volcanic rocks of the Vodlozero Domain.

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The age of the Sumian–Sariolian (Early Paleoproterozoic) volcanic rocks of the

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Fennoscandian Shield remained undetermined for a long time. Their upper age limit was mainly

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constrained by the age of the overlying felsic rocks at 2442 ± 15 Ma (Zlobin et al., 2003),

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whereas their lower age boundary remained unclear and the rocks were sometimes ascribed to

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the Archean (Levchenkov et al., 2001). Exceptions are volcanic rocks of the Vetreny Belt, which

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were dated by Sm–Nd isochron method at 2410 ± 34 Ma and 2449 ± 35 Ma and by U–Pb zircon

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method at 2437 ± 3 Ma (Puchtel et al., 1996, 1997). Recent Re-Os data more accurately

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constrained the age of the Vetreny Belt volcanics to 2407±6 Ma (Puchtel et al., 2016), which

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corresponds to the Sariolian. Their specific geological position, some compositional

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peculiarities, including komatiitic affinity, unusually fresh appearance, and different structures of

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the sequence complicate their direct correlation with other Early Paleoproterozoic volcanic rocks

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of the Karelian Craton.

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More accurate geological and geochronological constraints are available for the Sumian

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volcanic rocks of the Imandra–Varzuga Belt, where basal conglomerates of the Paleoproterozoic

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sequence rest directly on the Monchegorsk layered complex dated at 2504 ± 1.5 Ma (Amelin et

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al., 1995; and others), thus limiting the lower age boundary. With progress in the U–Pb SHRIMP

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(sensitive high-resolution ion microprobe) zircon dating, several dates were obtained directly for

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the Sumian and Sariolian mafic rocks of the Karelian and Kola cratons. In particular, andesites of

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the Arvarench structure previously ascribed to the Archean were dated by the U–Pb method at

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2424 ± 12 Ma (Vrevsky, 2011). Two age values were obtained by U–Pb method for basaltic

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andesites of the Shomba structure: 2423 ± 31 Ma (Zlobin et al., 2010) and 2439 ± 21 Ma

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(Myskova et al., 2013). The Sariolian basalts of the Lekhta structure were dated by U–Pb zircon

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method at 2412 ± 17 Ma (Myskova et al., 2012). Unfortunately, no age constraints are available

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for the Sumian rocks of the Vodlozero Domain. All attempts to date these rocks remained

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unsuccessful. At present, they are arbitrarily ascribed to the Sumian–Sariolian system based on

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their compositional correlation with the volcanic rocks of the Vetreny Belt and other structures

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(Hanski, 2012).

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3. Geological background and Petrography Early Paleoproterozoic (Sumian) mafic to intermediate rocks are widespread in the

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Koikary, Elmus, Semch, Kumsa, and Krasnaya Rechka structures in the Vodlozero Domain,

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where they form volcanogenic sequences made up of the alternation of lava flows and tuffs up to

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1500 m thick.

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The stratified lava sequences unconformably overlie the Meso–Neoarchean complexes

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represented by the komatiite–basaltic and basalt–andesite–dacite-rhyolite volcanic associations

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and Archean granitoids of different affinities (Svetov et al., 2004). The Sumian rocks are divided

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into the Glubokozero and Kumsa formations (Golubev et al., 2011). Each formation consists of

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the lower sedimentary and the upper volcanic units. In the Semch and Koikary structures, the

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base of the sequence is represented by eluvium–talus breccias. The volcanic sequence rests on

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the sedimentary unit (150 m thick) made up of siltstones, sandstones, and tuffites in the Kumsa

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structure and overlies the weathered rocks after granites in the Krasnaya Rechka structure. The

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sequences are overlain by the Sariolian polymictic conglomerates with interbeds of sandstones

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and gravelstones. The sequences are mainly made of massive, amygdaloidal, and less common

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pillow lava flows and piles (9–12 to 40–45 m thick) intercalated with thin (0.4–4.5 m) tuff

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horizons (agglomerate-size).

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The inner stratification of the rock association is similar in all studied structures.

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However, the degree of preservation of the reconstructed sections is sharply different. In

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particular, lava sequences are represented by 16 flows with a total thickness of 630 m in the

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Koikary structure, 20 flows (700 m thick) in the Elmus structure, 21 flows (650–800 m thick) in

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the Semch structure, 35 lava flows (1200 m thick) in the Kumsa structure, and 12 flows forming

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a 270-m-thick volcanic sequence in the Krasnaya Rechka structure. Their correlated sections are

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presented in (Melezhik et al., 2012).

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The lavas are characterized by different degrees of gas saturation, which is reflected in the

variable proportion of amygdules in the flows and foam at the roof.

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The Sumian rock association was metamorphosed under green-schist facies conditions.

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The mineral assemblage of the association is composed of secondary minerals represented

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mainly by albite, albite–oligoclase, quartz, actinolite, hornblende, chlorite, epidote, biotite, and

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accessories (magnetite, titanite, pyrite, chalcopyrite, and hematite). The rocks are marked by the

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relict diabase, porphyritic, variolitic, amygdaloidal, and brecciated textures (Fig. 2). Plagioclase

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is represented by the narrow prismatic laths (0.2–0.3 to 0.6–0.8 mm in size) with scarce

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polysynthetic twins, which are almost completely replaced by sericite, chlorite, and epidote.

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4. Samples and methods

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Samples used in this study were collected from the Krasnaya Rechka, Elmus, Koikary, and Kumsa structures in the Vodlozero Domain of the Karelian craton (Fig. 1). The major-element composition was analyzed by X-ray fluorescence on a PW-2400 Philips (Analytical BV) spectrometer at the Institute of Geology of Ore Deposits, Petrography,

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Mineralogy, and Geochemistry, Russian Academy of Sciences (IGEM RAS) (analyst A.I.

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Yakushev). The measurement accuracy was 1–5 rel.% for concentrations more than 0.5 wt.% and

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up to 12 rel.% for concentrations below 0.5 wt.%.

The trace element composition of rocks was also analyzed at the IGEM RAS (analyst

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Ya.V. Bychkova) on an ionization X-Series II ICP–MS mass spectrometer. For measurements,

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samples were digested in a mixture of acids (HF, HNO3, HCl) in closed beakers using a

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Milestone microwave and were conserved by HNO3 (3%). Sensitivity of the device over the

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entire measurement period was controlled using 68-element standard certified solutions (ICP–

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MS-68A, HPS, solutions A and B), which included all elements analyzed in the samples. The

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detection limit of elements was from 0.1 ng/g for heavy and medium weight elements to 1 ng/g

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for light elements. The measurement error was 1–3 rel.%. Element concentrations were

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calculated using calibration solutions prepared from the standard solution ICP–MS-68A, HPS (A

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and B) with concentrations 0.03–10 ppb.

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The Nd isotope composition was analyzed at the Geological Institute, Kola Science

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Centre, Russian Academy of Sciences on a seven-channel Finnigan MAT-262 (PRQ) mass

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spectrometer using a technique described in (Bayanova, 2004). The reproducibility of

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measurements for the isotopic ratios of the La Jolla standard, 143Nd/144Nd = 0.511857 ± 6 (n =

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11) was no higher than 0.002% (2σ). The uncertainty of the 147Sm/144Nd ratio was calculated

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from seven measurements of the BCR-2 standard as 0.2% (2σ). The laboratory blank was 0.3 ng

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for Nd and 0.06 ng for Sm. The Nd isotope ratios were normalized to 146Nd/144Nd = 0.7219 and

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corrected for mass fractionation based on analyses of the la Jolla Nd standard. The model age

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was calculated with respect to the depleted mantle with an age of 4.56 Ga using current

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characteristics: 147Sm/144Nd = 0.2135, 143Nd/144Nd = 0.513151 (Goldstein and Jacobsen, 1988).

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The Nd isotope composition was also analyzed at the Center for Isotope Research of the

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Karpinskii All-Russia Geological Research Institute on a Triton (Thermo) solid phase

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multichannel high-resolution mass spectrometer in a static mode. The isotope composition of Nd

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was corrected for mass fractionation relative to 146Nd/144Nd = 0.7219 accepted for natural

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compositions. The value of 143Nd/144Nd in the JNdi-1 international isotope standard during

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measurements was 0.512126 ± 2 (n = 25). Laboratory blank was less than 0.01 ng for Nd and

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Sm. The chemical separation of elements, acid digestion of silicate samples, and data processing

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are described in (Krivolutskaya et al., 2010).

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5. Geochemistry and Nd isotope composition Major and trace element contents are listed in Table 1. Due to the possible mobility of

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alkalis and silica, the samples were classified using immobile elements in the Zr/TiO2–Nb/Y

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diagram (Winchester and Floyd, 1977, Fig. 3). The rocks correspond to basalt, subalkaline

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basalt, basaltic andesite, and andesite, falling between the tholeiitic and alkaline series. The Mg#

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parameter [100×Mg/(Mg + Fe2+)] shows wide variations between 67 and 33. High Mg# (>60)

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rocks, which correspond in composition to high-Mg basalts, are close to the parental melts and

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can be used in genetic considerations. They are also characterized by extremely high Cr contents

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(up to 1123 ppm). Compared to the Cr-depleted low-Mg varieties, the high-Cr high-Mg rocks

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(Mg# > 60) are characterized by lower Zr, Y, Nb, Th, U (Fig. 4), and higher Cr, Co, and Ni

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contents. The total REE content in the high-Cr varieties is slightly lower. Their REE patterns are

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usually less fractionated, but sometimes overlap partly with those of low-Cr rocks: (La/Yb)n

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(subscript n represents chondrite normalization) up to 7.5, (La/Sm)n = up to 2.8, (Gd/Yb)n =

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around 2.0 against (La/Yb)n = 5.44–12.34, (La/Sm)n = 2.03–4.4, and (Gd/Yb)n = 1.36–2.71) (Fig.

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5). The high-Mg varieties, which were found mainly in the Elmus and to a lesser extent in the

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Koikary structures (Table 1), are similar to rocks of the Vetreny Belt.

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Other Early Paleoproterozoic mafic rocks of the Karelian Craton (Paanajarvi, Lekhta, and

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Shomba structures) are usually represented by the highly fractionated rocks (Mg# <55) and

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cannot be used for genetic considerations (Buiko et al., 1995; Bogina et al., 2015).

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Variations of the major and minor components were considered relative to Mg# as a

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fractionation index. The rocks show a clear positive correlation of Mg# with Ni, Cr, and Sr, and a

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negative correlation with Zr, Nb, Y, Th, TiO2, and P2O5 (Fig. 4). The positive correlation of Ni

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with Mg# indicates the fractionation of olivine. The positive correlation of Cr and Sr with Mg#

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suggests the fractionation of clinopyroxene (Fig. 4). Thus, the negative correlation of Mg# with

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TiO2 and P2O5 along with the non-systematic variations of Fe2O3 and V suggest that the

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fractionation of Fe–Ti oxides and apatite was not important in the studied rock series. An

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increase of incompatible elements, such as Zr and Th, with increasing degree of fractionation can

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be caused by crustal contamination. Some increase in the Y content with the decrease of Mg#

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argues against the amphibole fractionation. The most significant variations are observed for Cr

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(45 times, 25–1123 ppm) and Ni (ca. 20 times, 20–376 ppm).

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The chondrite-normalized REE patterns show LREE enrichment with very insignificant

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or no Eu anomaly. The primitive mantle (PM)-normalized multielements diagrams show variable

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enrichment in LILE, negative Nb and Sr anomalies, and no or positive Ti anomaly (Fig. 5). All

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rocks define near parallel patterns, which suggest the comagmatic origin of the rock series

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through fractionation. High Gd/Yb ratio (2.04–2.77) in the high-Mg rocks points to a garnet-

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bearing source (e.g., Bryant et al., 2006). The rocks are characterized by the slightly negative εNd (2445), which varies from –0.7 to

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–2.9, increasing in the high-Mg# rocks (Table 2). These values, in general, are close to εNd data

283

on Early Paleoproterozoic SHMS rocks around the world (Puchtel et al., 1997, 2016; Vrevsky,

284

2011; Ketchum et al., 2013; Bogina et al., 2015).

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6. Discussion

As mentioned above, Early Paleoproterozoic SHMS rocks have been reported from

288

different cratons at the Archean–Paleoproterozoic boundary. However, data on volcanic rocks of

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this age are rarer due to their lesser abundance as compared to dikes and intrusions of the same

290

age. Representative data on these rocks from different domains of the Karelian Craton, Cola

291

Craton, Canadian Shield, and Indian Craton are compiled and shown in (Figs. 6 and 7). The

292

comparison of these volcanic rocks revealed a general similarity in their major, trace, and Nd

293

isotopic characteristics. Most of the rocks correspond to basalt and basaltic andesite. Most

294

primitive high-Mg rocks (Mg# up to 67) of this series occur in the Vodlozero Domain. The

295

studied rocks are characterized by enrichment in SiO2 at a given MgO content, but low TiO2 and

296

high Cr contents. On this basis, they are frequently termed as boninite-like rocks (Sharkov et al.,

297

2005). However, the TiO2 content is higher relative to boninites, while the SiO2 is lower. The

298

TiO2 is lower than 1.6 (generally <1.2). Note that the lowest TiO2 content (< 1%) is noted in the

299

high-Mg rocks. Its content possibly increases during the crystallization differentiation.

300

Correspondingly, spidergrams show from weakly negative to weakly positive Ti anomalies in the

301

more fractionated varieties. All of the rocks are characterized by the REE enrichment, which is

302

also typical feature of the SHMS rocks, and practically parallel trace-element patterns (Fig. 7).

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Exceptions are the rocks from the lower part of the Seidorechka Formation (Kola

304

Craton), as already mentioned in Hanski et al. (2012). They are marked by the following specific

305

features: lower SiO2 content; higher FeO, Al2O3, and Cr2O3 contents; weakly fractionated REE

306

patterns as compared to other rocks from this locality (Figs. 6 and 7), and positive εNd (+1.0 and

307

+1.4, Hanski, 2012). According to Fedotov et al. (2009), these rocks can genetically be related to

308

scarce low-Ti tholeiitic dikes found in this area.

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All SHMS rocks define a negative Nb anomaly of variable depth (Nb/Nb* usually below

310

0.5). Their εNd values are weakly negative, ranging mainly between 0 and –3. In particular, the

311

εNd in rocks of the Vodlozero Domain (this study) varies from –0.7 to –2.9, which is close to that

312

in basalts of the Vetreny Belt (from –0.82 to –2.0) (Puchtel et al., 2016). The εNd value in the

313

basaltic rocks from the Central Domain is between –0.4 and –2.3 (Bogina et al., 2015). The Nd

314

isotope composition in the mafic rocks of the Seidorechka Formation varies from –0.9 to –2.8

315

(Chashchin et al., 2008; Hanski, 2012). The εNd value in the Mantyvaara area, central Finnish

316

Lapland (Hanski and Huhma, 2005) also varies similarly from –0.2 to –3.2. Basaltic andesites

317

from the Thessalon Formation (Jolly et al., 1992) have εNd from –0.5 to –2.2. Samples with

318

extremely low εNd values (–4.5 and –4.9) falling beyond this range were found in basalts, which

319

are developed at the base of this rock sequence and interlayered with sedimentary rocks, can be

320

related with intense in situ assimilation of host rocks.

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To sum up, all Early Paleoproterozoic SHMS rocks are characterized by close

322

petrochemical composition, similar trace element patterns (Fig. 7) with the LREE enrichment,

323

negative Nb anomalies, and weakly negative Nd isotope compositions varying within a relatively

324

narrow range. This implies that they were generated by similar processes. The LREE enrichment

325

and negative εNd values are indicative of either crustal contamination or derivation from the

326

metasomatized (reworked by previous subduction events) subcontinental lithospheric mantle

327

(SCLM). At the same time, Paleopoterozoic SHMS rocks are localized in domains and shields

328

with different crustal evolution histories. They were intruded into heterogeneous basements,

329

which would seemingly result in different compositions. Two important problems are discussed

330

here based on the obtained and available literature data: (1) what factors were responsible for

331

such a remarkable global similarity of the Early Paleoproterozoic mafic volcanic (SHMS) rocks

332

in the heterogeneous basement of different cratons; (2) which mechanism produced variations in

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the studied rock series of the Vodlozero Domain from the high-Cr high-Mg to the low-Cr low-

334

Mg basalts.

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The geochemical and isotope geochemical similarities suggest a common formation

336

mechanism. Therefore, it is hardly possible that such a global similarity would be provided by

337

their derivation from the variably metasomatized SCLM. In addition, based on the recently

338

obtained Os isotope data, Yang et al. (2016) concluded that the SCLM cannot be the mantle

339

source for Early Paleoproterozoic intrusions, and correspondingly, for genetically related and

340

compositionally similar volcanic rocks. It is also unlikely that the contribution of crustal rocks of

341

different compositions and ages would produce the observed narrow compositional ranges.

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342

Based on new data obtained in this study, we propose a universal mechanism, which

343

explains both the ubiquitous development of the Paleoproterozoic SHMS around the world and

344

their geochemical similarities. The most primitive least contaminated samples (E-7, E-8) with maximum Mg# (ca. 67)

346

and the highest contents of compatible elements (Cr, Ni) presumably are closer to primary

347

magmas in equilibrium with the mantle peridotite (Frey et al., 1978). Correspondingly, these

348

samples can be used to estimate a mantle source. Their HREE fractionation marked by the high

349

Gd/Yb is indicative of residual garnet in the source during the partial melting and, hence, their

350

generation from a garnet-bearing source, i.e., under a pressure of more than 15 kbar. The low Sc,

351

Y, and Lu/Hf, and the fractionated HREE pattern in our samples are also consistent with the

352

presence of garnet in the source. At the same time, their negative εNd suggests the contribution of

353

a crustal component. To test this hypothesis, we carried out model calculations using a simple

354

mixing model. An important point was a choice of end members. With allowance for the LREE

355

enrichment and the trace element evidence for the presence of garnet in their source, we assumed

356

that the mafic end member was a partial melt from a depleted mantle. Using the formula of

357

modal batch melting, the primary melt was calculated as 30% partial melt from depleted mantle

358

(Salters and Stracke, 2004) in equilibrium with garnet peridotite restite (53% Ol + 34% Cpx +

359

8% Opx + 5% Grt). Mineral modes and partition coefficients were also taken from (Salters and

360

Stracke, 2004). The resultant melt contained 2.3 ppm of Nd. The εNd value of 4.7 for this melt

361

was calculated at 2.5 Ga from the DM evolution curve (Goldstein et al., 1984). The 3.2 Ga

362

granite gneisses, 2.8 Ga TTG, and 2.7 Ga sanukitoids of the Vodlozero Domain were taken as

363

the possible crustal end members. Specific samples taken for the modeling are mentioned in

364

notes to Fig. 8. As calculated for 2.45 Ga, the Archean sanukitoids are usually characterized by

365

sufficiently narrow and weakly negative εNd at high Nd content; Neoarchean (2.74) adakites at

366

close εNd have lower LREE contents. Mesoarchean (3.2 Ga) TTG gneisses of the Vodlozero

367

Domain are similar to adakites in terms of the LREE content, but have extremely low εNd values

368

(Lobach-Zhuchenko et al., 1993). It is seen in the diagram that hybrid melts produced by the

369

mixing of DM with adakites and TTG gneisses fall away from the data points of Early

370

Paleoproterozoic rocks, whereas the mixing of partial melt of depleted mantle with sanukitoids

371

could provide the sufficiently uniform Nd isotopic composition of the Early Paleoproterozoic

372

mafic rocks. Moreover, the model of simple mixing between the most primitive sanukitoids and

373

the depleted mantle demonstrates that, regardless of the mixing percentage, the Nd isotope

374

composition remains practically at the same level (weakly negative values), thus providing

375

relative isotope-geochemical homogeneity of the studied Early Paleoproterozoic rocks (Fig. 8).

376

As known, sanukitoids were identified at all cratons: Superior Province of the Canadian Shield

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(e.g., Stern and Hanson, 1991), Fennoscandian Shield (e.g., Larionova et al., 2007; Heilimo,

378

2010), South Indian Dharwar Craton (Sarvothaman, 2001), Western Greenland (Steenfelt et al.,

379

2005), and others. In the Western Superior Province, for instance, Archean sanukitoids compose

380

up to 15% of exposed rocks (Stevenson et al., 1999). On the basis of their geochemical

381

signatures, Stern et al. (1989) and Stern and Hanson (1991) argued that this rock series was

382

formed by the melting of an enriched mantle wedge above a subsiding slab. As a result, these

383

rocks are characterized by very intense LREE enrichment, which is usually attributed to

384

reworking of the sanukitoid source by the LILE-rich fluids derived by slab dehydration during

385

previous subduction events. In our opinion, by the beginning of the Early Paleoproterozoic,

386

rocks of this series presumably composed the lower crust beneath most Archean cratons,

387

whereas the earlier TTG granitoids and supracrustal rocks were localized at higher levels. Based

388

on the ubiquitous presence of sanukitoids in the world, we suggest that their most primitive

389

varieties localized in the lower crust served as a universal contaminant for the Early

390

Paleoproterozoic plume-related mantle melts. Lauri et al. (2012) considered the Archean

391

sanukitoids as a possible source for the A-type granites, which are frequently associated with

392

Early Paleoproterozoic SHMS rocks, forming a bimodal series. These authors proposed that

393

sanukitoids served as a potential fertilizer of the lower crust at that time. The large-scale

394

emplacement of mafic melts in the lower crust in the Paleoproterozoic could cause the melting of

395

lower-crustal sanukitoids with the formation of A-type granites and the simultaneous

396

contamination of mantle magmas. The indirect evidence for the presence of sanukitoids in the

397

lower crust comes from finds of the lower-crustal xenoliths with the protolith age of 2.75–2.65

398

Ga in the carbonatite dike near Kandalaksha on the Kola Peninsula (Claesson et al., 2000).

399

Analysis of the lower crustal garnet-bearing granulite xenoliths from Elovy Island showed that

400

some of them are enriched in Ba, Sr, LREE, and Cr (Kempton et al., 1995), which are typical

401

signatures of sanukitoids. Sanukitoids are also widespread in the high-grade western part of the

402

Marmion terrane, Superior Province, Canadian Shield (Tomlinson et al., 2004). The high-REE

403

Cpx enderbites found among the lower crustal Nilgiri Hills granulites and formed at a pressure of

404

9–10 kbar were recognized in the Dharwar Craton, India. They are interpreted as the former

405

igneous Mg-K suite of sanukitoid affinity (Peucat et al., 2013). All these data support the idea

406

that sanukitoids were present the lower crust in the Early Paleoproterozoic and could serve as a

407

universal contaminant during the formation of the SHMS rocks.

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Geochemical data indicate the accumulation of incompatible elements in late derivatives,

409

and some increase of REE contents from the high-Cr high-Mg basalts to basaltic andesites and

410

andesites, which is generally consistent with the model of fractional crystallization. As

411

mentioned above, the positive correlation of Ni and Cr with Mg# indicates fractionation of

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olivine and clinopyroxene, respectively (Wang et al., 2013). Analysis of isotope data revealed a

413

positive correlation of εNd with Cr and Mg#, the negative correlation with Th, and slight decrease

414

of εNd in the more evolved varieties. All these facts indicate that the rock series was formed by

415

the coupled assimilation and fractional crystallization. The fractionating assemblage was

416

estimated by the major and Fe group element modeling using a Comagmat-3.52 Software

417

(Ariskin et al., 1993). It was established that the low-Cr rocks can be obtained from high-Cr rock

418

by the subsequent fractionation of olivine, clinopyroxene, and plagioclase. Then, we performed

419

the assimilation-fractional crystallization modeling (AFC) for Cr and V using the Excel

420

spreadsheet (Ersoy and Helvaci, 2010). The average lower crust and differentiated rocks of the

421

sanukitoid series of the Bergaul batholith, Vodlozero Domain (middle–upper crustal levels)

422

(Larionova et al., 2007) were used as contaminants. The high-Mg basalt with the highest Cr and

423

Ni contents from the Elmus structure was taken as the starting composition. The composition of

424

the studied volcanic rock series is well consistent with the AFC model (Fig. 9); the application of

425

the average lower crust (Taylor and McLennan, 1985) as contaminant was unsuccessful. Hereby,

426

the rocks experienced contamination at the middle–upper crustal levels by the felsic (more

427

fractionated) derivatives of the sanukitoid series of the Bergaul batholith in the Vodlozero

428

Domain. This is supported by some increase of the Y content and decrease of Gd/Yb ratios in the

429

more evolved rocks. At the same time, the isotope shift was insignificant, indicating that the rock

430

composition was mainly controlled by fractionation. The trace-element modeling showed that

431

the rock series may be derived by the AFC mechanism (60% Cpx + 20% Ol + 20% Pl; r = 0.02).

432

The 2.7-Ga fractionated high-Si member of the sanukitoid series emplaced at the upper crustal

433

level of the Vodlozero Domain was used as contaminant (Fig. 7; Larionova et al., 2007).

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Thus, the obtained data and modeling suggest that the parental melt for the studied

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volcanic rocks was derived by the partial melting of a depleted mantle. This is consistent with

436

previous works suggesting that the Early Paleoproterozoic magmatism was related to the melting

437

of DM plume source without a significant input of the recycled material (Yang et al., 2016 and

438

references therein). The presence of a DMM (Depleted MORB Mantle) component in the

439

magma source agrees well with scarce finds of low-Ti tholeiitic dikes in the Karelian and Kola

440

cratons (Fedotov et al., 2005; Vuollo and Huhma, 2005) and volcanic rocks in the lower part of

441

the Seidorechka Formation with weakly or no fractionated REE patterns and slightly positive εNd

442

values. Owing to the ubiquitous development of the sanukitoid series in all cratons, the proposed

443

model seems to be universal and can be applied to explain the formation of the Early

444

Paleoproterozoic SHMS around the world.

445

7. Conclusions

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(1) Early Paleoproterozoic mafic volcanic rocks of the Vodlozero Domain widely vary in

447

composition from the high-Mg# (>60) basalts with high Cr contents (up to 1200 ppm) to

448

moderate and low-Mg basalts with the moderate and low Cr contents. They are characterized by

449

LREE-enriched patterns with (La/Yb)n up to 12.34 and weakly negative εNd (from –0.7 to –2.9).

450

(2) Chondrite normalized REE pattern in the high-Mg rocks of the Vodlozero Domain suggests

451

their generation from a garnet-bearing source. At the same time, their negative εNd in

452

combination with the LREE enrichment points to their crustal contamination. Their generation is

453

consistent with contamination of the high-degree (30%) partial melt derived from a depleted

454

mantle. Primitive rocks of the sanukitoid series could serve as a universal contaminant.

455

(3) All rock series from the high- to low-Cr rocks were generated by the fractionation of high-Cr

456

rocks coupled with an insignificant crustal assimilation of felsic end members of the sanukitoid

457

suite of the Vodlozero Domain. This statement is supported by the positive correlation of εNd

458

with Cr and Mg#, negative correlation with Th, and slight decrease of εNd in the more evolved

459

varieties. The main fractionating minerals were olivine, clinopyroxene, and plagioclase.

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Acknowledgments

462

We greatly thank two anonymous reviewers whose very constructive comments essentially

463

improved our manuscript. This study was supported by the Russian Foundation for Basic

464

Research (Project No. 16-05-00708 and 15-05-01214).

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the

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Fennoscandian

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subalkaline

granitoids)

in

the

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Craton:

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Zlobin, V.L., Konilov, A., Mints, M.V., Kaulina, T.V., Bulatov, B.S., 2003. Paleoproterozoic

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ignimbrites of the Lekhta ignimbrites, Central Karelia, Proceedings of 2nd All-Russia

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volcanics: Doklady Earth Sciences 435, 1415–1419.

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694 695

FIGURE CAPTIONS. Fig. 1. Geological map of the Karelian Craton with indication of the occurrences of the

696

Sumian and Sariolian volcanic rocks (modified after Bogina et al., 2015). Early Paleoproterozoic

698

supracrustal structures: (1) Krasnaya Rechka, Semch, Koikary, (2) Kumsa, (3) Lekhta, (4)

699

Shomba, (5) Vetreny Belt, (6) Paanajarvi, (7) Imandra–Varzuga, (8) Arvarench, (9) Pechenga;

700

Lapland greenstone belt: (10) Kittila, (11) Salla, (12) Kuusamo (Koilismaa Complex), (13)

701

Perapohja; Layered massifs: (14) Burakovsky Massif, (15) Oulanka Complex, (16) Monchegorsk

702

Complex, (17) Fedorova-Pana Massif, (18) Mt. General’skaya Massif, (19) Portimo, (20) Kemi,

703

(21) Elovy xenoliths locality. Domains of the Karelian Craton: VD–Vodlozero, CD–Central

704

Karelian, WD–Western Karelian.

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Fig. 2. Morphology of the Sumian rock lithotypes: (a) amygdaloidal basaltic andesite

705

lavas with polygonal jointing (Koikary structure); (b) lava breccias (Semch structure); (c) “giant”

707

amygdules in a lava flow (Kumsa structure); (d) quartz–albite amygdules in the roof breccias of

708

massive basaltic andesite lava flow (Kumsa structure), (e, f) pillow lavas of basaltic andesite,

709

interpillow space is filled with tuff material with admixture of sandstone. Matrix is subjected to

710

late carbonatization (Elmus structure).

711 712

Fig. 3. The Zr/TiO2 vs. Nb/Y classification diagram (after Winchester and Floyd, 1977) for Early Paleoproterozoic volcanic rocks of the Vodlozero domain.

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Fig. 4. Variations of abundances of selected major (in wt.%) and trace elements (in ppm)

713 714

in whole rock samples of Early Paleoproterozoic volcanic rocks of the Vodlozero Domain vs.

715

Mg#.

Fig. 5. Chondrite- and PM-normalized REE and trace element abundances in the whole

717

rocks samples of Early Paleoproterozoic volcanic rocks of the Vodlozero Domain. The low-Cr

718

rocks of the Elmus structure are shown by empty crosses. Normalizing values are from Evensen

719

et al.(1978) and Hofmann (1988).

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Fig. 6. Selected major and trace element compositions of SHMS rocks from the

721

Vodlozero Domain (this study) as compared with the SHMS rocks from the Shomba and Lekhta

722

structure, Central Domain, Karelian Craton (Bogina et al., 2015), from the Vetreny Belt

723

Formation, Karelian Craton (Puchtel et al., 1997), from the Seidorechka Formation, Imandra–

724

Varzuga structure, Kola Craton (Hanski et al., 2012; Melezhik et al., 2013), from the Thessalon

725

Formation, Canadian Shield (Ketchum et al., 2013), and Dongargarh Supergroup, India

726

(Sensarma et al., 2002).

727 728

Fig. 7. Primitive mantle-normalized abundances of selected lithophile trace elements for SHMS rocks from the Vodlozero Domain, Karelian Craton (this study), Shomba and Lekhta

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structures, Central Domain, Karelian Craton (Bogina et al., 2015), Seidorechka Formation,

730

Karelian Craton (Hanski et al., 2012; Melezhik et al., 2013), from the Thessalon Formation,

731

Canadian Shield (Ketchum et al., 2013), and Dongargarh Supergroup, India (Sensarma et al.,

732

2002).

733

Fig. 8. Simple mixing modeling showing the formation of Nd isotope composition of Early Paleoproterozoic volcanic rocks of the Vodlozero Domain. As end-members, we took 30

735

% partial melt of depleted mantle after (Salters and Stracke, 2004) in equilibrium with Grt

736

peridotite restite (formed at approximately 3 GPa) and primitive high-Mg sanukitoid from the

737

Bergaul Massif in the Vodlozero Domain (sample K-17-2, Larionova et al., 2007). 2.85-Ga

738

adakites (Sample 8#94, Puchtel et al., 1999) and 3.2-Ga TTG granites from the Vodlozero

739

Domain (Sample K-13/85, Lobach-Zhuchenko et al., 1993) were also used as end members for

740

comparison. Mineral modes and partition coefficients used for calculation of the partial melt

741

were taken from (Salters and Stracke, 2004). The values of Nd and εNd (2.45) in end members

742

were calculated from the DM partial melting.

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in Early Paleoproterozoic volcanic series of the Vodlozero Domain.

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Fig. 9. Model calculations showing the effect of AFC on the Cr and Ni content (in ppm)

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Table 1. Representative analyses of Sumian volcanic rocks from the Vodlozero block (oxides in wt %, elements in ppm)

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1

2

3

4

5

6

7

8

E-1/13

E-3/13

E-4/13

E-7/13

E-8/13

E-09

E-14

KOJ-01

49.71 0.55 10.31 12.08

49.78 0.75 17.55 11.11

58.07 0.55 11.51 8.02

50.6 0.61 12.92 10.80

45.14 0.64 11.24 12.57

48.02 0.94 11.04 14.58

54.43 0.71 14.26 10.17

50.43 0.75 14.87 10.48

54.89 1.15 13.44 12.26

0.21 12.14 11.36 0.78 0.22 0.07

0.19 6.37 6.94 3.77 0.38 0.1

0.14 8.64 6.54 3.83 0.50 0.07

0.17 9.98 6.84 3.92 0.14 0.08

0.19 13.28 13.48 0.17 0.09 0.08

0.21 11.52 7.81 2.48 0.24 0.12

0.15 6.11 6.93 4.26 0.6 0.11

0.21 9.51 4.93 4.68 0.18 0.09

2.30 99.73 67 30 192 784 68 258 4 108 12 69 2 71 9.41 18.97 2.39 10.15 2.09 0.57 2.26 0.33 2.03 0.42 1.30 0.17 1.01 0.12 0.76 0.10 2.66 1.20 0.10

2.88 99.82 53 36 262 170 48 73 7 237 27 120 4 117 40.30 71.26 8.38 32.77 5.70 1.97 6.66 0.83 4.69 0.92 2.85 0.35 2.20 0.26 1.38 0.20 4.50 3.30 0.70

1.86 99.73 68 31 159 773 45 215 12 97 8 71 2 256 7.01 14.59 1.91 7.83 1.63 0.31 1.82 0.25 1.58 0.30 0.96 0.12 0.81 0.11 0.94 0.06 2.59 1.00 0.00

3.65 99.71 65 32 182 831 63 265 2 108 11 78 3 35 9.18 20.48 2.61 11.13 2.31 0.46 2.40 0.34 2.04 0.40 1.22 0.16 0.97 0.12 0.97 0.12 4.55 1.40 0.12

2.81 99.69 68 33 213 928 66 268 1 253 15 80 3 29 14.52 31.16 3.75 15.38 3.10 0.95 3.44 0.48 2.82 0.53 1.53 0.19 1.19 0.14 0.92 0.12 3.51 1.49 0.14

2.74 99.7 61 32 207 908 72 340 4 51 14 58 6 68 10.91 27.27 3.76 16.05 3.74 0.80 3.36 0.50 2.88 0.58 1.53 0.22 1.29 0.17 1.81 0.39 2.05 2.94 0.66

2.11 99.84 54 27 176 181 45 91 17 142 12 81 3 221 13.17 28.80 3.55 13.97 2.79 0.66 2.50 0.43 2.40 0.50 1.43 0.21 1.32 0.17 1.44 0.23 6.29 3.17 0.59

3.58 99.71 64 42 232 573 27 244 3 168 14 55 3 93 7.56 15.64 2.14 9.24 2.27 0.45 2.38 0.42 2.69 0.60 1.72 0.24 1.57 0.23 1.70 0.16 22.88 0.65 0.13

9

10

53.46 0.64 14.78 11.22

11 K-014a/15 47.27 0.96 11.00 14.72

12 Gir-061/11 51.67 1.26 11.89 13.20

13 Gir07/11 52.71 0.98 9.82 11.89

0.19 4.13 4.44 4.5 1.84 0.15

0.18 5.54 6.85 4.05 0.41 0.11

0.20 10.17 9.12 0.49 1.86 0.09

0.16 7.54 7.76 3.33 0.42 0.14

0.18 9.00 8.52 2.96 0.41 0.08

2.79 99.78 40 17 186 31 49 95 51 141 15 103 8 444 21.86 48.15 6.07 23.51 4.83 1.33 3.99 0.63 3.35 0.65 1.65 0.23 1.45 0.19 3.59 0.50 5.00 3.68 0.84

2.01 99.27 49 29 171 114 45 109 14 124 11 125 2 180 12.73 30.15 3.75 14.65 2.92 0.81 2.84 0.43 2.19 0.46 1.34 0.18 1.20 0.19 1.95 0.29 44.46 3.92 0.54

3.67 99.67 58 32 198 775 70 364 30 132 13 106 4 603 12.69 31.95 4.17 17.34 3.81 1.17 3.98 0.54 2.68 0.55 1.36 0.18 1.17 0.16 1.92 0.38 22.29 2.54 0.43

2.40 99.77 52 26 169 318 47 177 11 133 16 135 6 93 14.87 36.27 4.43 19.15 4.27 1.29 3.99 0.58 3.12 0.58 1.58 0.20 1.31 0.20 3.05 0.43 4.28 2.75 0.56

3.13 99.68 59 27 165 524 55 252 12 63 12 104 5 125 20.48 41.94 4.69 19.05 3.48 1.25 3.11 0.43 2.35 0.46 1.26 0.17 1.09 0.16 2.25 0.32 2.41 2.10 0.46

SC

RI PT

KOJ-09 Se-12/15

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EP

AC C

Sample No. SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 H2O L.O.I. Total # Mg Sc V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

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52.34 1.45 13.47 15.01

0.16 3.77 5.45 3.98 1.82 0.16

0.20 5.03 5.69 3.53 0.83 0.16

0.19 10.19 9.4 0.62 2.00 0.09

0.16 7.35 8.23 3.37 0.56 0.14

0.15 4.32 7.94 3.8 0.52 0.15

0.15 3.9 8.03 4.04 0.53 0.15

0.15 5.4 10.15 3.82 0.36 0.12

0.19 3.57 7.18 3.79 0.42 0.17

0.88 99.79 36 19 161 10 39 59 63 247 14 165 7 342 17.96 40.11 4.86 19.71 3.90 0.90 3.50 0.49 2.75 0.53 1.46 0.19 1.26 0.18 3.37 0.49 4.27 2.91 0.59

2.07 99.78 39 21 169 27 45 74 12 103 15 166 7 373 22.75 54.90 6.31 25.29 4.83 0.91 3.96 0.57 2.95 0.55 1.53 0.21 1.33 0.19 3.55 0.51 2.79 3.19 0.63

1.58 99.61 61 29 165 576 66 264 44 238 13 99 5 819 12.22 29.08 3.64 15.39 3.07 0.81 3.03 0.43 2.37 0.45 1.23 0.17 1.09 0.16 2.07 0.38 3.63 2.14 0.47

2.06 99.74 53 20 157 327 53 176 16 227 16 128 6 176 13.25 30.75 3.81 16.19 3.40 0.96 3.20 0.48 2.66 0.52 1.43 0.19 1.21 0.18 2.69 0.48 2.92 2.85 0.56

2.24 99.76 41 17 168 22 43 62 11 181 17 143 7 153 23.12 54.37 6.20 25.52 4.70 1.58 4.41 0.59 3.10 0.60 1.56 0.20 1.32 0.20 3.10 0.57 6.68 3.40 0.68

0.61 99.74 42 17 162 16 35 61 9 349 16 154 7 169 17.44 41.32 5.01 21.13 4.02 1.34 3.80 0.51 2.90 0.56 1.52 0.21 1.30 0.19 3.24 0.58 6.34 3.50 0.69

0.78 99.58 49 25 178 369 43 189 6 1114 16 138 7 143 17.74 39.28 4.46 19.17 3.87 1.37 3.41 0.50 2.90 0.56 1.54 0.19 1.25 0.17 2.60 0.43 6.57 2.39 0.59

1.19 99.75 34 18 194 14 46 58 8 485 17 172 10 102 19.22 46.12 5.49 23.62 4.74 1.28 4.07 0.59 3.22 0.62 1.76 0.23 1.49 0.21 3.67 0.77 4.70 2.98 0.76

AC C

19 Med15/11 56.73 1.14 13.90 10.56

20

21

22

23

24

25

Kum-40

Kum-48

904

1013-1

1008-2

1007-4

54.12 1.06 12.77 10.85

54.32 1.43 14.12 13.37

54.30 1.55 16.72 2.72 9.34 0.15 2.22 2.38 6.10 2.82

53.34 1.12 13.96 1.65 9.94 0.18 3.85 4.97 4.56 1.75

52.60 1.12 14.77 2.09 10.54 0.25 4.60 5.48 4.03 1.80

54.92 1.07 14.14 2.16 10.3 0.14 4.40 5.26 2.87 2.65

0.12 1.74 100.16 31

0.10 2.17 97.59 45

0.14 2.33 99.75 47

0.14 1.74 99.79 46

8 56 25 192 7 47 9.24 25.34 3.34 14.50 3.00 1.10 3.90 0.72 4.52 0.88 2.84 0.38 2.68 0.36 5.32 0.56 2.60 5.96 1.32

49 281 19 101 7 449 22.30 53.00 6.70 28.04 5.44 1.44 5.50 0.80 4.02 0.72 2.06 0.22 1.50 0.18 3.00 0.58 1.90 3.28 0.80

49 175 22 79 7 727 29.70 69.84 8.82 35.72 6.92 1.76 6.72 0.98 4.78 0.88 2.28 0.26 1.60 0.18 2.12 0.52 2.24 3.76 0.88

69 318 16 71 6 1064 23.68 50.74 6.64 26.02 5.20 1.44 4.94 0.74 3.52 0.68 1.86 0.20 1.32 0.14 1.78 0.52 2.80 3.52 0.72

SC

18 Med14/11 54.88 1.20 12.36 12.2

M AN U

Gir-11/11 Gir-15/11

17 Med09/11 52.88 1.14 11.13 12.72

RI PT

55.74 1.39 13.2 13.24

16 Med08/11 52.59 0.97 9.08 12.9

TE D

15

EP

14 Sample No. SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O P2O5 H2O L.O.I. Total # Mg Sc V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U

Note: Samples 1-7 are from the Elmus structure; samples 8-11 from the Koikary structure; samples 12-15 from the Krasnaya Rechka structure; samples 16-25 from the Kumsa structure.

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Table 2. Sm-Nd isotope data on the early Paleoproterozoic volcanics of the Vodlozero block, Karelian Craton

1013-1 1008-2 1007-4

Sm

Metabasalt Metabasalt Metabasalt Metabasalt Metabasalt Metaandesite Metabasalt Metabasalt Basaltic andesite basaltic andesite basaltic andesite basaltic andesite

Nd

Isotope ratios 147

144

Sm/ Nd

143

Nd/144Nd

εNd(2.45 Ga) -2.9 -0.9 -1.3 -1.8 -2.0 -1.2 -0.7 -1.8 -2.0

6.519 1.698 1.93 4.95 6.624 3.968 3.491 4.524 5.448

35.62 8.103 9.13 24.32 34.328 21.245 17.043 22.893 28.04

0.11062 0.12668 0.12781 0.12304 0.116640 0.112902 0.123800 0.119461 0.1174

0.511098 (±5) 0.511460 (±18) 0,5114578 (±6) 0.5511357 (±10) 0.511242 (±6) 0.511225 (±6) 0.511423 (±12) 0.511296 (±14) 0.511252

5.129

24.63

0.1259

0.511395

-1.9

6.522

33.44

0.1179

0.511299

-1.3

4.296

21.98

0.1182

-1.9

RI PT

E-3/13 E-4/13 E-7/13 Koi-09/13 Gir 15/11 Gir 07/11 Med 08/11 Med 15/11 904

Content (ppm) Rock

SC

Sample No.

0.51127

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Samples E-3/13, E-4/13, and E-7/13 are from the Elmus structure; sample Koi-09/13 from the Koikary structure; samples Gir 15/11 and Gir 07/11 from the Krasnaya Rechka structure; samples Med 08/11, Med15/11, 904, 1013-1, 1008-2, and 1007-4 are from the Kumsa structure. Samples 904, 1013-1, 1008-2 and 1007-4 were analyzed at the Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences.

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~2.45 Ga mafic volcanics of the SHMS series of the Vodlozero block are studied. Their derivation via mixing of depleted mantle and adakitic lower crust is proposed.

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The entire rock series is modeled by AFC with Ol , Cpx, and Pl as liquidus phases