Shoshonitic magmatism in the Paleoproterozoic of the south-western Siberian Craton: An analogue of the modern post-collision setting

Lithos 328–329 (2019) 88–100

Contents lists available at ScienceDirect

Lithos journal homepage: www.elsevier.com/locate/lithos

Shoshonitic magmatism in the Paleoproterozoic of the south-western Siberian Craton: An analogue of the modern post-collision setting Alexei V. Ivanov a,⁎, Ivan V. Levitskii b, Valery I. Levitskii b, Fernando Corfu c, Elena I. Demonterova a, Leonid Z. Reznitskii a, Ludmila A. Pavlova b, Vadim S. Kamenetsky d, Valery M. Savatenkov e,f, Vladislav I. Powerman a,g a

Institute of the Earth's Crust, Siberian Branch of the Russian Academy of Sciences, Irkutsk, Russia A.P. Vinogradov Institute of Geochemistry, Siberian Branch of the Russian Academy of Sciences, Irkutsk, Russia c University of Oslo, Department of Geosciences and CEED, Oslo, Norway d University of Tasmania, School of Natural Sciences, Hobart, Tasmania, Australia e Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences, Saint-Petersburg, Russia f Institute of the Earth's Sciences, Saint-Petersburg State University, Saint-Petersburg, Russia g Schmidt Institute of Physics of the Earth RAS, Moscow, Russia b

a r t i c l e

i n f o

Article history: Received 22 March 2018 Accepted 12 January 2019 Available online 22 January 2019 Keywords: Siberian Craton Paleoproterozoic Shoshonitic series U-Pb dating Geochemistry Sr-Nd-Pb isotopes

a b s t r a c t The Siberian Craton was assembled in a Paleoproterozoic episode at about 1.88 Ga by the collision of older blocks, followed at about 1.86 Ga by post-collisional felsic magmatism. We have found a set of extremely fresh micabearing lamprophyre-looking rocks within the Sharyzhalgay metamorphic complex of the south-western Siberian Craton. Zircon from these rocks yields a U\\Pb TIMS age of 1864.7 ± 1.8 Ma, which coincides perfectly with the peak of the post-collisional granite ages and postdates by ~15 Ma the peak of ages obtained for metamorphism. The same ages were reported earlier for a mafic dyke with ocean island basalt (OIB) geochemical signatures and a Pt-bearing mafic-ultramafic intrusion found in the same region. Mineralogy, major and trace element geochemistry and Sr-Nd-Pb isotopes show that the studied rocks (1) have shoshonitic affinity, (2) are hybrid rocks with mineral assemblages which could not be in equilibrium, (3) where derived by recycling of an Archean crustal source and (4) resemble post-collision Tibetan shoshonitic series. The genesis of these rocks is considered to be due to melting of crustal lithologies and metasomatized lithospheric mantle within a subducted slab. Some of the resulting melts ascended through the lithospheric column and fractionated to low-Mg absarokites, whereas other melts were contaminated by orthopyroxenitic mantle material and attained unusual high-Mg mafic compositions. According to our model, the post-collisional magmatism (shoshonite- and OIB-type) occurred due to upwelling of hot asthenosphere through a slab window, when the active collision ceased as a result of the slab break off and loss of the slab pull force. Overall, our study shows that in the Paleoproterozoic shoshonitic melts were emplaced within a similar tectonic setting as seen today in modern orogenic systems. © 2019 Elsevier B.V. All rights reserved.

1. Introduction ‘The present is the key to the past’. Uniformitarianism, whose basic principles were formulated by James Hutton more than two centuries ago, is one of the pillars of geological sciences. However, were the processes in the earliest Precambrian the same as in the modern Earth today? For example, it is debated when the modern tectonic style started – already in the Hadean or the Eoarchean (de Wit, 1998; Turner et al., 2014), at ~2 Ga in the Paleoproterozoic (Hamilton, 2003), much later at about 1 Ga in the Neoproterozoic (Davies, 1992; Stern, ⁎ Corresponding author. E-mail address: [email protected] (A.V. Ivanov).

https://doi.org/10.1016/j.lithos.2019.01.015 0024-4937/© 2019 Elsevier B.V. All rights reserved.

2005), or was the process of transformation from earlier to modern tectonic styles more gradual (Stern, 2007)? Igneous rocks are commonly suggested as proxies of tectonic settings (Kay and Kay, 1993; Pearce, 1976, 1982, 2008; Pearce et al., 1984; Pearce and Norry, 1979; Verma and Agrawal, 2011; Verma et al., 2006; Vermeesch, 2006) by assuming that rock chemical compositions adequately reflect the processes by which magma generated – passive decompression within oceanic and continental rifts, active mantle upwellings under intraplate regions, volatile flux addition above subducting slabs, crustal recycling due to lithospheric delamination in orogens etc. One of the limiting factors for using igneous rocks for tectonic identification is the rock alteration (Sheth, 2008), which is often the case for the rocks of the distant past. In this paper, we report the discovery of Paleoproterozoic, yet unusually

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

fresh mica-bearing igneous rocks with a completely preserved original mineral assemblage. These rocks occur within the boundaries of the Archean and Paleoproterozic Sharyzhalgay metamorphic complex of the south-western Siberian Craton (Fig. 1). Precise U\\Pb dating has revealed an age of ~1865 Ma, which is younger than the age of the preceding metamorphism. Geochemical and isotopic data show that the rocks have much in common with shoshonites erupted in the Late Cenozoic in Tibet during the post-collisional tectonic stage. Therefore, we interpret these data as indicating that tectonics at ~1865 Ma were akin to the modern tectonic style and the south-western Siberian Craton is a Paleoproterozoic analogue of modern orogenic systems. 2. Geology of the Sharyzhalgay metamorphic complex and samples for this study It is accepted, that the Siberian Craton was assembled during the Paleoproterozoic collision of few blocks, which were then intruded by post-collisional felsic magma (e.g. Gladkochub et al., 2006; Levitskii et al., 2002; Poller et al., 2004, 2005; Rojas-Agramonte et al., 2011; Turkina et al., 2012). In general, the time of the collision has been constrained by dating of granulites, whereas the geochemistry of later granites has allowed distinguishing post-collisional magmatism. The timing of metamorphism and magmatism is considered to be in a

Fig. 1. Location of the Sharyzhalgay high-grade metamorphic complex in the global pattern of cratons (A), generalized map of the south-western Siberian Craton (B), and a schematic cross-section in the studied area (C). In A – black shading indicates outcropping parts of the cratons, grey shadings are cratonic parts covered by sediments. Acronyms for some Archean terranes after Condie (1981): Al – Aldan, An – Anabar, C – Chinese, E – Ethiopian, G – Guiana, EA – East Antarctic, I – Indian, K – Kola, Kl – Kaapvaal, L – Liberian, M – Mauritanian, NA – North Atlantic, P – Pilbara, Q – Quzzalian, R – Rhodesian, SF – Sao Francisco, Sr – Superior, Sv – Slave, U – Ukranian, W – Wyoming, Y – Yilgarn.

89

range of 1.84–1.88 Ga, though obviously post-collisional granites postponed the peak of metamorphism. Paleoproterozoic postmetamorphic mafic rocks (mainly dykes), which mark an episode of post-collisional extension, are also known (Ernst et al., 2016; Gladkochub et al., 2010, 2013; Shokhonova et al., 2010). Recently, it was proved that some Pt-bearing mafic-ultramafic massifs belong to the same episode of extension (Mekhonoshin et al., 2016). The Sharyzhalgay metamorphic complex belongs to the Siberian Craton (Fig. 1A). It consists of a number of blocks (Fig. 1B), which are differentiated by the nature of rocks and the grade of metamorphism. The oldest rocks within the Sharyzhalgay metamorphic complex were found in the Onot green-schist belt (Bibikova et al., 1982; Sandimirova et al., 1992). These are tonalite-trondjhemite-granite-diorite intrusions with an age of 3.29–3.39 Ga (Bibikova et al., 2006; Turkina et al., 2013). The Kitoy block, on which our study is focused (Fig. 1B), is dominated by granulites and amphibolites. Two distinct metamorphic episodes in the Neoarchean (2.48–2.56 Ga) and Paleoproterozoic (1.86–1.90 Ga) were recognized (Glebovitskii et al., 2011; Levchenkov et al., 2012; Levitskii et al., 2010; Poller et al., 2004). Reconstruction of protolith compositions for the Kitoy block shows that up to 70% are sedimentary rocks, among which some 20% are carbonates. Igneous proholiths in the Kitoy block are subordinate. The Irkut block is characterized by the same two metamorphic episodes, but the protolith is mainly igneous rocks (80%) with subordinate sedimentary rocks (20%) (Poller et al., 2005; Sal'nikova et al., 2007; Turkina et al., 2012; Urmantseva et al., 2012). Postcollisional granites in both blocks have ages from 1.84 to 1.88 Ga (Aftalion et al., 1991; Didenko et al., 2003; Donskaya et al., 2002; Kirnozova et al., 2003; Levitskii et al., 2002, 2004; Poller et al., 2004, 2005; Sal'nikova et al., 2007). Prior to our study, Paleoproterozoic post-metamorphic mafic rocks in the Kitoy block were known in the form of a dyke swarm with an age of 1864 ± 4 Ma (U\\Pb TIMS on baddelyite on one dyke sample; Gladkochub et al., 2010) and of the Pt-bearing plagioclase peridotite – gabbronorite Malozadoisky massif (1863 ± 1 Ma by U\\Pb TIMS on baddeleyite; Mekhonoshin et al., 2016). A small massif of mafic rocks with a total length of about 400 m was found in the middle branch of the Kitoy river (Fig. 1C). Already in the field two general groups of rocks were recognized, which could be further subdivided into few different rock types on the basis of petrography and chemistry. The first, volumetrically dominant, group is dolerite with large fresh plagioclase crystals. The second group forms a 100 m wide body and few meter-thick dykes. The 100 m wide body contains amphibole-bearing rocks, whereas the dykes have visible mica crystals, which in the field gives them a lamprophyric appearance. The first group of rocks consists of picritic basalts according to major element chemistry and shows trace element affinities of boninitic rocks (Ivanov et al., 2017). Seven of such picritic basalts were subjected to preliminary K\\Ar dating yielding ages from 1941 ± 42 to 3213 ± 110 Ma. In one of the picritic basalt samples 16 zircon grains were recovered, which were analysed by laser ablation inductively plasma mass spectrometry for U\\Pb dating (see details in Supplementary Table S1 and Supplementary Fig. S1). After filtering for concordance only six ages were retained: 1800 ± 7 Ma, 1815 ± 9 Ma, 1915 ± 4 Ma, 2430 ± 13 Ma, 2440 ± 50 Ma and 2485 ± 7 Ma (propagated errors reported at 1σ level). The bimodal distribution of ages is in general agreement with the age of the country rocks and shows that all the zircon grains in the basalts are inherited. Since, the age of the picritic basalts remains unconstrained we do not consider them further in the paper. 3. Methods Whole-rockmajor-element compositions were determined by x-ray fluorescence (XRF) on fused pellets (Afonin et al., 1984) with adaptations to a S4 Pioneer (Bruker ASX, Germany) spectrometer. Trace elements in all studied rocks, except samples D170, D186 and D219,

90

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

were analysed by three methods: Li, Rb – by flame photometry (FP), B, Be, Sc, V, Cr, Co, Ni, Zn, Cu, Ga, Ge, Sr, Ba, Zr, Mo, Sn, Tl, Pb, La, Ce, Nd, Yb, Y – by atomic emission spectrometry (AES) and V, Cr, Co, Zn, Ga, Ge, Rb, Sr, Y, Zr, Nb, Mo, Sn, Cs, Ba, La, Ce, Pr, Nd, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, Lu, Hf, Ta, W, Tl, Pb, Th, U – by inductively coupled plasma mass spectrometry (ICP-MS). Analytical procedures for FP and AES were routine and are described in Fabrikova (1961) and Vasil'eva et al. (1997), respectively. Sample preparation for ICP-MS was lithium metaborate fusion described in details by Panteeva et al. (2003). When elements were analysed by two different methods, the results were in good agreement for all elements except V. In all diagrams, however, we use values only

Table 1 Major (wt%) and trace (ppm) elements in the studied rocks. Type

Low-Mg group

Number

D220

D222

C34

D219

D231

D170

D225

D186

SiO2 TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 LOI Total

49.10 0.99 13.45 9.41 0.14 7.37 8.02 2.46 3.02 0.56 5.41 99.92

50.08 1.07 12.94 10.07 0.14 8.90 9.60 2.39 2.51 0.72 1.26 99.68

54.39 1.27 15.11 9.16 0.14 5.08 7.84 2.38 2.60 0.49 1.76 100.22

49.38 1.09 13.87 10.14 0.14 8.86 9.43 2.36 2.61 0.72 1.38 99.69

50.51 0.77 11.58 9.28 0.15 12.54 9.46 1.85 2.71 0.46 0.46 99.76

50.67 0.32 11.58 10.07 0.17 15.02 8.91 1.19 0.97 0.16 0.80 99.86

50.72 0.33 11.80 10.08 0.17 14.77 8.99 1.31 0.97 0.18 0.50 99.81

49.39 0.73 10.42 9.72 0.17 15.18 9.80 0.48 2.78 0.36 0.77 99.805

B Be Sc V Cr Co Ni Zn Cu

FP 34 AES 2.2 1.5 24 180 520 46 150 120 40

FP 44 AES 2.1 1.5 29 160 520 47 170 100 38

FP 40 AES 14.5 2 35 140 130 31 68 110 18

FP 172 AES 1.7 1.5 26 160 580 45 130 91 32

FP 20 AES 5.2 2.2 25 180 1500 54 200 110 62

FP 30 AES 9.8 b 0.8 41 160 2400 67 390 100 110

FP 30 AES 5.4 0.4 37 210 2000 60 220 74 80

FP 76 AES 3.2 2.8 21 190 1800 58 340 69 51

ICP-MS 15.9 1.54 80.1 818 30.4 175 9.00 1.2 2 1.67 1835 103 224 24.4 93.9 16.5 3.35 11.3 1.34 6.13 1.09 2.80 0.379 2.19 0.340 4.48 0.456 0.433 0.607 24.1 10.2 1.92

ICP-MS 16.4 1.57 79.1 756 30.7 211 9.17 1.2 2 1.96 1756 107 236 25.5 98.6 17.0 3.36 11.5 1.35 6.31 1.18 2.84 0.380 2.32 0.350 5.10 0.462 0.394 1.12 21.2 9.89 1.89

ICP-MS 18.4 1.60 91.6 563 38.6 240 11.2 0.5 2 3.18 1376 70 147 16.2 62.8 12.3 2.63 9.41 1.33 6.96 1.42 3.65 0.509 3.28 0.461 6.08 0.660 0.748 0.431 19.4 7.38 1.33

ICP-MS 13.9 1.48 119 580 27.5 205 8.09 2.5 2.4 2.58 1658 74.8 155 16.8 64.6 13.3 2.86 9.51 1.18 5.63 0.925 2.30 0.299 1.76 0.250 5.19 0.441 0.736 0.685 21.6 14.9 2.39

ICP-MS

Ga Ge Rb Sr Y Zr Nb Mo Sn Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta W Tl Pb Th U

ICP-MS 11.6 1.71 36.7 527 17.0 67.3 3.65 0.34 1.7 2.07 1275 91.4 181 17.6 62.3 9.49 1.72 6.46 0.737 3.42 0.613 1.63 0.231 1.53 0.250 1.74 0.298 0.689 0.357 26.3 12.4 2.24

Li

High-Mg group

14 1.3 74 364 27 286 1.4 1.5 1 1930 100 190 91

2.5

0.67 23

28 442 14 40 1.2 1.8 1170 81 160 37

1.8

27

1.2 196 456 22 183

15 1250 59 130 56

1.5

4.7

from one method as indicated in Table 1. The 87Sr/86Sr, 143Nd/144Nd, Rb/86Sr, and 147Sm/144Nd isotopic ratios (Table 2) were obtained by isotope dilution TIMS using a Finnigan MAT 262 mass-spectrometer (see for details of the procedure Ivanov et al., 2018). The lead isotope analysis of ancient rocks is not trivial. Therefore, crushed rock fractions of 0.5–0.25 mm size were initially pretreated by 2.2 N HCl on a hot plate for one hour, then washed by distilled water, dried and powdered. Powders were split into two aliquots for each sample. One aliquot was digested by using a mixture of concentrated HF-HNO3-HClO4 in a closed vessel during two days at 120 °C. Another aliquot was subjected to sequential leaching – first by 5 N HCl and then by concentrated HNO3. Only the final leachate after HNO3 was used for analysis. Further procedure followed that of Manhes et al. (1984). Isotope ratios were measured using a ThermoFisher Triton mass-spectrometer. Lead isotopic data are provided in Table 3. For K\\Ar dating, argon isotope ratios and potassium concentrations were measured on two aliquots using ARGUS VI mass-spectrometer and SOLAAR M atomic absorption spectrometer, respectively. Prior the K\\Ar dating, all samples were treated by 4 N HNO3. Description of the analytical procedure was described elsewhere (Peretyazhko et al., 2018). Conventional 40K decay constants were used for the age calculation (Steiger and Jäger, 1977). U\\Pb dating was performed on air abraded zircon grains using isotope dilution TIMS following the procedure described by Corfu (2004). One zircon grain was chemically abraded as indicated in Table 4. Examples of zircons before and after abrasion are shown in Supplementary Fig. S2. The tracer composition for the ID-TIMS measurements has been harmonized with that of the EARTHTIME ET235 spike (Corfu et al., 2016). The instrument used was a Finnigan MAT 262 massspectrometer. Electron microprobe (EMP) analyses using JXA-8200 analyser and scanning electron microscopy with energy dispersive X-ray spectrometry (EM-EDS) analyses using Hitachi SU-70 were conducted for assessing mineral chemistry in two rock samples as described below in a section on mineralogy. Some of the analytical work was done using equipment belonging to three analytical centres in Irkutsk, Russia – Centre for Geodynamics and Geochronology, Institute of the Earth's Crust SB RAS (Sr-Nd-isotopes and K\\Ar dating), Centre of geochemical and isotopic studies, A.P. Vinogradov Institute of Geochemistry SB RAS (XRF, FP, AES, EMP), Centre of Ultramicroanalysis, Limnological Institute SB RAS (ICP-MS). Lead isotope analyses were conducted at the Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences, St-Petersburg. EM-EDS analyses were performed at the University of Tasmania. U\\Pb dating was done at the University of Oslo. 87

4. Results 4.1. Rock petrography, mineralogy and classification using major and trace elements Already in the field, the group of lamprophyric-looking rocks was subdivided into three types. The first type, which is referred to as lowMgO type, consists of massive, fine-grained dykes of black and dark brownish colours. They contain tiny plates of plagioclase and abundant small crystals of mica. A view of this rock type under the polarizing microscope, using sample D220 as an example, is provided in Fig. 2A,B. This sample consists of (from higher to lower abundances) plagioclase, mica, orthopyroxene, olivine, K-feldspar and quartz. Accessory minerals are various sulphides, apatite, monazite and opaque oxides and hydrooxides. Alteration is minor, though some samples contain secondary colourless or light green amphiboles, light brown biotite, light green epidote and white mica. In a SiO2-K2O classification diagram rocks of this type fall into the field of absarokites and they have relatively low-MgO (b 10 wt%) and moderately high TiO2 (N1 wt%) concentrations (Fig. 3).

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

91

Table 2 Sr-Nd-isotopic data for the studied samples. Sample

87

Rb/86Sr

87

Sr/86Srmeas

D220 D222 C34 D170 D225 D231

0.2509 0.3066 0.4373 0.1805 0.2181 0.4691

0.711778 0.712813 0.717539 0.710510 0.710490 0.718463

87

Sr/86Srt

0.705044 0.704585 0.705804 0.705666 0.704635 0.705875

Sm/144Nd

εSrt

147

+38.9 +32.4 +49.7 +47.8 +33.1 +50.8

0.0772 0.0801 0.0779 0.1192 0.1428 0.1099

The second type, referred to as high-MgO type and closely associated with the first type in form of small dykes, was distinguished in the field by its dark green colour. A representative thin section of sample D170 is shown in Fig. 2C,D. This sample consists of (from higher to lower abundances) clinopyroxene, orthopyroxene, plagioclase, mica and olivine. Some samples contain minute grains of carbonates (calcite and ankerite). Accessory minerals are zircon, titanite, ilmenite, apatite, monazite and orthite. Some samples, or zones within samples, contain secondary colourless or light green amphiboles, light green epidote, zoisite, chlorite and light brown biotite. The rocks of these type are higher in MgO (N12 wt%), lower in TiO2 (b 0.8 wt%) and in the SiO2-K2O classification diagram they fall into the field of absarokites or some in the field of basalts (Fig. 3). Sample C34, collected from a 100-m wide magmatic body, is distinctly richer in hornblende. Major minerals in this sample are the following (from higher to lower abundances): plagioclase, hornblende, clinopyroxene, orthopyroxene, mica, quartz, olivine, K-feldspar, ilmenite, magnetite and various sulphides. Accessory minerals are apatite, zircon, titanite, monazite and orthite. A view of a thin section of sample C34 under polarized microscope is shown in Fig. 2E,F. This sample plots in the field of high-K basaltic andesites in the SiO2-K2O classification diagram, and it has the lowest MgO and highest TiO2 concentrations among the set of rocks considered (Fig. 3). However, because sample C34 is the only sample with hornblende in our collection, we do not separate it in a special third type and group it together with other low-MgO lamprophyric samples on all diagrams. According to the classification principles of the International Union of Geological Sciences (Le Maitre et al., 2002) our samples are classified as kersantite (high- and low-MgO groups) and spessartite (sample C34). However, irrespective of the names given, they are all characterized by shoshonitic geochemical affinities and resemble Tibetan postcollision shoshonitic series (Figs. 4–6). The exception is sample C34, which in the Th/Yb vs Ta/Yb diagram plots in the field of calc-alkaline rocks (Fig. 5). However, in the Ce/Yb vs Ta/Yb diagram this sample is

143

Nd/144Ndmeas

143

0.511144 0.511038 0.511179 0.510742 0.511030 0.510712

Nd/144Ndt

0.510196 0.510055 0.510223 0.509280 0.509278 0.209363

εNdt

TDM, Ga

−0.5 −3.3 −0.0 −18.5 −18.6 −16.9

2.1 2.3 2.1 3.8 4.5 3.5

also located in the shoshonitic field (Fig. 5). The most puzzling are the compositions of high-MgO lamprophyric samples D170 and D225. Although they are located in the field of Tibetan shoshonitic series in the major element ratio diagram (Fig. 4), in the trace element ratio diagram (Fig. 5) and in the primitive-mantle normalized diagram (Fig. 6), their potassium content is too low for the shoshonitic series (Fig. 3). K2O/ Na2O ratio in these two samples is b1 (Table 1), whereas shoshonitic series is characterized by K2O/Na2O ratio N 2. According to the EMP and EM-EDS analyses, samples D170 and D225 are similar to each other, though sample D170 has a wider association of minerals. Rock-forming minerals in both samples are: enstatite, olivine, augite, plagioclases (bytownite, oligoclase) and phlogopite. Accessory minerals are the following: chromite, fluorapatite, monazite-(Ce), pyrite (with Ni up to 0.7 wt%), magnetite (with Ni up to 0.7 wt%), Timagnetite, zircon, ferrodolomite, quartz and ferrobrucite. Selected chemical compositions of the rock-forming minerals and chromite are provided in Supplementary Tables S2–S5. In addition to the above mentioned minerals sample D170 includes chromium-rich pyroxenes, amphiboles and muscovite (though it cannot be totally excluded hat Cr is not from tiny chromite inclusions). It also contains large amount of rare earth element phosphates: Monazite-(Ce, La), monazite-(Th, Ce, La), rhabdophane-(La, Ce), britholite-(Ce), brockite, cerite and a number of not identified phosphates. Also identified were calcite, barite, chalcopyrite, samaniite, fluorite and the rare alumosilicates anandite, trattnerite and cronstedtite. The spectrum of accessory and rare minerals is very unusual, for example, ferrobrucite, which is found as an intergrowth with orthopyroxene and plagioclase, and as inclusions in plagioclase (Fig. 7). Most likely, ferrobrucite is the product of alteration of ferropericlase, which can crystallize at the upper mantle conditions in regions with low silica activity (Brey et al., 2004). The importance of this mineral will be discussed later. The mineralogical composition of the high-MgO lamprohyres merits a special mineralogically-focused study. What is important for our study

Table 3 Uranium-lead isotope data for the studied samples. Sample

Pb, ppm

U, ppm

238 204

D-170 D-170 D-225 D-225 D-186 D-186 D-219 D-219 D-220 D-220 D-222 D-222 D-231 D-231

WR L WR L WR L WR L WR L WR L WR L

17

0.761

15

0.626

2.2

0.339

17

0.447

19

0.669

15

0.639

16

0.528

U Pb

2.80 7.15 2.47 5.13 9.89 13.8 1.54 2.95 2.07 2.77 2.48 1.76 1.99 1.48

206 204

Pb Pb

16.036 17.131 15.393 16.418 18.362 20.667 15.374 16.438 15.549 16.920 15.528 16.489 15.401 16.427

WR - whole rock, L - leachate in HNO3. ⁎ Initial ratios, corresponding to age of 1865 Ma. ⁎⁎ Model age (Stacey-Kramers two stage model) in Ga. ⁎⁎⁎ Model μ (Stacey-Kramers two stage model). ⁎⁎⁎⁎ Pb-Pb isochron age and MSWD.

2 s, %

207 204

0.06 0.06 0.06 0.06 0.06 0.12 0.06 0.06 0.06 0.06 0.06 0.06 0.06 0.06

Pb Pb

15.181 15.3 15.097 15.21 15.588 15.877 15.191 15.319 15.224 15.391 15.211 15.344 15.160 15.353

2 s, %

208 204

0.09 0.09 0.09 0.09 0.09 0.12 0.09 0.09 0.09 0.09 0.09 0.09 0.09 0.09

Pb Pb

38.753 46.461 37.77 42.836 38.008 41.243 36.409 43.511 36.678 42.407 36.768 39.725 35.919 38.306

2 s, %

208 204

0.12 0.12 0.12 0.12 0.12 0.15 0.12 0.12 0.12 0.12 0.12 0.12 0.12 0.12

Pb Pb(t)⁎

208 204

Pb Pb(t)⁎

TSK⁎⁎

mSK⁎⁎⁎

t, Ga⁎⁎⁎⁎

MSWD

1.88±0.16

0.18

2.12±0.04

0.84

15.09

15.07

1.86

8.95

14.56

15.00

2.24

9.48

15.03

15.21

2.13

10.13

14.86

15.13

2.17

9.89

14.85

15.14

2.20

10.02

14.69

15.12

2.30

10.19

14.73

15.08

2.21

9.78

92

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Table 4 U\ \Pb isotope data for the dated zircon grains. Weight [ug] Pb ppm

U ppm

Th/U Pbc ppm

Pbcom Pg

206 204

Pb Pb

207 235

Pb U

2s abs

206 238

Pb U

2s abs

rho

207 206

Pb Pb

2s abs

206 238

Pb 2 s U abs

207 235

Pb 2 s U abs

207 206

Pb 2 s Disc. Pb abs %

Sample D170 BR AA 2 BR AA 4 BR AA 2

1034 605 558

3279 0.10 1879 0.10 1722 0.08

0 0.06 0

5.5 6.0 2.2

23,353 24,998 31,725

5.027 0.015 5.138 0.014 5.203 0.014

0.3213 0.00084 0.97 0.1135 0.3277 0.00083 0.97 0.1137 0.3314 0.00077 0.97 0.1139

0.0001 0.0001 0.0001

1796 1827 1845

4.1 4.0 3.7

1824 1842 1853

2.5 2.4 2.2

1856 1860 1862

1.4 1.2 1.2

3.7 2.0 1.1

Sample D225 BR AA 1 BR-CL AA 1 LP CA 1 CL AA 1

1231 1203 143 1862

3859 3907 490 6978

0 6.24 0 20.48

4.0 12 1.7 27

19,452 6083 5123 4220

5.123 4.890 4.602 4.121

0.3269 0.3139 0.2970 0.2699

0.0001 0.0001 0.0001 0.0001

1823 1760 1676 1540

5.1 4.2 3.4 6.5

1840 1800 1750 1658

2.9 2.6 2.2 4.1

1859 1848 1838 1811

1.3 1.6 1.6 2.0

2.2 5.4 10 17

0.07 0.08 0.12 0.10

0.018 0.015 0.012 0.021

0.00105 0.00086 0.00069 0.00128

0.98 0.96 0.94 0.98

0.1137 0.1130 0.1124 0.1107

AA = air abrasion. CA = chemical abrasion.

Fig. 2. Thin sections of lamprophyric-looking rocks under a polarized microscope. The figure shows variable amount of magmatic mica in the samples as well as little signs of alteration. Mineral abbreviation: Cpx – clinopyroxene, Fsp – K-feldspar, Ol – olivine, Opx – orthopyroxene, Pl – plagioclase, Phl – phlogopite, Qtz – quartz.

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Fig. 3. K2O vs SiO2 and TiO2 vs MgO diagrams for the studied samples in comparison with Tibetan shoshonite series rocks, including related ultrapotassic rocks (Campbell et al., 2014; Hebert et al., 2014; Turner et al., 1996; Williams et al., 2004; Xu et al., 2001). Rock classification lines are after Peccerillo and Taylor (1976).

is that such a variable mineral assemblage could not crystallize in equilibrium from the same magma. The high-MgO rocks are hybrid in origin. 4.2. Geochronology Because the rocks found in the Kitoy river area (Fig. 1C) are fresh and devoid of significant post-eruption modifications, it was initially not

Fig. 4. K2O/TiO2 vs Al2O3/TiO2 diagram for the studied samples in comparison with Tibetan shoshonite series rocks, including related ultrapotassic rocks (the same samples as in Fig. 3 with SiO2 b 55 wt% and K2O b 6 wt%). Fields of orogenic and anorogenic lamproites are shown with a formal divider at Al2O3/TiO2 = 4 (Ivanov et al., 2018). Among large globally distributed samples of anarogenic lamproites, only Mesozoic lamproites from the Siberian Craton plot in the field of orogenic lamproites (Ivanov et al., 2018).

93

clear if they belong to a Precambrian, Paleozoic or Mesozoic magmatic event. For a preliminary assessment of their age, two samples, namely low-MgO lamprophyre (sample D220) and high-MgO lamprophyre (sample D225), were subjected to K\\Ar dating. They yielded surprisingly old (for their fresh-looking petrography) ages of 1919 ± 42 Ma and 1814 ± 58 Ma, respectively. As a father test, two high-MgO lamprohyres (samples D170 and D225), which contain accessory zircons, were processed for U\\Pb dating. The zircon grains are mostly anhedral and subequant and rarely euhedral prismatic, have a brownish to white appearance reflecting very high U contents (mostly 1000 to 7000 ppm) and metamictization, but they are not fractured, and are generally unaltered (Fig. S2). The analyses from the two samples lie on the same discordia line with an upper intercept with concordia at ~1865 Ma and an approximately Devonian lower intercept age (Fig. 8A). The slight scatter is likely due to minor secondary Pb loss remaining in some of the grains after abrasion. The three analyses of sample D170 alone are 1 to 4% discordant and perfectly collinear (MSWD = 0.01) yielding an upper intercept age of 1864.7 ± 1.8 Ma (Fig. 8B). The peculiar but consistent features of the zircon grains, such as the very high U contents, low Th/U, anhedral to locally euhedral shapes but lack of fracturing (reflecting lack of internal zoning), and relatively low degree of discordance in spite of the advanced metamictization strongly support an indigeneous magmatic origin of the zircons. If they were inherited from the country rocks or the source in a much younger magma they would likely form a heterogeneous population enriched in low-U zircon grains and with much more discordant U\\Pb systems as seen with the example of inherited zircons from one of the picritic basalt samples (Supplementary Fig. S1 and Supplementary Table S1). These arguments show that 1864.7 Ma is the true age of the high-MgO lamprophyres. Interestingly, the K\\Ar age for the same sample overlaps within analytical error with the U\\Pb age, and thus testifies to the lack of significant alteration in this sample. The K\\Ar age of sample D220 is slightly older if only the analytical errors are considered. The K\\Ar ages calculated with conventional 40K decay constants (Steiger and Jäger, 1977) are systematically younger by 1% compared to U\\Pb ages (e.g. Ivanov, 2006; Min et al., 2000), and thus the K\\Ar age is even older relative to U\\Pb age. However, it is not expected that the K\\Ar system may remain truly closed for Paleoproterozoic rocks. The older K\\Ar age may be due to either some K-loss but radiogenic argon retained, or some excess of 40Ar. 4.3. Sr-Nd-Pb isotopes In the 206Pb/204Pb -207Pb/204Pb diagram (Fig. 9), the whole-rock and HNO3 leachate fractions of samples D225 and D170 form an isochron with a slope of 1.88 ± 0.16 Ga (Table 3), which within (though large) error agrees with the U\\Pb zircon ages of these two samples (Fig. 8). The whole-rock and HNO3 leachate fractions of the other samples, irrespective of their high- or low-MgO composition, lie on a line with a slope of 2.12 ± 0.04 Ga (Fig. 9, Table 3). This date predates the age of the metamorphism, whereas the samples show no sign of metamorphic modifications. Thus, from a geological point of view, this date is meaningless. The good linearity for other than the D225 and D170 samples is due to mixing of melts derived from two sources with different lead isotope compositions. Stacey and Kramers (1975) model ages for the studied samples vary from 1.86 to 2.32 Ga (Table 3). The older than real model ages can be explained by a three-stage lead evolution (Fig. 10A): t1 – derivation of crust with μ = 9.7 from the mantle at ~3.9 Ga, t2 – loss of U relative to Pb due to formation of the lower crust and setting new μ = 5.5 at ~ 2.9 Ga, t3 – melting of the lower crust and enriched mantle sources and mixing at 1.865 Ga which lead to variations in initial lead isotope ratios in the studied samples (Fig. 10B). Sample D186 with the highest initial 206Pb/204Pb ratio is an exception. It lies away from the mixing line between the lower crust and enriched mantle sources discussed above

94

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Fig. 5. Ce/Yb and Th/Yb vs Ta/Yb diagrams for the studied samples in comparison with Tibetan shoshonite series rocks, including related ultrapotassic rocks (the same samples as in Fig. 3 with SiO2 b 55 wt% and K2O b 6 wt%). Boundaries for tholeiitic, sub-alkaline and shoshonitic series area after Pearce (1982).

but is located close to a mixing line between the upper crust and depleted mantle sources (Fig. 10B). Knowing the age of the rocks allows recalculation of Sr\\Nd isotopic ratios to initial values. From Table 2 and Fig. 11, it is seen that the lowand high-MgO lamprophyres are principally different in terms of εNdt. The low-MgO lamprophyres are characterized by zero or just slightly negative values, whereas the high-MgO lamprophyres have negative values as low as −18, which requires either low crustal or metasomatized mantle lithologies in the source (e.g. Menzies, 1989). However, there is no difference in terms of εSrt values. In εSrt vs εNdt diagram the high- and ultraK rocks worldwide define two general trends (e.g. Bergman, 1987; Foley et al., 1987; Ivanov et al., 2018; Lustrino et al., 2016). The trend of negative correlation with larger variations in εSr and moderate variations in εNd is a remarkable isotopic characteristic of orogenic lamproites and shoshonites (Fig. 11). The trend of larger variations in εNd and less significant variations in εSr is an isotopic feature of anorogenic lamproites (Fig. 11). Generally speaking, the Paleoproterozoic low-MgO and high-MgO lamprophyres from the south-western Siberian Craton are located close to the orogenic and anorogenic trends, respectively (Fig. 11).

5. Discussion 5.1. Unique freshness of studied lamprophyres The studied lamprophyres are exceptionally fresh for being Paleoproterozoic rocks, which makes them a unique object for study. For example, selected images of thin sections shown in Fig. 2 are free of recognizable secondary minerals. Phlogopite, plagioclase and olivine grains, which can be easily altered, are not substituted by secondary minerals (Fig. 2). The minute grains of low temperature minerals (e.g. carbonates, chlorite, cronstendtite, ferrobrucite), which were found in some of the samples either under polarizing microscope or using EMP and EM-EDS, apparently are the result of posteruption autometasomatic processes. The absence of secondary modifications and later thermal overprints is also evident in the practically identical U\\Pb age and whole-rocks K\\Ar ages. Lamprophyres cut through picritic dolerites of unknown age with a boninite geochemical flavour (Fig. 1C). These picritic dolerites are also free of secondary alterations (Fig. 2 in Ivanov et al., 2017). By contrast, rocks chemically similar to lamprophyres, but severely altered and even metamorphosed to granulite facies have been found within the Irkut block of the Sharyzhalgay complex (Levitskii, 2012). Thus, the unique freshness of the lamprophyres is due to the unique preservation conditions of the Kitoy block. It appears, that at 1885 Ma this block was already in subsurface position and since it has never been buried or significantly uplifted. 5.2. Assessment of crustal contamination and crystal fractionation

Fig. 6. Primitive mantle (McDonough and Sun, 1995) normalized trace element diagram for studied samples in comparison with Tibetan shoshonite series rocks, including related ultrapotassic rocks (the same samples as in Fig. 3 with SiO2 b 55 wt% and K2O b 6 wt%).

Alkaline melts are aggressive and may react with ambient rocks. The studied low-MgO lamprophyre samples contain quartz and K-feldspar grains (Fig. 2), which at first glance may be considered as evidence of crustal contamination. However, the lowest εNdt values, down to −18, are found in high-MgO lamproite samples, which contain neither quartz nor K-feldspar, whereas quartz- and K-feldspar-bearing lamprophyres are characterized by εNdt values close to zero. In case of crustal contamination, we would expect the lowest εNdt values in rocks with quartz and K-feldspar, not the opposite. Thus we consider that there was no significant crustal contamination, which affected mineralogy, trace element geochemistry and isotope composition of the studied lamprophyre samples. High-MgO and low-MgO lamprophyres belong to two isotopically distinct groups and are not related by fractional crystallization process. The high-MgO samples contain nearly primitive olivine with Fo number

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

95

Fig. 7. Back scattered electron (BSE) image of samples D170 (A) and corresponding spectrum of probable ferrobrucite (B) and BSE image of sample D225 (C) and corresponding spectra of ferrobrucite (D) and plagioclase (E). Samples D170 and D225 were analysed with a microprobe JXA 8200 and scanning electron microscope with energy dispersive X-ray spectrometer Hitachi SU-70. Mineral acronyms in (A) are Fbrc – ferrobrucite, Ol – olivine, En – enstatite, Pl – plagioclase (bytownite-oligoclase), Act – actinolite.

N 0.90 and thus underwent minor crystal fractionation. In the low-MgO samples olivine is less rich in magnesium, but still its Fo number is N87. Though some fractional crystallization occurred, it did not significantly affect the chemical composition of the final melts. The low-MgO samples in addition to olivine contain quartz, which could not have crystallized in equilibrium. 5.3. Origin of high- and low-Mg lamprophyric melts For the studied rocks there is a positive correlation between MgO and Ni and Cr, no correlation between MgO and most of the incompatible trace elements, as exemplified by the MgO vs Ba diagram, and there is a negative correlation between MgO and heavy rare earth elements, such as exemplified by the MgO-Yb diagram (Fig. 12). At the same time, high-Mg lamprophyres are characterized by negative εNdt values (Fig. 11, Table 2). Calculated Nd TDM one-stage model ages for low-Mg and high-Mg lamprophyre samples are 2.1–2.3 Ga and 3.5–4.5 Ga, respectively (Table 2). The oldest ages within the Sharyzhalgay block are limited to two intervals of 2.48–2.56 and 3.29–3.39 Ga (Bibikova

et al., 2006; Levchenkov et al., 2012; Poller et al., 2005; Turkina et al., 2013). Thus, the Nd TDM model ages do not match the ages of the Sharyzhalgay block (especially in the case of high-MgO lamprophyres) and likely show that the source was not within the lithosphere of the Siberian Craton. (Note: use of a two-stage neodymium evolution model shifts the model ages to even older values, including the meaningless age for one of the high-MgO samples, which becomes older than the age of Earth). Both low- and high-Mg lamprophyres have high concentrations of the majority of highly incompatible trace elements and a notable depletion in high field strength elements relative to neighbouring elements, as seen in the primitive mantle normalized diagram (Fig. 6). To reconcile such unusual features, we suggest that high-Mg lamprophyres crystallized from a hybrid melt, produced first by melting of recycled crustal lithology within the subducted lithospheric plate and then by contamination in the orthopyroxene-rich mantle wedge. This is similar to a model suggested recently for the origin of East Tibetan shoshonites, which have similar trace elements signatures to our high-MgO lamprophyres (Campbell et al., 2014).

96

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Fig. 8. Concordia plot for samples D170 and D225. A – zircons from both samples lie on the same discordia with upper and lower intercepts at ~1865 and ~380 Ma, respectively. Symbol sizes in A are large that 2 sigma analytical error. B – concordia plot for the sample D170. The upper intercept of the discordia line at 1864.7 ± 1.8 Ma is considered as the crystallization age. Ellipses in B represent 2 sigma analytical errors.

The hybrid nature of the high-Mg lamprophyres is evident from their mineral variability, presence of disequilibrium mineral assemblage (forsteritic olivine and quartz, for example). Ferrobrucite, which likely is the secondary mineral of ferropericlase, was found as inclusions in plagioclase and in intergrowth with orthopyroxene and plagioclase (Fig. 7). Ferropericlase was also found as inclusions in some kimberlitic diamonds (Kaminsky, 2012; Stachel et al., 2000; Zedgenizov et al., 2001) and this was taken as evidence for diamond crystallization at lower mantle conditions (Kaminsky, 2012). In our case, it is improbable that ferropericlase (precursor of ferrobrucite) is a mineral of the lower mantle because of the association with plagioclase and orthopyroxene. Most

Fig. 10. Three-stage model for the evolution of lead isotopic compositions in 206Pb/204Pb -207Pb/204Pb diagram. DMK-T – model curve of the isotopic lead composition for depleted upper mantle after (Kramers and Tolstikhin, 1997). Other model isotopic lead evolution curves: EM – in enriched mantle, LC – in the lower crust, UC – in the upper crust. Sample points are initial lead isotope ratios calculated to 1865 Ma.

Fig. 9. 206Pb/204Pb -207Pb/204Pb diagram for measured values of whole-rock and HNO3 leachate fractions. Regression lines with the calculated ages for the subsets of the samples are shown in grey. Leachates and whole-rock analyses of the same samples are connected by black lines.

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

97

Fig. 11. εSr vs εNd for the studied samples in comparison with Tibetan shoshonite series rocks, including related ultrapotassic rocks (the same samples as in Fig. 3 with SiO2 b 55 wt% and K2O b 6 wt%) and Mesozoic lamproites from different parts of the Siberian Craton (Davies et al., 2006; Ivanov et al., 2018).

likely, it crystallized in the upper mantle in regions with low silica activity such as carbonated dunites (Brey et al., 2004). Abundant orthopyroxene, plagioclase and probably olivine in the high-Mg rocks are, at least partially, xenocrystic in origin. Alternatively, high-MgO lamprophyres could be derived from metasomatized lithospheric mantle. Because of the Nd TDM model ages mismatch with the real ages of the Siberian Craton, such metasomatized lithospheric mantle must have been a part of the subducted slab. A similar subducted lithospheric source has been suggested, for example, for the origin of Variscan lamprophyres of SW England (Dupuis et al., 2016). The low-Mg lamprophyres, though exhibiting Sr\\Nd isotope signatures that approach those of mantle-derived melts, are too low in MgO to have been produced from melting of mantle lithologies. At the same time, in the primitive mantle normalized diagram their trace element patterns show a notable similarity with those of the high-Mg lamprophyres (Fig. 6). We speculate that the low-Mg lamprophyres crystallized from melts, which were also derived from melting of crustal lithologies at mantle depth, but these rocks had positive εNdt values at 1865 Ma. The only candidate for such rocks is a mixture between mafic oceanic crust and subducted sediments. A contribution of the subducted sediments into the mantle source of mafic melts is usually assessed by using radiogenic and stable isotopes. Such studies limit the amount of the recycled sediments to a few per cent for basalts and andesites (Hergt et al., 1991; Nebel et al., 2011; Plank and Langmuir, 1998; Stern and Kilian, 1996), but allow up to 100% of sediments for K-rich melts, such as lamproitic melts (e.g. Ivanov et al., 2018; Murphy et al., 2002). Assuming εNdt values of +10 and − 20 for the mafic oceanic crust and subducted sediments, respectively, paired with Nd concentrations of average mid-ocean ridge basalts (Nd ~ 12, Gale et al., 2013) and bulk continental crust (Nd ~ 16, McLennan, 2001), a mixture of ~7:3 of these two components is required to yield the εNdt values measured for the low-MgO lamprophyres. Melts produced from such a mixed source apparently did not pass through the mantle wedge and was not contaminated by orthopyroxenitic mantle. It could have been contaminated by crustal material, however, when the shoshonite melt stagnated in the crust and fractionated to the low-Mg composition. In general, this picture agrees with the variations of lead isotope compositions (Fig. 10B), which suggests contribution of lower crustal material (mafic oceanic crust) and enriched mantle material (here, mantle wedge). A graphical representation of the origin of the highand low-Mg lamprophyres is shown in Fig. 13B. Both types of melts

Fig. 12. Yb, Ba, Cr, Ni vs MgO for the studied samples. Dashed lines connect the low-MgO lamprophyric sample C34 with εNdt = 0 to the hypothetic average composition of xenocrystic mantle material with high MgO, Cr and Ni, and low concentrations of heavy rare earth elements. The projected zero concentrations of Yb limits MgO in such an average composition to less than ~25 wt%.

were derived from melting of subducted material, but only some of the resulted melts entrained mantle material from the mantle wedge, which was modified from initially dominantly olivine to dominantly orthopyroxene lithology due to interaction of felsic melts with olivine at a previous stage of subduction and collision. 5.4. From collision to extension Post-collisional magmatism is often explained by slab break-off, which allows hotter asthenospheric flow to penetrate underneath overriding plate, where it heats and interacts with lithosphere and hydrated mantle wedge (e.g. Keskin, 2003; Liegeois and Black, 1987; von Blanckenburg and Davies, 1995; Zhu et al., 2015). Numerical modeling

98

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Fig. 13. A model of change from collisional (A) to post-collisional (B) tectonic stages. This model was inspired by work of Zhu et al. (2015) for development of the India-Asian collision system, but it takes into account the known facts about metamorphism and magmatism of the south-western Siberian Craton. Acronyms: CAV – continental arc type volcanoes, OIB – ocean island basalt type volcanoes, CC – continental crust, SCLM – subcontinental lithospheric mantle, MW – mantle wedge, MC – mafic crust, SOLM – suboceanic lithospheric mantle.

suggests that the time span between initiation of continental collision and slab break-off varies from 10 to N20 million years for young (weak) and old (strong) slabs, respectively (van Hunen and Allen, 2011). The time difference from collision to post-collisional extension in our example fits well within this time interval: from the peak of metamorphism to the peak of extension it took about 15 million years (Fig. 14). A model for such a rapid change from one to another tectonic regime is shown in Fig. 13. According to this model, which is based on the modern example of the India-Asia system (Zhu et al., 2015), the collision of two continental blocks was driven by slab pull force, but when the continental blocks finally collided the slab broke off. This resulted in the loss of the slab pull force causing mountain growth and stress relaxation. Hot asthenospheric material penetrated through the slab window to the mantle wedge and heated metasomatized sublithospheric

Fig. 14. Summary of available geochronological data for the south-western Siberian Craton. Original data for metamorphism and post-collisional granites (Aftalion et al., 1991; Didenko et al., 2003; Donskaya et al., 2002; Kirnozova et al., 2003; Levchenkov et al., 2012; Levitskii et al., 2002, 2004; Poller et al., 2004, 2005; Sal'nikova et al., 2007; Turkina et al., 2009, 2012) are represented in form of kernel density estimation plot using DensityPlotter (Vermeesch, 2012) with setting adaptive mode and bandwidth of 10 Ma. Data for OIB-type dykes and Pt-bearing mafic-ultramafic intrusions are after Gladkochub et al. (2010) and Mekhonoshin et al. (2016), respectively. The age for the shoshonitic-type lamprophyric dykes is from this study. The vertical grey bar represents the mean of the three ages.

mantle producing OIB-type magma. At the same time, heating of the upper portion of the slab attached to the lithosphere resulted in melting of metamorphosed sediments and oceanic crust with the production of shoshonitic melts. Relaxation of the lithospheric stress provided the possibility for melting of the crust, which was already hot, and formation of post-collisional felsic melt that intruded the orogeny and crystallized as granites.

6. Conclusions Unusually fresh lamprophyric-looking rocks of in the middle branch of the Kitoy river within the boundaries of the Sharyzhalgay metamorphic complex of the south-western Siberian Craton were dated by the U\\Pb method on zircon at 1864.7 ± 1.8 Ma. This age postdates the peak of metamorphism (collision) by 15 million years, and together with other published ages for OIB-type mafic rocks, some Pt-bearing intrusions and postcollisional granites, marks the stage of post-collisional extension. The studied rocks can be split into low- and high-MgO lamprophyres. The low-MgO lamprophyres are formally classified as absarokite – the mafic counterpart of the shoshonitic series. It shows strong shoshonitic affinities with the Tibetan post-collisional volcanic rocks on various major and trace element diagrams. Some of the high-MgO lamprophyres are too low in potassium (K2O/Na2O b 1) to be considered within the shoshonitic series, but they show various geochemical signatures of the shoshonitic series as well. The diversity of the minerals in the high-MgO lamprophyres implies that the assemblage cannot have crystallized in equilibrium. The highMgO lamprophyres are a hybrid rock type, containing minerals which crystallized from a felsic melt and minerals of a shallow, orthopyroxenitic mantle wedge. The Sr, Nd and Pb isotopes show that the studied lamprophyres crystallized from melts derived from a crustal or strongly metasomatized lithospheric source. We suggest that the source was subducted to mantle depth during collision and melted as a result of post-collisional extension. Some melts derived from the subducted plate entrained mantle wedge minerals on the melt passage to the surface and attained the high-MgO composition, yet retaining recycled isotopic signatures. Overall, our study shows that in the Paleoproterozoic shoshonitic melts were emplaced within a similar tectonic setting as seen today in modern orogenic systems. Supplementary data to this article can be found online at https://doi. org/10.1016/j.lithos.2019.01.015.

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

Acknowledgements We thank Jean-Paul Liegeois and an anonymous reviewer for critical comments and Hafiz Ur Rehman for editorial handling. The study was conducted under the financial support of the Russian Science Foundation (project no. 16-17-10068) to A.V.I., E.I.D. and S.V.M. The study of LIV, LVI and PLA, on one hand, and V.I.P., on the other hand, was performed by the governmental assignments in terms of the tasks IX.129.1.3 (№ 0350-2016-0029) and № 0144-2014-0091, respectively. References Afonin, V.P., Gunicheva, T.N., Piskunova, L.F., 1984. X-ray fluorescence silicate analysis. Novosibirsk, Nauka. P. 228 (in Russian). Aftalion, M., Bibikova, E.V., Bowes, D.R., Hopgood, A.M., Perchuk, L.L., 1991. Timing of Early Proterozoic collisional and extensional events in the granulite-gneiss-charnokitegenite complex, Lake Baikal, USSR – a U-Pb, Rb-Sr, and Sm-Nd isotopic study. Journal of Geology 99, 851–861. https://doi.org/10.1086/629556. Bergman, S.C., 1987. Lamproites and other potassium-rich igneous rocks: a review of their occurrence, mineralogy and geochemistry. In: Fitton, J.G., Upton, B.G.J. (Eds.), Alkaline Igneous Rocks. vol. 30. Geological Society Special Publication, pp. 103–190. Bibikova, E.V., Khiltova, V.J., Gracheva, T.V., Makarov, V.A., 1982. Age of greenstone belts of the Prisayanie. Doklady Akademii Nauk SSSR 267, 1171–1174 in Russian. 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, 310–315. https://doi.org/10.1134/ S0016702906030098. von Blanckenburg, F., Davies, J.H., 1995. Slab breakoff: a model for syncollisional magmatism and tectonics in Alps. Tectonics 14, 120–131 10.1029|94TC02051. Brey, G.P., Bulatov, V., Girnis, A., Harris, J.W., Stachel, T., 2004. Ferropericlase – a lower mantle phase in the upper mantle. Lithos 77, 655–663. https://doi.org/10.1016/j. lithos.2004.03.013. Campbell, I.H., Stepanov, A.S., Liang, H.-Y., Allen, C.M., Norman, M.D., Zhang, Y.-Q., Xie, Y.W., 2014. The origin of shoshonites: new insights from the Tertiary high-potassium intrusions of eastern Tibet. Contributions to Mineralogy and Petrology 167. https:// doi.org/10.1007/s00410-014-0983-9 no. 983. Condie, K.C., 1981. Archean Greenstone Belts. Developments in Precambrian Geology 3, 1–434. Corfu, F., 2004. U–Pb age, setting and tectonic significance of the anorthosite-mangerite charnockite granite suite, Lofoten-Vesteralen. Norway Journal of Petrology 45, 1799–1819. https://doi.org/10.1093/petrology/egh034. Corfu, F., Svensen, H., and Mazzini, A., 2016. Comment to paper: Evaluating the temporal link between the Karoo LIP and climatic-biologic events of the Toarcian Stage with high precision U–Pb geochronology by Bryan Sell, Maria Ovtcharova, Jean Guex, Annachiara Bartolini, Fred Jourdan, Jorge E. Spangenberg, Jean-Claude Vicente, Urs Schaltegger in Earth and Planetary Science Letters 408 (2014) 48–56. Earth and Planetary Science Letters 434, 349–352, doi: https://doi.org/10.1016/j.epsl.2015.07.010. Davies, G.F., 1992. On the emergence of plate tectonics. Geology 20, 963–966. https://doi. org/10.1130/0091-7613(1992)020b0963:OTEOPTN2.3.CO;2. Davies, G.R., Stolz, A.J., Mahotkin, I.L., Nowell, G.M., Pearson, D.G., 2006. Trace element and Sr–Pb–Nd–Hf isotope evidence for ancient, fluid-dominated enrichment of the source of Aldan Shield lamproites. Journal of Petrology 47, 1119–1146. https://doi.org/ 10.1093/petrology/eg1005. Didenko, A.N., Kozakov, I.K., Bibikova, E.V., Vodovozov, V.Y., Khil'tova, V.Y., Reznitskii, L.Z., Ivanov, A.V., Levitskii, V.I., Travin, A.V., Shevchenko, D.O., Rasskazov, S.V., 2003. Paleoproterozoic granites of the Sharyzhalgai block, Siberian Craton: Paleomagnetism and geodynamic inferences. Doklady Earth Sciences 390, 510–515. 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 postcollision magmatism at the southern flank of the Siberian Craton: New geochronological data and geodynamic implications. Doklady Earth Sciences 383, 125–128. Dupuis, N.E., Murphy, J.B., Braid, J.A., Shail, R.K., Nance, R.D., 2016. Mantle evolution in the Variscides of SW England: Geochemical and isotopic constraints from mafic rocks. Tectonophysics 681, 353–363. https://doi.org/10.1016/j.tecto.2016.02.044. Ernst, R.E., Hamilton, M.A., Söderlund, U., Hanes, J.A., Gladkochub, D.P., Okrugin, A.V., Kolotilina, T., Mekhonoshin, A.S., Bleeker, W., LeCheminant, A.N., Buchan, K.L., Chamberlain, K.R., Didenko, A.N., 2016. Long-lived connection between southern Siberia and northern Laurentia in the Proterozoic. Nature Geoscience 9, 464–469. https://doi.org/10.1038/NGEO2700. Fabrikova, E.A., 1961. Flame-photometry determination of alkaline metals in silicate rocks and minerals. IMGRE AN USSR Transactions (6), 118 (In Russian). Foley, S.F., Venturelli, G., Green, D.H., Toscani, L., 1987. The ultrapotassic rocks: characteristics, classification and constraints for petrogenetic models. Earth-Science Reviews 24, 81–134. https://doi.org/10.1016/0012-8252(87)90001-8. Gale, A., Dalton, C.A., Langmuir, C.H., Su, Y., Schilling, J.-G., 2013. The mean composition of ocean ridge basalts. Geochemistry, Geophysics, Geosystems 14, 489–518. https://doi. org/10.1029/2012GC004334. Gladkochub, D.P., Pisarevsky, S.A., Donskaya, T.V., Natapov, L.M., Mazukabzov, A.M., Stanevich, A.M., Sklyarov, E.V., 2006. The Siberian Craton and its evolution in terms of the Rodinia hypothesis. Episodes 29, 169–174. Gladkochub, D.P., Pisarevsky, S.A., Donskaya, T.V., Ernst, R.E., Wingate, M.T.D., Söderlund, U., Mazukabzov, A.M., Sklyarov, E.V., Hamilton, M.A., Hanes, J.A., 2010. Proterozoic mafic magmatism in Siberian craton: an overview and implications for

99

paleocontinental reconstruction. Precambrian Research 183, 660–668. https://doi. org/10.1016/j.precamres.2010.02.023. Gladkochub, D.P., Pisarevsky, S.A., Mazukabzov, A.M., Söderlund, U., Sklyarov, E.V., Donskaya, T.V., Ernst, R.E., Stanevich, A.M., 2013. The first evidence of Paleoproterozoic late-collision basite magmatism in the near-Sayan salient of the Siberian craton basement. Doklady Earth Sciences 450, 583–586. https://doi.org/ 10.1134/S1028334X13060019. Glebovitskii, V.A., Levchenkov, O.A., Levitskii, V.I., Rizvanova, N.G., Levskii, L.K., Bogomolov, E.S., Levitskii, I.V., 2011. Age stages of metamorphism at the Kitoi sillimanate schist deposit, southeastern Prisayan'e. Doklady Earth Sciences 436, 13–17. https://doi.org/10.1134/S1028334X11010247. Hamilton, W.B., 2003. An Alternative Earth: GSA Today 13, 4–12. Hebert, R., Guilmette, C., Dostal, J., Bezard, R., Lesage, G., Bedard, E., Wang, C., 2014. Miocene post-collisional shoshonites and their crustal xenoliths, Yarlung Zangbo Suture Zone southern Tibet: Geodynamic implications. Gondwana Research 25, 1263–1271. https://doi.org/10.1016/jgr.2013.05.013. Hergt, J.M., Peate, D.W., Hawkesworth, C.J., 1991. The petrogenesis of Mesozoic Gondwana low-Ti flood basalts. Earth and Planetary Science Letters 105, 134–148. https://doi.org/10.1016/0012-821X(91)90126-3. van Hunen, J., Allen, M.B., 2011. Continental collision and slab break-off: a comparison of 3-D numerical models with observations. Earth and Planetary Science Letters 302, 27–37. https://doi.org/10.1016/j.epsl.2010.11.035. Ivanov, A.V., 2006. Systematic differences between U–Pb and 40Ar/39Ar dates: reasons and evaluation techniques. Geochemistry International 44, 1041–1047. Ivanov, A.V., Levitsky, I.V., Levitsky, V.I., Demonterova, E.I., Reznitsky, L.Z., 2017. Paleoarchean mafic rocks of the Southwestern Siberian craton: preliminary geochronology and geochemical characterization. Geodynamics & Tectonophysics 8, 477–480. https://doi.org/10.5800/GT-2017-8-3-0267. Ivanov, A.V., Demonterova, E.I., Savatenkov, V.M., Perepelov, A.B., Ryabov, V.V., Shevko, A.Y., 2018. Late Triassic (Carnian) lamproites from Noril'sk, polar Siberia: evidence for melting of the recycled Archean crust and the question of lamproite source for some placer diamond deposits of the Siberian Craton. Lithos 296-299, 67–78. https://doi.org/10.1016/j.lithos.2017.10.021. Kaminsky, F., 2012. Mineralogy of the lower mantle: a review of 'super-deep' mineral inclusions in diamond. Earth-Science Reviews 110, 127–147. https://doi.org/10.1016/j. earscirev.2011.10.005. Kay, R.W., Kay, S.M., 1993. Delamination and delamination magmatism. Tectonophysics 219, 177–189. https://doi.org/10.1016/0040-1951(93)90295-U. Keskin, M., 2003. Magma generation by slab steepening and breakoff beneath a subduction-accretion complex: an alternative model for collision-related volcanism in Eastern Anatolia, Turkey. Geophysical Research Letters 30, 8046. https://doi.org/ 10.1029/2003GL018019. Kirnozova, T.I., Bibikova, E.V., Kozakov, I.K., Kovach, V.P., Didenko, A.N., Reznitsky, L.Z., 2003. Early Proterozoic post-collisional granitoids of the Pri-Sayan uplift of the Siberian Platform: U-Pb geochronological and Sm-Nd isotope data. Abstract of the II Russian conference under isotopic geochronology. pp. 193–195 (in Russian). Kramers, J.D., Tolstikhin, I.N., 1997. Two terrestrial lead isotope paradoxes, forward transport modelling, core formation and the history of the continental crust. Chemical Geology 139, 75–110. https://doi.org/10.1016/S0009-2541(97)00027-2. Le Maitre, R.W., Streckeisen, A., Zanettin, B., Le Bas, M.J., Bonin, B., Bateman, P., Bellieni, G., Dudek, A., Efremova, S., Keller, J., Lameyre, J., Sabine, P.A., Schmid, R., Sorensen, H., Woolley, A.R., 2002. Igneous Rocks. A classification and glossary terms. Recommendations of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks, 2nd edition Cambridge University Press (236 p). Levchenkov, O.A., Levitskii, V.I., Rizvanova, N.G., Kovach, V.P., Sergeeva, N.A., Levskii, L.K., 2012. Age of the Irkut block of the Prisayan uplift of the Siberian platform basement: Dating minerals from metamorphic rocks. Petrology 20, 86–92. https://doi.org/ 10.1134/S0869591112010031. Levitskii, I.V., 2012. Geochemistry of Granulite and Greenschist Complexes of Prisayan Salient of the Siberian Platform Basement Candidate of Science Dissertation. 276 p. (In Russian). Levitskii, V.I., Mel'nikov, A.I., Reznitskii, L.Z., Bibikova, E.V., Kirnozova, T.I., Kozakov, I.K., Makarov, V.A., Plotkina, Y.V., 2002. Early Proterozoic postcollisional granitoids in southwestern Siberian craton. Geologiya i Geofizika 43, 717–731. Levitskii, V.I., Sal'nikova, E.B., Kotov, A.B., Reznitskii, L.Z., Barash, I.G., Yakovleva, S.Z., Kovach, V.P., Mel'nikov, A.I., Plotkina, Y.V., 2004. Age of formation of apocarbonate metasomatites of the sharyzhalgai uplift of the Siberian craton basement, southwestern Baikal region: U-P. Doklady Earth Sciences 399, 650–654. Levitskii, V.I., Reznitskii, L.Z., Sal'nikova, E.B., Levitskii, I.V., Kotov, A.B., Barash, I.G., Yakovleva, S.Z., Anisimova, I.V., Plotkina, Y.V., 2010. Age and origin of Kitoi sillimanite schist deposit, eastern Siberia. Doklady Earth Sciences 431, 394–398. https://doi.org/ 10.1134/S1028334X1003027X. Liegeois, J.P., Black, R., 1987. Alkaline magmatism subsequent to collision in the PanAfrican belt of the Adrar des Iforas (Mali). Geological Society of London, Special Publication 30, 381–401. https://doi.org/10.1144/GSL.SP.1987.030.01.18. Lustrino, M., Agostini, S., Chalal, Y., Fedele, L., Stagno, V., Colombi, F., Bouguerra, A., 2016. Exotic lamproites or normal ultrapotassic rocks? The Late Miocene volcanic rocks from Kef Hahouner, NE Algeria, in the frame of the circum-Mediterranean lamproites. Journal of Volcanology and Geothermal Research 327, 539–553. https://doi.org/ 10.1016/j.jvolgeores.2016.09.021. Manhes, G., Allègre, C.J., Provost, A., 1984. U-Th-Pb systematics of the eucrite “Juvinas”. Precise age determination and evidence for exotic lead. Geochimica et Cosmochimica Acta 48, 2247–2264. McDonough, W.F., Sun, S.-S., 1995. The composition of the earth. Chemical Geology 120, 223–253. https://doi.org/10.1016/0009-2541(94)00140-4.

100

A.V. Ivanov et al. / Lithos 328–329 (2019) 88–100

McLennan, S.M., 2001. Relationships between the trace element composition of sedimentary rocks and upper continental crust. Geochemistry, Geophysics, Geosystems 2. https://doi.org/10.1029/2000GC000109 2000GC000109. Mekhonoshin, A.S., Ernst, R., Söderlund, U., Hamilton, M.A., Kolotilina, T.B., Izokh, A.E., Polyakov, G.V., Tolstykh, N.D., 2016. Relationship between platinum-bearing ultramafic-mafic intrusions and large igneous provinces (exemplified by the Siberian Craton). Russian Geology and Geophysics 57, 822–833. https://doi.org/10.1016/j. rgg.2015.09.020. Menzies, M.A., 1989. Cratonic, circumcratonic and oceanic mantle domains beneath the Western United States. Journal of Geophysical Research 94, 7899–7915. https://doi. org/10.1029/JB094iB06p07899. Min, K., Mundil, R., Renne, P.R., Ludwig, K.R., 2000. A test for systematic errors in 40Ar/39Ar geochronology through comparison with U/Pb analysis of a 1.1-Ga rhyolite. Geochimica et Cosmochimica Acta 64, 73–98. https://doi.org/10.1016/S0016-7037 (99)00204-5. Murphy, D.T., Collerson, K.D., Kamber, B.S., 2002. Lamproites from Gaussberg, Antarctica: possible transition zone melts of Archean subducted sediments. Journal of Petrology 43, 981–1001. https://doi.org/10.1093/petrology/43.6.981. Nebel, O., Vroon, P.Z., van Westrenen, W., Iizuka, T., Davies, G.R., 2011. The effect of sediment recycling in subduction zones on the Hf isotope character of new arc crust, Banda arc, Indonesia. Earth and Planetary Science Letters 303, 240–250. https://doi. org/10.1016/j.epsl.2010.12.053. Panteeva, S.V., Gladkochub, D.P., Donskaya, T.V., Markova, V.V., Sandimirova, G.P., 2003. Determination of 24 trace elements in felsic rocks by inductively coupled plasma mass spectrometry after lithium metaborate fusion. Spectrochimica Acta, Atomic Spectroscopy 58, 341–350. https://doi.org/10.1016/S0584-8547(02)00151-9. Pearce, J.A., 1976. Statistical analysis of major element patterns in basalts. Journal of Petrology 17, 15–43. https://doi.org/10.1093/petrology/17.1.15. Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: Thorpe, R.S. (Ed.), Andesites. Wiley, NY, pp. 525–548. Pearce, J.A., 2008. Geochemical fingerprinting of oceanic basalts with applications to ophiolite classification and the search for Archean oceanic crust. Lithos 100, 14–48. https://doi.org/10.1016/j.lithos.2007.06.016. Pearce, J.A., Norry, M.J., 1979. Petrogenic implications of Ti, Zr, Y and Nb variations in volcanic rocks. Contributions to Mineralogy and Petrology 69, 33–47. https://doi.org/ 10.1007/BF00375192. Pearce, J.A., Harris, N.B.W., Tindle, A.G., 1984. Trace-element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology 25, 956–983. https://doi.org/10.1093/petrology/25.4.956. Peccerillo, A., Taylor, S.R., 1976. Geochemistry of Eocene calc-alkaline volcanic rocks from the Kastomonon area, northern Turkey. Contributions to Mineralogy and Petrology 58, 63–81. Peretyazhko, I.S., Savina, E.A., Khromova, E.A., Karmanov, N.S., Ivanov, A.V., 2018. Unique Clinkers and Paralavas from a New Nyalga Combustion Metamorphic Complex in Central Mongolia: Mineralogy, Geochemistry, and Genesis. Petrology 26, 181–211. https://doi.org/10.1134/S0869591118020054. Plank, T., Langmuir, C.H., 1998. The chemical composition of subducting sediment and its consequences for the crust and mantle. Chemical Geology 145, 325–394. https://doi. org/10.1016/S0009-2541(97)00150-2. Poller, U., Gladkochub, D.P., Donskaya, T.V., Mazukabzov, A.M., Sklyarov, E.V., Todt, W., 2004. Timing of early Proterozoic magmatism along the Southern margin of the Siberian Craton (Kitoy area). Transactions of the Royal Society of Edinburgh: Earth Sciences 95, 215–225. Poller, U., Gladkochub, D., Donskaya, T., Mazukabzov, A., Sklyarov, E., 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. https://doi.org/10.1016/j.precamres.2004.12.003. Rojas-Agramonte, Y., Kröner, A., Demoux, A., Xia, X., Wang, W., Donskaya, T., Liu, D., Sun, M., 2011. Detrital and xenocrystic zircon ages from Neoproterozoic to Palaeozoic arc terranes of Mongolia: significance for the origin of crustal fragments in the Central Asian Orogenic Belt. Gondwana Research 19, 751–763. https://doi.org/10.1016/j. gr.2010.10.004. 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. https://doi.org/10.1134/ S0869593807040016. Sandimirova, G.P., Livitsky, V.I., Pakholchenko, Y.A., 1992. Rb-Sr geochronology of the oldest granitoids of the East Sayan. Doklady Akademii Nauk 326, 696–700. Sheth, H.R., 2008. Do major oxide tectonic discrimination diagrams work? Evaluating new log-ratio and discriminant-analysis-based diagrams with Indian Ocean mafic volcanic and Asian ophiolites. Terra Nova 20, 229–236. https://doi.org/10.1111/j.13653121.2008.00811.x. Shokhonova, M.N., Donskaya, T.V., Gladkochub, D.P., Mazukabzov, A.M., Paderin, I.P., 2010. Paleoproterozoic basaltoids in the North Baikal volcanoplutonic belt of the Siberian craton: age and petrogenesis. Russian Geology and Geophysics 51, 815–832. https://doi.org/10.1016/j.rgg.2010.07.001.

Stacey, J.S., Kramers, J.D., 1975. Approximation of terrestrial lead isotope evolution by a 2stage model. Earth and Planetary Science Letters 26, 207–221. https://doi.org/ 10.1016/0012-821X(75)90088-6. Stachel, T., Harris, J.W., Brey, G.P., Joswig, W., 2000. Kankan diamonds (Guinea) II: lower mantle inclusions parageneses. Contributions to Mineralogy and Petrology 140, 16–27. https://doi.org/10.1007/s004100000174. Steiger, R.H., Jäger, E., 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters 36, 359–362. Stern, R.J., 2005. Evidence from ophiolites, blueschists, and ultra-high pressure metamorphic terranes that the modern episode of subduction tectonics began in Neoproterozoic time. Geology 33, 557–560. https://doi.org/10.1130/G21365.1. Stern, R.J., 2007. When and how did plate tectonic begin? Theoretical and empirical considerations. Chinese Science Bulletin 52, 578–591. https://doi.org/10.1007/s11434007-0073-8. Stern, C.R., Kilian, R., 1996. Role of the subducted slab, mantle wedge and continental crust in the generation of adakites from the Andean Austral Volcanic Zone. Contributions to Mineralogy and Petrology 123, 263–281. Turkina, O.M., Berezhnaya, N.G., Urmantseva, L.N., Paderin, I.P., Skublov, S.G., 2009. U-Pb isotope and REE composition of zircon from the pyroxene crystalline schists of the Irkut terrane, Sharyzhalgai Uplift: evidence for the Neoarchean magmatic and metamorphic events. Doklady Earth Sciences 429, 527–533. https://doi.org/10.1134/ S1028334X09090207. 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. https://doi.org/ 10.1016/j.gr.2011.09.012. 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. https://doi.org/10.1016/j.rgg.2013.02.003. Turner, S., Arnaud, N., Liu, J., Rogers, N., Hawkesworth, C., Harris, N., Kelley, S., Van Calsteren, P., Deng, W., 1996. Post-collision, shoshonitic volcanism on the Tibetian Plateau: Implications for convective thinning of the lithosphere and the source of ocean island basalts. Journal of Petrology 37, 45–71. Turner, S., Rushmer, T., Reagan, M., Moyen, J.-F., 2014. Heading down early on? Start of subduction on Earth. Geology 42, 139–142. https://doi.org/10.1130/G34886.1. Urmantseva, L.N., Turkina, O.M., Kapitonov, I.N., 2012. Protoliths of Paleoproterozoic calciphyres from the Irkut block (Sharyzhalgai uplift of the Siberian craton): composition and origin. Russian Geology and Geophysics 53, 1291–1303. https://doi.org/ 10.1016/j.rgg.2012.10.003. Vasil'eva, I.E., Kuznetsov, A.M., Vasil'ev, I.L., Shabanova, E.L., 1997. Calibration of procedures of atomic emission spectroscopy with computer processing of spectra. Journal of Analytical Chemistry 52, 1118–1127. Verma, S.P., Agrawal, S., 2011. New tectonic discrimination diagrams for basic and ultrabasic volcanic rocks through log-transformed ratios of high field strength elements and implications for petrogenetic processes. Revista Mexicana de Ciencias Geologicas 28, 24–44. Verma, S.P., Guevara, M., Agrawal, S., 2006. Discriminating four tectonic settings: five new geochemical diagrams for basic and ultrabasic volcanic rocks based on log-ratio transformation of major element data. Journal of Earth System Science 115, 485–528. https://doi.org/10.1007/BF02702907. Vermeesch, P., 2006. Tectonic discrimination diagrams revisited. Geochemistry, Geophysics, Geosystems 7, Q06017. https://doi.org/10.1029/2005GC001092. Vermeesch, P., 2012. On the visualisation of detrital age distributions. Chemical Geology 312-313, 190–194. https://doi.org/10.1016/j.chemgeo.2012.04.021. Williams, H.M., Turner, S.P., Pearce, J.A., Kelley, S.P., Harris, N.B.W., 2004. Nature of the source regions for post-collisional, potassic magmatism in southern and northern Tibet from geochemical variations and inverse trace element modelling. Journal of Petrology 45, 555–607. https://doi.org/10.1093/petrology/cgg094. de Wit, M.J., 1998. On Archean granites, greenstones, cratons and tectonics: does the evidence demand a verdict? Precambrian Research 91, 181–226. https://doi.org/ 10.1016/S0301-9268(98)00043-6. Xu, Y.G., Menzies, M.A., Thirlwall, M.F., Xie, G.H., 2001. Exotic lithosphere mantle beneath the western Yangtze craton: petrogenetic links to Tibet using highly magnesian ultrapotassic rocks. Geology 29, 863–866. https://doi.org/10.1130/0091-7613(2001) 029b0863:ELMBTWN2.0.CO;2. Zedgenizov, D.A., Yefimova, E.S., Logvinova, A.M., Shatsky, V.S., Sobolev, N.V., 2001. Ferropericlase inclusions in a diamond microcrystal from the Udachnaya kimberlite pipe, Yakutia. Doklady Earth Sciences 377, 319–321. Zhu, D.-C., Wang, Q., Zhao, Z.-D., Chung, S.-L., Cawood, P.A., Niu, Y., Liu, S.-A., Wu, F.-Y., Mo, X.-X., 2015. Magmatic record of India-Asia collision. Scientific Reports 5, 14289. https://doi.org/10.1038/srep14289.