Scholarly article on topic 'Mid-Miocene thermal impact on the lithosphere by sub-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia'

Mid-Miocene thermal impact on the lithosphere by sub-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia Academic research paper on "Earth and related environmental sciences"

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Geoscience Frontiers
{Volcanism / Geodynamics / Cenozoic / Asia / Asthenosphere / Lithosphere}

Abstract of research paper on Earth and related environmental sciences, author of scientific article — Irina Chuvashova, Sergei Rasskazov, Tatyana Yasnygina

Abstract High-Mg lavas are characteristic of the mid-Miocene volcanism in Inner Asia. In the Vitim Plateau, small volume high-Mg volcanics erupted at 16–14 Ma, and were followed with voluminous moderate-Mg lavas at 14–13 Ma. In the former unit, we have recorded a sequence of (1) initial basaltic melts, contaminated by crustal material, (2) uncontaminated high-Mg basanites and basalts of transitional (K–Na–K) compositions, and (3) picrobasalts and basalts of K series; in the latter unit a sequence of (1) initial basalts and basaltic andesites of transitional (Na–K–Na) compositions and (2) basalts and trachybasalts of K–Na series. From pressure estimation, we infer that the high-Mg melts were derived from the sub-lithospheric mantle as deep as 150 km, unlike the moderate-Mg melts that were produced at the shallow mantle. The 14–13 Ma rock sequence shows that initial melts equilibrated in a garnet-free mantle source with subsequently reduced degree of melting garnet-bearing material. No melting of relatively depleted lithospheric material, evidenced by mantle xenoliths, was involved in melting, however. We suggest that the studied transition from high- to moderate-Mg magmatism was due to the mid-Miocene thermal impact on the lithosphere by hot sub-lithospheric mantle material from the Transbaikalian low-velocity (melting) domain that had a potential temperature as high as 1510 °С. This thermal impact triggered rifting in the lithosphere of the Baikal Rift Zone.

Academic research paper on topic "Mid-Miocene thermal impact on the lithosphere by sub-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia"


Accepted Manuscript

Mid-Miocene thermal impact on the lithosphere by sub-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia

Irina Chuvashova, Sergei Rasskazov, Tatyana Yasnygina

PII: S1674-9871(16)30058-5

DOI: 10.1016/j.gsf.2016.05.011

Reference: GSF 464

To appear in: Geoscience Frontiers

Received Date: 7 July 2015 Revised Date: 10 May 2016 Accepted Date: 21 May 2016

Please cite this article as: Chuvashova, I., Rasskazov, S., Yasnygina, T., Mid-Miocene thermal impact on the lithosphere by sub-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia, Geoscience Frontiers (2016), doi: 10.1016/ j.gsf.2016.05.011.

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Initiation of rifting

16-14 Ma

High-Mg magmatism

i Contamination of /S'IvÎV mantle liquids by crustal material

Grt— Grt+

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Partial arrest — of liquids in shallow mantle

Solid state diffusion in the lithospheric mantle

Grt = 6-12%

Advanced rifting

14-13 Ma

Moderate-Mg magmatism

Melting at shallow /' F-0.15-0.20 mantle source /

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Integrated thermal effect in the shallow lithosphere

Motion of the lithosphere, rate 2-3 cm yr1

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b-lithospheric convective mantle material: Transition from high- to moderate-Mg magmatism beneath Vitim Plateau, Siberia

Irina Chuvashovaa,b,\ Sergei Rasskazova,b, Tatyana Yasnyginaa

a Institute of the Earth's Crust, Siberian Branch of RAS, Irkutsk, Russia bIrkutsk State University, Irkutsk, Russia

* Corresponding author e-mail address:


High-Mg lavas are characteristic of the mid-Miocene volcanism in Inner Asia. In the Vitim Plateau, small volume high-Mg volcanics erupted at 16-14 Ma, and were followed with voluminous moderate-Mg lavas at 14-13 Ma. In the former unit, we have recorded a sequence of (1) initial basaltic melts, contaminated by crustal material, (2) uncontaminated high-Mg basanites and basalts of transitional (K-Na-K) compositions, and (3) picrobasalts and basalts of K series; in the latter unit a sequence of (1) initial basalts and basaltic andesites of transitional (Na-K-Na) compositions and (2) basalts and trachybasalts of K-Na series. From pressure estimation, we infer that the high-Mg melts were derived from the sub-lithospheric mantle as deep as 150 km, unlike the moderate-Mg melts that were produced at the shallow mantle. The 14-13 Ma rock sequence shows that initial melts equilibrated in a garnet-free mantle source with subsequently reduced degree of melting garnet-bearing material. No melting of relatively depleted lithospheric material, evidenced by mantle xenoliths, was involved in melting, however. We suggest that the studied transition from high- to moderate-Mg magmatism was due to the Mid-Miocene thermal impact on the lithosphere by hot sub-lithospheric mantle material from the Transbaikalian low-velocity (melting) domain that had a potential temperature as high as 1510°С. This thermal impact triggered rifting in the lithosphere of the Baikal Rift Zone.

Key words: Volcanism; Geodynamics, Cenozoic; Asia; Asthenosphere; Lithosphere; Picrobasalt; Basanite; Basalt; Trace elements; Isotopes

1. Introduction

The explanation for the origin of within-plate volcanism from deep mantle sources is currently based on global seismic tomography images in combination with results of ultrahigh pressure experiments (Anderson, 2007; Ghosh et al., 2007, 2009; Maruyama et al., 2007; Karato, 2012). The emphasis in the ongoing debate about hot mantle plumes rising from the core-mantle boundary is placed on the crucial value of seismic velocity structure of the mantle (Hofmann, 1997; Foulger, 2010). The structure of the mantle and components of volcanic rocks associated with the Hawaiian plume are adequately interpreted by high-temperature processes responsible for generating high-Mg magmatic liquids (Sobolev et al., 2005, 2011; Wang and Gaetani, 2008; Zhao, 2009; Hanano et al., 2010; Herzberg, 2011). The origin of basalts from continents and other ocean islands, where no high-Mg liquids erupted, is attributed to extension of the lithosphere and precursor metasomatism without high temperature processes (Griffin et al., 1988; Barry et al., 2003).

The Emeishan and Siberian large igneous provinces (LIPs) that formed in Asia in the late Permian and at the Permian-Triassic boundary were dominated by moderate-Mg (moderate-temperature) basaltic melts with minor high-Mg (high-temperature) meimechites and, probably, komatiites (Vasiliev and Zolotukhin, 1995; Fedorenko et al., 1996; Fedorenko and Czamanske, 1997; Kogarko and Ryabchikov, 2000; Hanski et al., 2004; Carlson et al., 2006; Zhang et al.,



simulated by a thermochemical plume in the upper mantle and transition zone with a diameter of nearly 1000 km (Sobolev et al., 2011). However, no low-velocity region that might correspond to the Siberian LIP is recorded (Koulakov and Bushenkova, 2010).

Cenozoic volcanic fields of Asia are often underlain by low-velocity anomalies residing in the upper mantle (Rasskazov et al., 2003a; Lei and Zhao, 2005; Zhao et al., 2009; Chuvasova et al., 2012; Wei et al., 2012). Decreasing velocities were firstly detected therein at a shallow level of the upper mantle through a temporal delay of seismic waves from nuclear explosions in Nevada, USA (Rogozhina and Kozhevnikov, 1979). At the beginning of the 1990s, some volcanic areas in Central and East Asia were referred to deep mantle plumes (Nakamura et al., 1990; Yarmolyuk et al., 1990, 1994; Rasskazov, 1991, 1994). There was also an attempt to locate mantle plumes through geophysical data (Zorin et al., 2003, 2006). It was suggested that the dispersed within-plate Cenozoic volcanism of Asia was derived due to a broad asthenospheric upwelling from a deeper mantle level ("hot region") (Tatsumi et al., 1990) or convective flow upraised from the core-mantle boundary ("hot mantle field" (Zonenshain et al., 1991). The extensive asthenospheric upwelling was assumed to be marked by the DUPAL isotopic anomaly (Flower et al., 1998). However, from the measured Pb-isotope ratios in basalts, no anomaly was detected in Cenozoic basalts erupted within the Caledonian terranes of Inner Asia, although the anomaly was traced in lavas erupted through the Tuva-Mongolian and other Riphean micro-continents that separated from East Gondwana and drifted across Paleo-Asian Ocean from the Southern to Northern Hemisphere in the late Precambrian. It was inferred that the anomaly was brought by these tectonic units and could not belong to the deep "hot mantle region" (Rasskazov et al., 2002).

The S-wave model of Asia by Yanovskaya and Kozhevnikov (2003) was used for interpreting spatial-temporal distribution of Cenozoic basalts in terms of large low-velocity (melting) domains of two mantle levels: the Transbaikalian of 200-410 km and the Sayan-Mongolian, Sea of Okhotsk, and Philippine Sea of 50-200 km (Rasskazov et al., 2003a, 2004). The identities of the Sayan-Mongolian and Transbaikalian melting domains were highlighted by both a drastic lateral change of S-wave velocities at the depth of 250 km and high- to moderate-Mg transition of lava compositions (Fig. 1).

Most of the high-Mg low-volume lavas with 12-32 wt.% MgO (Le Bas, 2000; Kerr and Arndt, 2001) erupted in the outline of the Transbaikalian domain in the short time interval of 1614 Ma. Voluminous moderate-Mg lavas flooded large areas afterwards and lasted for millions of years. In the Kamar volcanic field, a single high-Mg lava flow was found within the 16-15 Ma unit that was preceded and followed by moderate-Mg basalts dated at 18.1-17.6 and 13-12 Ma, respectively (Rasskazov et al., 2003b, 2013b). In the Udokan volcanic field, the initial low-volume (few km3) eruptions of high-Mg olivine melaleucitites occurred at ca. 14 Ma, whereas the subsequent high-volume (450 km3) alkali olivine basalt-trachyte and basanites-tephriphonolite volcanic series flooded the area of 3 000 km in the last 8 Ma (Rasskazov et al., 1997, 2000). Small outcrops of high-Mg rocks and extensive moderate-Mg volcanic strata were identified also in the Vitim and Dariganga volcanic fields (Aschepkov, 1991; Aschepkov et al., 2003; Chuvashova et al., 2012, 2016). In this paper, we present results of a comparative study of Mid-Miocene rocks that erupted in the Vitim volcanic field at 16-14 and 14-13 Ma and exhibited high- and moderate-Mg liquids, respectively.

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2. Geologic background and sampling

The Vitim Plateau is situated within the Baikal-Vitim Fold System, the geological structure of which has long been the object of lively debate. Comprehensive geological mapping of the area showed sedimentary and volcano-sedimentary units of late Precambrian and early Paleozoic designated as Baikalides and early Caledonides, respectively (Belichenko, 1977). Later on, the whole area of Western Transbaikal was referred to a single Barguzin super-terrane, much of c,oss-i,nked by .he Ma granites of the huge Angara-Vitim batholith

(Melnikov et al., 1994). Due to thorough geological study, the area was subdivided, nevertheless, into early Paleozoic terranes of different origin (Belichenko et al., 2006). According to this interpretation, the Vitim Plateau is located between the Ikat back-arc turbidite and West Stanovoy composite terranes within the Eravna island-arc tectonic unit that was affected by volcanism in the Mesozoic and Cenozoic (Fig. 2). Stratigraphic and paleontological records in small fragments disseminated in the Angara-Vitim batholith showed also Upper Silurian-Carboniferous sedimentary rocks (Ruzhentsev et al., 2012) that could originate on a continental margin of the closing Mongol-Okhotsk Bay of the Paleo-Pacific (Sengor and Natal'in, 1996; Parphenov et al., 1999).

The Cenozoic volcanism of the Vitim Plateau has become known due to numerous publications on deep-seated inclusions from basanites and picrobasalts in this area (Volyanyuk et al., 1976; Kiselev et al., 1979; Rasskazov, 1985, 1993; Aschepkov, 1991; Ionov et al., 1993, 2005; Litasov et al., 2000; Aschepkov, Andre, 2002; Litasov and Taniguchi, 2002; Aschepkov et al., 2011; Goncharov and Ionov, 2012). Cenozoic volcanic and volcano-sedimentary sequences of the Vitim Plateau have been studied since the beginning of the 1980s using borehole cores for K-Ar dating of volcanic layers and paleontological records of sediments (Rasskazov and Batyrmurzaev, 1985; Rasskazov et al., 2000, 2001, 2007; Chernyaeva et al., 2007; Usoltseva et al., 2010). The sequences were assigned as the Kularikta, Dzhilinda, Khoygot, and Bereya Formations dated back to the Oligocene, middle-upper Miocene, Pliocene, and Pleistocene, respectively.

The Kularikta Formation accumulated in conditions of slightly dissected relief. No reliable geological and geochronological evidence has so far been obtained on the Oligocene volcanic activity in the area. The Dzhilinda Formation was deposited within erosional paleovalleys as deep as 500 m that were filled firstly with coarse sediments, containing few volcanic layers, and then with lavas and lacustrine sediments. From borehole records of the Dzhilinda Formation, the major Northern, Southern, and Central paleovalleys have been traced over 120, 100, and 50 km, respectively. After partial erosion of lavas and sediments of the Dzhilinda Formation, the new erosional channels were buried beneath lavas and sediments of the Khoygot, and Bereya Formations.

The Vitim volcanic field occupies an area of about 5000 km . The total volume of volcanic rocks exceeds 1500 km3. Slag-lava edifices, eighty eight of which erected on the surface of the volcanic plane and another eighteen edifices buried beneath lavas, belong to seven large and six small volcanic centers (Fig. 2a). Each large volcano with diameter of 16-20 km consists of between seven to seventeen slag-lava edifices. A volcano of smaller size comprises two to four slag-lava cones. A comparative study of volcanic sequences in the eastern and western parts of the Vitim volcanic field demonstrated similar lava compositions in the units formed between 13.0-9.5 Ma, significant difference between the earlier and later rocks (Chuvashova et al., 2016). In the Bereya volcanic center, located in the eastern part of the volcanic field, the early volcanic interval is exhibited by the 16-14 and 14-13 Ma volcano-sedimentary units. The volcanic center is marked by seven volcanic cones that overtop a lava shield in an area of at least 250 km . The representative data on rock compositions of the volcanic center were obtained from both outcrops and borehole cores (Figs. 3 and 4).

The 16-14 Ma high-Mg rocks occur as fragments, pillow lavas, and lenses in the layered tuffaceous-sedimentary strata, exposed in a quarry at the 76th km of the Romanovka to Bagdarin road (Aschepkov, 1991). This sequence was formed when the relief has not yet been dissected. At present the tuffaceous-sedimentary strata occupies a relatively high altitude (>1000 m). Fragments of volcanic rocks and pillow lavas are composed primarily of volcanic glass. Some rocks are aphyric; others contain large olivine crystals or fragments of disintegrated mantle and crustal rocks.

The volcanic units were sampled from the tuffaceous-sedimentary sequence in the open quarry as it was deepening. The top layer was opened in September of 1986, the middle one in

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found in the upper layer, high-Mg basanites and basalts as well as pillow lavas of the same rocks - in the middle layer, and high-Mg basalts with inclusions of crustal material - in the lower one (Fig. 5).

The erosional paleovalleys, deepened into the basement to the altitude of 850 m, were filled

with moderate-Mg lavas erupted between 14-13 Ma. A volume of the accumulated lava unit as

thick as 150 m exceeded 15 km . About 70% of the stratum exhibited the group of basalts and basaltic andesites throughout the volcanic center. The remaining 30% of the 14-13 Ma unit are basalts and trachybasalts that occur above the basalt-basaltic andesite package in the northern part of the volcanic center. This unit was sampled in both boreholes and outcrops.

3. Sample preparation and analytical procedures

Collected samples were prepared for analytical work by crushing in an agate mortar about 400-500 g of rock fragments. From high-Mg rocks, xenocrystic material was removed under a binocular microscope. The effect of this removal was tested by analyzing aliquots with and without xenocrysts for 5 samples selected from the lower layer of the 16-14 Ma unit (Fig. 5).

Analytical work was performed in the laboratory for isotopic and geochronological studies of the Institute of the Earth's Crust according to analytical procedures described by Rasskazov et al. (2012) and Yasnygina et al. (2015). Trace elements were determined by inductively coupled plasma mass spectrometer (ICP-MS) technique using an Agilent 7500 ce. Samples were prepared by microwave acid digestion with a mixture of HNO3 and HF. For a more complete decomposition, each sample was re-evaporated with HF and with additions H2O2. In addition, samples of high-Mg rocks were fused with lithium metaborate. Before measurement, In and Bi internal standards were introduced into the sample. Internal standard correction was done by interpolation. Standard reference solutions of HPS (High-Purity Standards) production and geological reference materials of the USGS (BIR-1a, DNC-1a, BHVO-2, AGV-2, and RGM-1) were used for calibration and validation of analyses. Measurement of Sr and Nd isotope ratios were performed on a Finnigan MAT 262 thermal ionization mass spectrometer (TIMS). To control the quality of measurements, standard reference samples JNd-1 (Japan) and NBS SRM-987 (USA) were used. The obtained ratio 143Nd/144Nd = 0.512104 ± 4 (2o) was close to the standard value 0.512103 (JNd-1), the ratio 87Sr/86Sr = 0.710248 ± 0.000013 (2o) was comparable to the standard value 0.71025 (NbS SRM-987) (Faure, 2001). Major oxides were measured by classic "wet chemistry". Data on volcanic rocks were included into the database (Rasskazov et al., 2013b). Additional data on volcanic rocks from the studied units were used after Aschepkov (1991), Harris (1998), Litasov and Taniguchi (2002), Johnson et al. (2005).

4. Major oxides

The volcanic rocks from the mid-Miocene units of the Bereya volcanic center are well grouped on the classification diagrams K2O vs. K2O/Na2O, K2O vs. SiO2, (Na2O+K2O) vs. SiO2, and MgO vs. SiO2 (Fig. 6). Representative compositions of the groups are placed in Table 1.

The K-Na ratio is usually used as the main criterion for definition of potassic (K2O>Na2O), sodic (K2O<0.25 xNa2O), and potassic-sodic (Na2O>K2O>0.25xNa2O) series (in wt.%) (Foley et al., 1987; Rollinson, 1993). In the diagram of K2O vs. K2O/Na2O, the mid-Miocene rocks from the Bereya volcanic center overlap a field of K-Na series and partly occupy potassic and sodic fields. Among the 16-14 Ma rocks, the high-Mg groups of picrobasalts-basalts and contaminated basalts-trachybasalts belong to K and K-Na series, respectively. The group of high-Mg basanites-basalts shows intermediate position, i.e. transition from K to K-Na series. Unlike the Ugii-Nur volcanic field of Central Mongolia, where potassic affinity was provided by increasing K concentrations in moderate-Mg lavas, the potassic high-Mg series of the Bereya


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group of basalts-trachybasalts exhibits K-Na series, the group of basalts-basaltic andesites demonstrates a transition from Na to K-Na series.

Similarly, rocks of 16-14 and 14-13 Ma are well separated on the total alkalis vs. silica (TAS) diagram. High-Mg basanites-basalts and picrobasalts-basalts show low SiO2 contents (42.8-45.0 and 44.2-45.6 wt.%), whereas high-Mg basalts-trachybasalts demonstrate the elevated SiO2 concentrations (46.2-50.2 wt.%). The most silica-rich sample VT-13-09 (SiO2 57.6 wt.%) falls into the trachyandesite field. We assign this rock to a trachyandesite-like composition. While compared to bulk compositions, groundmasses separated from the high-Mg basalts-trachybasalts reveal systematically lower SiO2 abundances in the range of 45.9-49.6 wt.%. Lavas of 14-13 Ma have the whole range of SiO2 from 47.3 to 53.5 wt.%. K-Na basalts-trachybasalts differ from the transitional (Na-K-Na) basalts-basaltic andesites by lower SiO2 contents.

MgO contents in high-Mg basanites-basalts and picrobasalts-basalts range from 12.5 to 20.4 wt.%. With increasing SiO2 contents in the bulk compositions of the high-Mg basalts-trachybasalts, MgO contents decrease slightly relative to high-Mg basanites-basalts and picrobasalts-basalts, ranging from 10.7 to 16.8 wt.%. The groundmass separates from the high-Mg basalts-trachybasalts contain 10.3-12.6 wt.% MgO. The 14-13 Ma lavas display a dramatic decrease of MgO contents to 6-10 wt.%.

The rock groups of 16-14 and 14-13 Ma are clearly seen in diagrams of normative minerals (Fig. 7). High-Mg basalts-trachybasalts differ from high-Mg basanites-basalts and picrobasalts-basalts by lower contents of normative anorthite 700xAn/(An+Ab). High-Mg basanites-basalts show 3-9% Ne, whereas picrobasalts-basalts 0-8% Hy. In terms of these petrochemical characteristics, picrobasalts-basalts are similar to absarokites from Southwest Japan (Tatsumi and Koyaguchi, 1989). The basalt-trachybasalt, Na basalt, and basalt-basaltic andesite groups yield sub-parallel trends of slightly increasing normative anorthite consistent with increasing 700*An/(An+Ab) and a general transition from Ae-normative to Atz-Hy-normative compositions. Simultaneous reduction of MgO concentrations and normative anorthite from 1614 to 14-13 Ma rocks demonstrates decreasing crystallization temperatures of magmatic liquids.

5. Trace elements

The diagrams of rare earth elements (REE) show distinct separation of rocks from the 1614 and 14-13 Ma units (Fig. 8). Light REE (LREE) decrease from the high-Mg basanite VT-13-14 and picrobasalt VT-13-19 to moderate-Mg rocks, whereas heavy REE (HREE) increase accordingly. As a result, REE patterns for high- and moderate-Mg rocks intersect each other. Chondrite-normalized La/Yb ratios range from 22 to 38 in the groups of high-Mg basanites-basalts and picrobasalts-basalts and from 5 to 18 in the groups of moderate-Mg basalts. The trachyandesite-like rock VT-13-9 is comparable to the high-Mg basanite VT-13-14 in terms of LREE and to moderate-Mg lavas in terms of HREE. Four samples of the high-Mg basalts-trachybasalts demonstrate elevated concentrations of REE in groundmass relative to bulk compositions (not shown in diagrams).

The rock grouping is supported by a multi-element diagram. Compared to moderate-Mg lavas, the high-Mg basanite VT-13-14 and picrobasalt VT-13-19 show elevated concentrations of incompatible elements from Ba to Sr, although peaks and troughs of these two samples are consistent with those in moderate-Mg lavas. Basalts of Na series reveal specific minimum of K relative to U and Nb. The trachyandesite-like rock VT-13-9 indicates an effect of crustal contamination by elevated Cs, Rb, K, Pb and reduced Nb, Ta, Sr, P, Zr, Hf, and Ti abundances.

6. Discussion

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moderate-Mg liquids, we associate the mid-Miocene appearance of high-Mg magmas with high temperatures in the Transbaikalian melting domain, define geochemical signatures of melts erupted from deep and shallow mantle sources and, finally, present a model of a thermal impact of a hot sub-lithospheric material on the lithosphere beneath the Vitim Plateau.

6.1 Crustal contamination

High-Mg volcanic rocks of the initial volcanic episode 16-14 Ma are sufficiently contaminated by crustal material. In diagrams, data points of high-Mg contaminated basalts-trachybasalts and the trachyandesite-like rock VT-13-9 are offset from the trends of high-Mg rocks. The contamination is expressed in increasing SiO2 and decreasing 100*An/(An+Ab) to values typical of moderate-Mg rocks (Fig. 6).

In the diagram Th/Yb vs. Ta/Yb, high-Mg contaminated basalts-trachybasalts and trachyandesite-like rock VT-13-9 are shifted above the mantle array. This shift is a characteristic of data points not only for bulk compositions of rocks, but also for groundmasses (Fig. 9a). Thus, crustal material was partially melted and introduced into crystallized liquids. Bulk composition of the trachyandesite-like rock VT-13-9 reveals lower Ta/Yb, while its groundmass is comparable with those of other contaminated species. This sample is the most contaminated by crustal material.

Crustal contamination of high-Mg basalts-trachybasalts and the trachyandesite-like rock VT-13-9 is reflected in high initial 87Sr/86Sr (Fig. 9b). In the diagram initial 87Sr/86Sr, vs. 1000/Sr, these rocks demonstrate a typical effect of crustal contamination (i.e. increasing both

Sr/ Sr and 1000/Sr) relative to uncontaminated 16-14 Ma rocks. Various groups of uncontaminated rocks are shifted along the ordinate with increasing 1000/Sr from picrobasalts-basalts through high-Mg basanites-basalts and moderate-Mg basalts-trachybasalts to the basaltic trachyandesite V-9. The shift is resulted from the generally increasing degree of partial melting in the mantle sources. Meanwhile, Sr isotope ratios may be overstated due to rock alteration. Although this effect can be reduced by acid treatment of a sample (Song et al., 1990; Rasskazov et al., 2015), no criterion for a strontium isotope ratio in an unaltered rock is substantiated and no true value of this parameter can be estimated.

6.2 Trends of high- and moderate-Mg rocks

On plots of compatible (Ni) and incompatible (Th, La, K2O, Rb, Sr, Ba) trace elements versus MgO, high- and moderate-Mg rocks yield individual groups (Fig. 9a-g). On the Ni vs. MgO plot, data points of bulk compositions for contaminated rocks show elevated Ni and MgO contents, as compared with groundmass compositions, except strongly contaminated trachyandesites-like rock VT-13-9. Basalts-trachybasalts from the lower 16-14 Ma unit resulted from partial assimilation of crust material accompanied with fractional crystallization of olivine. From an assimilation-fractional-crystallization (AFC) model (DePaolo, 1981) on the Ni vs. MgO plot (Fig. 10a), groundmasses of these rocks underwent fractional crystallization F = 913%. For picrobasalts this value decreases to a range of 3-10% and for basanites to a range of 04%.

The AFC model curve assumes a possible origin of the mid-Miocene rocks from a common primary high-Mg melt at the proportion of initial-contaminated compositions 1:4. This proportion is adopted from concentrations of MgO (16.98 wt.%), Ni (756 ppm) in the sample VT-13-8 and from concentrations of MgO (1.97 wt.%), Ni (78.5 ppm) in the sample VT-13-9, respectively. From trace-element modeling, a magma source is calculated to contain Ol (57%), Opx (25%), Cpx (5%), Grt (6%), and Amph (5%) (see details in section 6.3). Concentrations of MgO and Ni in minerals and mineral/melt distribution coefficients are presented in Tables 2 and 3, respectively. Bulk distribution coefficients Dmineral/melt(Ni) = 7.39, Dmineral/melt(Mg) = 4.48 are

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melting line of peridotite with the Ni concentration of 1900 ppm (Sobolev et al., 2005) demonstrates compositions with higher MgO contents. Deviations of MgO from 20.4 wt.% to 12.5 wt.% in high-Mg rocks appear to reflect relative variations mineral proportions in a source region (see Fig. 13a in section 6.3).

Assumption on the origin of both moderate and high-Mg rocks from a common primary high-Mg melt contradicts the distribution of incompatible elements. We suggest that the trend of high-Mg rocks reflects temperature variations in a sub-lithospheric source region, whereas the trend of moderate-Mg melts results from melting in a shallow mantle source. Increasing MgO consistent with decreasing concentrations of incompatible trace elements in the groups of high-Mg basanites-basalts and picrobasalts-basalts indicates increasing temperature and degree of partial melting. In the group of high-Mg contaminated basalts-trachybasalts, the highest concentrations of incompatible elements are defined in a groundmass of the basalt VT-13-21 (Th = 7.3 ppm, La = 68 ppm, Sr = 4912 ppm, Ba = 1508 ppm). Conodes of these elements in bulk compositions and groundmasses have steeper slopes than the trend of other high-Mg rocks. The conodes of uncontaminated high-Mg basaltic samples lay sub-parallel to those of the basalt VT-13-21 on the Th vs. MgO and La vs. MgO plots and more gentle on the Sr vs. MgO and Ba vs. MgO ones. Minimal concentrations of La, Sr, and Ba in bulk compositions and groundmasses of high-Mg basalts overlap each other. As MgO decreases in the trachyandesite-like rock VT-13-9 from the bulk composition to the groundmass, La and Th increase and Ba decreases with minor variations of Sr (Fig. 10b, c, f and g). Unlike Th, La, Sr, and Ba, concentrations of K and Rb depend substantially on crustal contamination. The group of high-Mg contaminated basalts-trachybasalts and trachyandesite-like rock VT-13-9 show enrichment by both elements relative to other high-Mg groups (Fig. 10d and e).

Grouping of volcanic rocks is expressed also in trace-element ratios. Zr/Nb is lower in high-Mg rocks than in moderate-Mg ones (2.3-4.3 and 5.0-12.2, respectively). Zr/Y is higher in picrobasalts-basalts than in high-Mg basanites-basalts. In the diagram Zr/Y vs. Zr/Nb, the high-Mg groups of basalts-trachybasalts, basanites-basalts, and the moderate-Mg basalt group of K-Na series show correlated trends of these ratios. Those of moderate-Mg transitional (Na-K-Na) basalts-basaltic andesites demonstrate sufficient variations of Zr/Nb and limited ranges of Zr/Y (Fig. 10h).

6.3 Components of mantle sources

In terms of Nd and Sr isotope ratios, sources of the 16-14 Ma high-Mg magmas have a common depleted isotopic component (depleted mantle, DM) and other less depleted compositions. This is indicative for convective isotopic homogenization of a sub-lithospheric mantle material and its fractional differentiation, accompanied by incompatible element enrichment of resulting melts (Figs. 11 and 12). The separated melts, therefore, are derived due to high-temperature sub-lithospheric convective processes in the mantle and decreasing melting temperatures connected with increasing incompatible element abundances in a source. This type of magma generation differs from the generation of komatiite magmas in a plume that shows depletion in both isotopic and elemental signatures (Arndt et al., 2008).

High-Mg liquids from the Vitim Plateau are enriched by LREE, Th, Zr, and other incompatible elements that are introduced into melts due to prevailing partitioning from various mineral phases. Peridotite xenoliths, extracted from the spinel-garnet transition zone by ascending liquids of this area, show evidence on melt infiltration in the upper mantle (Glaser et al., 1999). Various types of mantle xenoliths have mutual genetic relationships. It was suggested that the high-temperature garnet-bearing rocks result from repeated affection of basaltic melts on the ambient upper mantle material. From the AFC calculations, it was also inferred that megacrysts and shallow xenoliths crystalized during the rise of basaltic liquids to the surface (Ashchepkov and Andre, 2002; Ashchepkov et al., 2011). From these observations,

metasomatized lherzolites and pyroxenites are examined as possi Miocene liquids.

The group of high-Mg basanites-basalts has initial 143Nd/144Ndt = 0.512867-0.512892 and

Sr/ Sr(/) = 0.703831-0.704183 in the range of the depleted mantle. The group of picrobasalts-basalts reveals less depleted isotopic signatures: initial 143Nd/144Ndt = 0.512730-0.512842, initial

Sr/ Sr(;) = 0.703778-0.704406 with a relative shift of data points along the mantle array towards the primitive mantle composition. Crustal contamination of high-Mg basalts-

trachybasalts and the trachyandesite-like rock VT-13-9 is reflected both in high 8,Sr/86Srw and low 143Nd/144Ndt. The group of moderate-Mg basalts-trachybasalts overlaps both the high-Mg basanite-basalt and picrobasalt-basalt compositions. The basaltic trachyandesite V-9 shows a more depleted signature: 143Nd/144Ndt = 0.512934, 87Sr/86Sri = 0.703740 (Fig. 11a).

The differences between the eruption age of 16-14 Ma and Sm-Nd ages of 42-24 Ma obtained for garnet- clinopyroxene pairs of xenoliths from tuffaceous-sedimentary strata of the Bereya volcanic center was explained by incomplete isotopic equilibrium in the xenoliths at the time of eruption due to insufficiently rapid diffusional exchange in coarse-grained rocks. Sm-Nd and Rb-Sr isotope data on xenoliths have been interpreted in terms of a ~2 Gyr depletion event relative to the primitive mantle values after Zindler and Hart (1986). It was suggested that the mantle peridotites have been attached to the continental lithosphere since that time and survived major tectono-thermal events (Ionov et al., 2005). In the area of study, no geological units of ~2 Gyr have been recorded, though.

From Pb isotope data on Cenozoic volcanic rocks, it was inferred that the mantle of the Baikal-Mongolian region was isotopically homogenized by vigorous convection at about 700600 Ma in connection with disintegration of the supercontinent Rodinia (Chuvashova, 2013). This event seems to be displayed also in Nd isotopic composition of the lithospheric mantle beneath the Vitim Plateau. On Fig. 11c and d, data points of high-Mg basanites and mantle xenoliths with extremely high 143Nd/144Nd ratio are distributed along a line with a slope of 640 Ma. One can assume that the material of lithospheric mantle xenoliths is a complementary product of mantle depletion of convective material responsible for disintegration of the supercontinent Rodinia, while the high-Mg basanites show time-integrated 143Nd/144Nd ratio of sub-lithospheric convective mantle. Offset of individual data points of xenoliths with increasing 147Sm/144Nd ratio might reflect minor transformation of the lithospheric mantle, associated with the formation of the island-arc Eravna terrane at about 440 Ma.

Evolution of the convective mantle material at shallow depths is reflected in shifting data points parallel to abscissa-axis with increasing 147Sm/144Nd. Under generally depleted isotopic signatures of Nd in garnet and spinel peridotites from the 16-14 Ma tuffaceous-sedimentary stratum, the only composite xenolith of garnet-bearing xenolith 313-4 falls into a field of volcanic rocks. In terms of isotopic compositions of Nd, volcanic rocks are comparable also to minerals from veins of composite xenoliths. Consequently, the shallow-mantle volcanic rocks cannot be derived from a source of the depleted lithospheric mantle and might be related to the processes, exhibited by the formation of veins in mantle rocks.

The reported data on veins (Ionov et al., 2005) exhibit heterogeneities of the lithospheric mantle on the scale of individual samples. On the one hand, the isotopic differences that might exist among mineral grains of chemical layering on centimeter or meter presumably, will not be seen except in xenolithic studies. On the other hand, references to lithospheric mantle components imply volumes of rock large enough to represent the source of individual batches of magma, perhaps of the dimension of several cubic kilometers. For instance, volume of 15 km3 moderately-Mg magmas, erupted in the Bereya volcanic center at 14-13 Ma, assumes a 20% partial melting and segregation of magmas from a mantle batch as large as 75 km3.

Mantle components of volcanic rocks are identified by cooperative use of the diagrams 143Nd/144Ndt vs. 100/Nd and (La/Yb)N vs. YbN (Figs. 12 and 13). The high-Mg basanite 93VBS-4 with the lowest 100/Nd (1.7 ppm-1) and relatively high 143Nd/144Ndt (0.512892, £Nd = +5) is the starting point for trends of high-Mg basanites-basalts (BSN - high Grt), picrobasalts-basalts

ClilU 111^,11

U dodlll L

1 Cdoili^

VV^ Ul

composition is designated as the common component of the depleted mantle (DM). The trend of high-Mg contaminated rocks yields low 143Nd/144Ndt and 100/Nd (elevated Nd concentrations). The data points of the 14-13 Ma moderate-Mg lavas offset from compositions of the 16-14 Ma high-Mg rocks with elevated 100/Nd (reduced Nd concentrations). The common trend of moderately-Mg lavas and minerals from veins (Fig. 12) is indicative for origin of the lavas through melting of the vein-bearing shallow mantle source.

In the Bereya volcanic center, high-Mg liquids were followed with moderate-Mg ones within a time interval of 1-2 million years. There was drastic transition from high-Mg (high temperature) to moderate-Mg (moderate temperature) melts. In Miocene lavas from the Hannuoba province (North China), a similar contrast existed between the groups of evolved alkali olivine basalts with high La/Yb and three other groups of basalts (tholeiitic, transitional and primitive alkaline) with low La/Yb. In the La/Yb vs. MgO diagram, the three groups of Hannuoba basalts are compared with the moderate-Mg lavas from the Bereya volcanic center (Zhi et al., 1990) (Fig. 13b).

The group of evolved alkali olivine basalts from the Hannuoba province is shifted relative to the high-Mg rocks from the Bereya center with decreasing MgO content. The Hannuoba and Bereya high La/Yb trends are parallel to each other. They both can be interpreted in connection with increasing degree of partial melting at elevated temperatures. From model calculations by Chuvashova et al. (2012), it was inferred that the Hannuoba evolved alkali olivine basalts were derived by 2-5% partial melting of a source that had 4.5-6.0% of garnet. These parameters are close to those of a source for high-Mg rocks from the Bereya volcanic center.

Interestingly, that MgO contents strongly vary in the high-Mg Bereya rocks only under elevated La/Yb and decrease with decreasing La/Yb (i.e. at the elevated temperature and increasing degree of partial melting). The Hannuoba evolved alkali olivine basalts, affected with olivine fractionation (Zhi et al., 1990), had primary liquids compositionally similar to high-Mg rocks from the Bereya volcanic center and other localities of the Transbaikalian melting domain shown on Fig. 1. The unfractionated high-Mg liquids from the Bereya volcanic center might be overheated above a liquidus.

&1 Lid

6.4 Temperature estimates

The mantle melts at relatively high temperature and low pressure. Convective heat transfer in the asthenosphere provides vertical temperature gradient reflected in a material flux without additional or lost heat. Asthenospheric adiabatic gradient of about 10 K*GPa-1 corresponds to the potential temperature (Tp) of a mantle source, which expresses the mantle temperature, projected along the adiabatic line to the atmospheric pressure at the Earth's surface. This parameter can vary from 1250 to 1400°C in mid-ocean ridges and rise up to 1550°C in hot plumes such as Hawaii. Change to the conductive transfer of geothermal heat flow in the lithosphere causes a large temperature gradient, in which the temperature increases with depth for 500 Kx GPa-1 (16°Cxkm-1) or more. As a result, melting of the mantle is dominated by decompression mechanisms that cause eruptions of basaltic magma through the extended lithosphere (Asimov, 2005; Herzberg et al., 2007).

Rocks enriched in Ni and Cr and showing Mg#=Mg/(Mg+Fe2+)>70 with corrections Fe3+ = 0.15 Fetotal are usually related to the primary mantle liquids. Besides, the composition of primary melt should be in equilibrium with the most magnesian olivine (Herzberg, 2007). High-Mg volcanic rocks from the Bereya volcanic center are characterized by Ni and Cr as high as 788 ppm and 1503 ppm, respectively, and by interval of Mg#=66.3-77.5 (predominated values 7576). Some Mg# values decrease to 64 because of crustal contamination and increase to 78 due to variable mineral proportions in the mantle source (see Figs. 10a and 13a). Unlike high-Mg rocks with typical signatures of primary mantle liquids, moderate-Mg rocks show relatively low Cr and


Mg =53.9-63.2 that demonstrate generation of such magmas from a modified mantle source.

A maximum liquidus temperature of olivine crystallization from an erupted melt or primary eruption temperature (Tpe) is estimated from an assumption on anhydrous composition of magma by various equations. Three of these by Kutolin (1966), Herzberg et al. (2007), and Arndt et al. (2008): Tpe (°C) = 1056.6 + 17.34 x MgO, Tpe (°C) = 935 + 33 x MgO - 0.37 x MgO2, Tpe (°C) = 20 + 1,000 x MgO yield approximately comparable results. For the high-Mg basanite VT-13-2 (MgO = 17.1wt.%), that belong to the predominated group with Mg#=75-76, the maximum liquidus crystallization temperature obtained is in the range of 1340-1390°C. Calculation of a potential temperature using equation Tp (°C) = 1463 + 12.7 x MgO - 2924/MgO (Herzberg et al., 2007) provides more satisfactory approximation to the temperature conditions in the mantle. The high-Mg basanite VT-13-2 yields an estimate Tp = 1510°C.

The assessment of potential temperature for the mantle beneath the Vitim Plateau was at its maximum in the outline of the Transbaikalian melting domain. The Tp for generation of high-Mg olivine melaleucetites from the Udokan volcanic field does not exceed 1430°C. This estimate is comparable with the potential mantle temperature for rocks from Ontong Java Plateau and Iceland (Early Tertiary). The modern rocks from Iceland yield the lower Tp estimates. Velocity models are indicative for association of Ontong Java Plateau and Iceland with upper mantle melting anomalies, not extending below the transition layer (Foulger, 2010). It is noteworthy that the calculated Tp = 1550°C for Hawaii (Herzberg et al., 2007) exceeds the upper mantle Tp interval and is comparable with the higher Tp interval for high-Mg rocks from Gorgona Island (Fig. 14). These high temperatures might reflect raising hot plumes from the deeper mantle according to the model suggested by Morgan (1971).

HERE Fig. 14

The maximum temperatures in the mantle can be constrained also from temperature calculations of mineral equilibrium in mantle xenoliths. The highest temperatures were obtained for xenoliths of pegmatoid orthopyroxenites and websterites with lamellas in pyroxenes. These xenoliths were found in basanites of the Udokan and Kamar volcanic fields. The temperature interval of 1350-1450 °C was calculated for initial Ca solubility in orthopyroxene solid solutions (Rasskazov, 1985; Rasskazov and Chuvashova, 2013). In the Udokan volcanic field, the high-temperature xenoliths occur in moderate-Mg basanite lavas erupted along the axial part of Udokan Range in the last 4 Ma. These basanites (as well as their inclusions) were separated in time and space from the 14 Ma high-Mg melaleucitites located in the northern part of the volcanic field. The potential temperature of the latter Tp = 1430°C is consistent with the initial temperature in the mantle pegmatoid orthopyroxenites. In the Kamar volcanic field, the high-temperature xenoliths were sampled in basanites from the Sukhoi volcano erupted at 13-12 Ma. The calculated potential temperature Tp = 1470°C for these basanites is close to the initial temperature of the mantle-derived pegmatoid orthopyroxenites.

In the outline of the Sayan-Mongolian melting domain, calculations for a basalt with the highest MgO content from the Dzhida volcanic field (10.1 wt.%) yield crystallization temperature interval Tpe = 1200-1230°C and estimate of potential temperature Tp = 1300°C. Melt inclusions in olivines from basalts of this melting domain also show relatively low homogenization temperatures: in Eastern Sayans 1220-1250°C, in the Dzhida volcanic field 1175-1220°C, and in Hangay 1210-1300°C (Simonov et al., 2012). The highest value of the homogenization temperature is comparable with a maximum estimate of potential temperature in the Sayan-Mongolian domain.

Relatively low temperatures in the mantle liquids of the Sayan-Mongolian melting domain might reflect shallow depths of the latter, thermally provided by asthenospheric convection in the upper mantle no deeper than 200 km, unlike the Transbaikalian melting domain that demonstrates decreasing velocities and possible relation of convective processes to the lower

jai i ш uit upper mantle and transition layer. The evaluated potential temperature Tp = 1300°C for the Sayan-Mongolian melting domain is within the temperature range of basalts from mid-ocean ridges Tp = 1280-1400°C. From seismic tomography, it was inferred that low-velocity anomalies beneath mid-ocean ridges do not extend deeper than 150 km (Anderson et al., 1992). The shallow nature of sources for volcanic rocks in the Sayan-Mongolian melting domain is also confirmed by 3He/4He ratios in olivines from lavas of the Hangay and Dzhida areas comparable with the values of this parameter in basalts from mid-ocean ridges (Barry et al., 2007).

6.5 Depth estimates

LREE enrichment of magmas relative to HREE is controlled principally by the presence of garnet in a source region. From definition of Gd/Yb = 0.15 in a garnet, it was inferred that its occurrence in a residual solid maintains a higher Gd/Yb ratio in a partial melt as compared to a source (Walter, 1998). The value Gd/Yb = 1.6 was suggested as the lower limit for a liquid from a garnet-bearing source (Sobolev et al., 2011). On the one hand, from the Gd/Yb criterion, all the mid-Miocene melts, erupted in the Bereya volcanic center, should be assigned to garnet-bearing sources. Gd/Yb ranges from 2.5 to 3.4 in basalts-basaltic andesites and increases to the intervals 4.9-5.7, 3.6-6.5, 5.3-8.1, and 7.9-9.3 in Na basalts, basalts-trachybasalts, high-Mg basanites-basalts, and picrobasalts-basalts, respectively. On the other hand, the groups of basalts-basaltic andesites demonstrate the narrow (La/Yb)N, which is 7-8 times higher than that in the pyrolite, and Yb concentrations varying from 1.38 to 1.95 ppm. These are characteristics of a source enriched by LREE with no garnet or low garnet content.

Deep magma-generating processes in the sub-lithospheric mantle at 16-14 Ma are indicated by the potassic affinity of picrobasalts-basalts. The high activity of potassium in a deep part of the upper mantle is reflected in potassium-bearing clinopyroxenes in xenoliths from kimberlites and inclusions in diamonds, in the diamond-bearing pyroxene-garnet-carbonaceous rocks of the Kokchetav metamorphic complex and confirmed by experiments at high pressures and temperatures (Sobolev, 1974; Tsuruta and Takahashi, 1998; Perchuk et al., 2002; Shatsky et al., 2006; Foley et al., 2009). Natural assemblages with potassium-bearing clinopyroxenes are limited to the pressure range 5-7 GPa (Safonov et al., 2005) that corresponds to depths of 160220 km.

From the empirical equation P (kbar) = 213.6 - 4.05 x SiO2 (wt.%) (Scarrow and Cox, 1995), high-Mg basanites-basalts and picrobasalts-basalts yield pressure ranges of 36.4-48.6 and 41.4-44.3 kbar, respectively. These pressures correspond to generations of high-Mg melts beneath the Bereya volcanic center in the depth intervals 115-150 and 135-140 km during accumulation of the middle and upper layers of the 16-14 Ma stratum (Fig. 5). The maximum depth up to 150 km is characteristic of the sub-group of high-Mg basanites-basalts with an elevated degree of melting (BSN - high F), which is the best estimate for Tp.

A depth of lithospheric mantle sources, involved in melting about 14-13 Ma, is limited by a garnet-spinel transition. Experimental data of the CaO-MgO-Al2O3-SiO2 system (CMAS) show this transition below the solidus in the pressure ranges 1.8-2.0 GPa at 1200°C and 2.6-2.7 GPa at 1500°C (at depth 60-85 km). In Cr-bearing systems, spinel coexists with garnet. Increasing concentrations of Cr shift a reaction of garnet introduction to the higher pressures. The presence of Fe2+ has the opposite effect (Klemme, 2004; Klemme and O'Neill, 2000). Similar estimates for the garnet-spinel transition have been obtained through experiments on enriched and depleted peridotites. Under lower pressure conditions, melting is assumed to begin in garnet pyroxenite veins (Robinson and Wood, 1998).

6.6 Partial melting modeling

For modeling partial melting of enriched sources, we use mean compositions of minerals from Grt- and Grt-Phl lherzolites and also pyroxenites with phlogopite, ilmenite, and kaersutite,

as well as other ueep-seaieu xenuliuis mat satisfy lu the conditions. (1) a mineral assemblage Ol + Cpx + Grt + Amph ± Phl ± Ap represents a possible model source, (2) P-T parameters uf xenuliths correspond tu the upper mantle conditions (i.e. T>1000°C), and (3) analytical data fur trace-element concentrations in minerals are of high quality. The trace element concentrations were defined in xenuliths from picrobasalts and basanites of the Vitim Plateau (Iunuv et al., 1997; Glaser et al., 1999; Litasov et al., 2000; Ashchepkuv and Andre, 2002; Ashchepkuv et al., 2003, 2011). Additional trace-element data un Ta, Th, U concentrations in amphibole were taken fur xenuliths from basalts of the SW Tian Shan (Zheng et al., 2006). The data on the type-A apatite were used after O'Reilly and Griffin (2000) (Table 2).

Equilibrium partial melting was modeled using the equation (Shaw, 1970): Ci (0)

Ci =-, where C - concentration of element i in a melt; C/0) - concentration of

' Di + F(1 - Pi) ' ^

element i in the source; Di - bulk distribution coefficient for the source material, F - melt

fraction; Pi = £ XK{ - bulk distribution coefficient fur material involved in melting; X - weight

fraction of phase j; K - mineral/melt distribution coefficient fur phase j. Mineral/melt

distribution coefficients are given in Table 3. The clinopyroxene fraction in the source material

was increased about two times and the olivine fraction was decreased relative to their contents in

the source according to the experimental study of garnet-bearing rucks (Walter, 1998).

From results of trace-element modeling, we infer that the 16-14 Ma high-Mg basanites-basalts and picrobasalts-basalts were originated from a source that has: Ol (48-57%), Opx (25%), Grt (6-12%), Cpx (5%), Amph (5-7%), Phl (2-3%), and Ap (<0.07%). Some discrepancies between rucks and model melts were found in terms of Rb, Th, and K. Besides, the maximum of Sr in the trace element patterns are nut explained by this model. Melt fractions (F) vary from 0.04 to 0.06 (Fig. 15a and b).

In Fig. 13, some data points of high-Mg basanites-basalts fall at a line of the melt fraction F = ~0.045. Neodymium isotope ratios of this sub-group decrease as the garnet content in the source increases from 6 to 12%. Another sub-group of high-Mg basanites-basalts yields melt fractions increasing up to 0.075 in the source with 6-8% of garnet. The neodymium isotope ratios of this sub-group range in narrow interval as Nd concentrations decrease due to increasing melt fractions. The source of picrobasalts-basalts shows about 10% Grt and melt fraction F = 0.04-0.05.

The 14-13 Ma lavas were derived from sources with relatively low garnet (or garnet-free) and elevated clinopyroxene contents: basalts and trachybasalts - Ol (49.5%), Opx (25%), Grt (<3.5%), Cpx (12%), Amph (8%), Phl (2%), Ap (<0.07%), F = 0.05-0.1; basalts-basaltic andesites - Ol (54.5%), Opx (25%), Cpx (15%), Amph (5%), Phl (<0.5%), Ap (<0.05%), F = 0.15-0.20. In basalts-trachybasalts of K-Na series, LREE are variable in transitional (Na-K-Na) rocks, while HREE remain constant (Fig. 8b). These REE patterns are typical for equilibrium partial melting in a mantle source. Increasing La/Yb reflects decreasing degree of melting (Rollinson, 1993).

From the results of trace element modeling in the (La/Yb)N vs. YbN coordinates, high-Mg basanites-basalts are examined as derivatives from a source with the garnet content from 6 to 12%. A high-garnet trend (BSN - high Grt) in Fig. 12 corresponds to 6-12% of Grt at relatively low degree of melting (F = 0.042-0.046) in Fig. 13a. A similar source was identified for basanites from the Kamar volcanic field located near the South Baikal basin (Rasskazov et al., 2013b). In the source region, garnet concentrations might increase due to depth-dependent increasing rock density and due to garnet fractionation relative to less dense pyroxene and olivine. Some data points of the high-Mg basanite-basalt group from the Bereya volcanic center are shifted, however, along a line with garnet content of 6% towards melt fractions F up to 0.075.

X 11V^ 111 Cl/Vllll Cil VJV^JJlll V^O ULlllClL^iJ

to 150 km from these data points are consistent with increasing MgO content due to high-temperature magma generation. A trend for picrobasalts-basalts (PB) crosses the middle part of the BSN - high Grt one. Unlike the latter source, garnet content of the PB trend remains constant under varying degree of melting (F = 0.039-0.050). Contaminated basalts-trachybasalts with elevated (La/Yb)N fall at the line of sources with 12% of garnet.

In the (La/Yb)N vs. YbN diagram, the data points of lavas erupted about 14-13 Ma are plotted between model lines that show garnet content 3.5% and no garnet in a source region. Firstly, melted material from the shallow garnet-free source (or containing a small amount of garnet) provided voluminous eruptions (2/3 of the whole 14-13 Ma volume). The degree of partial melting was at a maximum (F = 0.15-0.20). The mantle source for coeval small portion of Na-basaltic lavas, derived from the garnet-bearing source (1-2% of Grt), experienced relatively weak fusion (F = 0.05-0.07). The final lava portion, which amounted to 1/4 of the volume of all the 14-13 Ma eruptions, was also derived from the garnet-bearing source (up to 3.5% Grt) with an elevated degree of partial melting (F = 0.05-0.10).

6.7 Change of magmatism at 16-13 Ma

We suggest that the mid-Miocene temporal change from high- to moderate-Mg liquids in the Bereya volcanic center was due to a thermal impact on the lithosphere from the convective mantle of the Transbaikalian melting domain (Fig. 16).

In the time interval of 16-14 Ma, initial high-Mg magmas were contaminated by disintegrated crustal material. The composition of primary magmas that rise from the mantle source was distorted due to the partial assimilation of this material. Perhaps, the primary mantle melts were similar in composition to the high-Mg basanite-basaltic melts erupted in the Bereya volcanic center after high-Mg contaminated basalts-trachybasalts.

The high-Mg basanite-basaltic and picrobasalts-basaltic magmas of the second and third episodes exhibited uncontaminated primary mantle liquids. In adiabatically ascending mantle material, melting depth depends on a heat content of a magmatic system. When heat supply is high, melting is initiated at a deeper level; when heat supply is low, melting is only provided in relatively shallow conditions (Fukuyama, 1985).

The mid-Miocene eruptions of high-Mg magmatic liquids in the Bereya volcanic center from the 115-150 km mantle source were due to adiabatic introduction of heat enough to melting at such a deep level. Created buoyancy of partially molten material exceeded that of crystalline phases in the ambient asthenospheric mantle. As a result, the melts, released from the matrix, moved upwards and erupted on the surface.

The deepest high-Mg basanites-basalts were produced through relatively homogeneous melting by excessive heat that caused moderate melting in the magmatic column F = 6%. An initial melt region had the upper limit at the depth 115 km governed by release of buoyant high-Mg liquids from a fractionated matrix with the garnet content varying from 6 to 12%. Shortly afterwards melting processes localized within a region that had more constant garnet content and yielded picrobasalt-basaltic liquids.

At the initial stage of 16-14 Ma, igneous channels provided heat supply to the shallow part of the lithospheric mantle. High-degree decompression melting of a garnet-free (or low-garnet) source caused thermal impact on the lithosphere, uplift and erosional dissection of the area and resulted in voluminous eruptions of moderate-Mg basalts and basaltic andesites about 14-13 Ma. This impulse of the shallow melting, however, was short-lived and changed to less voluminous melting in the deeper garnet-bearing part of the lithospheric mantle.

7. Conclusion

A comparative study of high- and moderate-Mg volcanic rocks erupted in the Bereya volcanic center of the Vitim Plateau, at 16-14 and 14-13 Ma, respectively, has revealed specific

(low-volume) unit shows the

temporal vanauuus of iucr cumpuíiuuna in

sequence of high-Mg rocks: (1) contaminated basalts-trachybasalts of K-Na series, (2) basanites-basalts of transitional (K-Na-K) compositions, and (3) picrobasalts-basalts of K series. The latter (voluminous) unit demonstrates the sequence of moderate-Mg rocks: (1) basalts-basaltic andesites of transitional (Na-K-Na) compositions and (2) basalts-trachybasalts of K-Na series.

It has been found that the initial portion of high-Mg liquids in the 16-14 Ma unit was strongly affected by crustal contamination. Later on, the uncontaminated melt portions of this unit revealed various temperature effects in a mantle anomaly; the elevated degree of melting in its source as deep as 150 km appeared to reflect adiabatic heat transfer, whereas the less advanced melting at shallower depths (up to 115 km) was caused by relative decreasing temperature beneath the lithosphere. Variations of major oxides, trace elements, and isotopes in the 14-13 Ma rock sequence showed initial melts equilibrated at the shallow garnet-free mantle source with subsequently reduced degree of melting of garnet-bearing material. The shallow melts revealed no melting of relatively depleted lithospheric material, evidenced by mantle xenoliths from the lithosphere.

The main purpose of this paper is to connect the mid-Miocene volcanic activity, initiated in deep and further developed in shallow mantle sources, with destabilization and rifting of the lithosphere in the Baikal Rift Zone. It is suggested that voluminous lava eruptions from the shallow lithospheric source at about 14-13 Ma reflected a thermal impact on the lithosphere beneath the Vitim volcanic field by hot convective material of the Transbaikalian melting domain that had a potential temperature Tp as high as 1510°C, comparable with that parameter in the upper mantle of the Ontong Java Plateau and Iceland hot spots. This thermal impact caused extension of the lithosphere, accompanied by subsidence of rift basins on the surface. In other volcanic fields of the Transbaikalian melting domain (Udokan, Dariganga, Kamar), Tp estimates decreased to 1430°C, whereas in the shallower Sayan-Mongolian melting domain, Tp did not exceed 1300°C.


This work was supported by the Russian Science Foundation for Basic Research (project 14-05-31328). Sample preparation for trace-element and isotope analytical work was performed by M.E. Markova and E.V. Saranina, isotope mass-spectrometric determinations by N.N. Fefelov, and major-oxide determinations by G.V. Bondareva and M.M. Samoilenko. Agilent 7500ce and Finnigan MAT 262 mass spectrometers are affiliated with the Common Use Centers "Ultramicroanalysis" and "Geodynamics and Geochronology", respectively. We thank Nick Roberts and two anonymous reviewers for constructive discussion of the manuscript.

Appendix A. Supplementary data

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V lllcllldlll, J3., JdlllcZIL/, n., JUIUI1, ilclUl, 1V1., I70I. _L/1SLI1UUL1UI1 L/UclllL/1 cilia U1 Illdjv

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Chuvashova Irina: Senior Research Scientist at the Institute of the Earth's Crust SB RAS, Irkutsk, Russia (since 2004), assistant professor of Dynamic Geology Chair at the Irkutsk State University (since 2015), M.S. (2006) from the Irkutsk State University, Ph.D. (2010) from the Institute of the Earth's Crust SB RAS. She worked in late Mesozoic and Cenozoic volcanic fields of Mongolia, China, Kyrgyzstan, Sakhalin, Amur region, and Siberia. Results of her research in geochemistry and petrology of volcanic rocks have been presented in 118 publications that include 6 monographs and 18 peer-reviewed papers.

Rasskazov Sergei: Head of Laboratory for isotopic and geochronological studies at the Institute of the Earth's Crust SB RAS, Irkutsk, Russia (since 1996), Head of Dynamic Geology Chair at the Irkutsk State University (since 2012), M.S. (1976) from the Irkutsk State University, Ph.D. (1982), Dr. (1992), and Professor (2005) of Perology and Volcanology from the Institute of the Earth's Crust SB RAS. He has studied Cenozoic rifts in Central and East Asia, North America, and East Africa. His main interest is origin of volcanic rocks, erupted at the latest geodynamic stage of the evolving Earth. He has published more than 500 works that include 17 monographs and 132 peer-reviewed papers.

Yasnygina Tatyana: Senior Research Scientist at the Institute of the Earth's Crust SB RAS, Irkutsk, Russia (since 2004), M.S. (1993) from the Irkutsk State University, Ph.D. (2004) from the Institute of the Earth's Crust SB RAS. Her scientific interests are geochemistry of volcanic rocks and crude oils, trace-element melting modeling. The results of her research have been presented in more than 110 publications that include 2 monographs and 29 peer-reviewed papers.

rlgUic LdpilUlls

Figure 1. Spatial distribution of high- and moderate-Mg mid-Miocene volcanic rocks in Inner Asia relative to the S-wave velocity image at a depth of 250 km after Yanovskaya and Kozhevnikov (2003). Volcanic fields with mid-Miocene rocks: high-Mg - Vitim (VT), Udokan (UD), Kamar (KM), Dariganga (DR); moderate-Mg - Tunka (TN), Dzhida (DZ), Eastern Sayans (ES), Khubsugul (KH), Dzabkhan (DB), Valley of Lakes (VL), Eastern Hangay (EH), Ugii-Nur (UN), Hannuoba (HN). Next to each of mid-Miocene locations of high- and moderate-Mg rocks, a maximum content of MgO (wt.%) is shown in parentheses.

Figure 2. Mid-Miocene river valleys buried beneath lavas of the Vitim volcanic field (a) and tectono-stratigraphic terranes of Western Transbaikal; (b) In the scheme (a), the Northern (N), Central (C) and Southern (S) paleovalleys and volcanic centers: Bereya (Br), Sharbahtay (Sh), Ingur (In), Salbulin (Sl), Antase (An), Kolichikan (Kl), Db (Dybryn), Nm (Namaru), Sirinikta (Sr), Muhal (Mh), Yaole (Ya), Khoigot (Kh), and Malyi Amalat (Ma) are shown after Rasskazov et al. (2000). In the scheme (b), the terrane boundaries of the island arc Eravna (Er), back-arc turbidite Ikat (Ik), Riphean Baikal-Muya (BM), and composite West Stanovoy (WS) are designated after Belichenko et al. (2006).

Figure 3. Location of the quarry and lines of drill holes that exposed, respectively, high-and moderate-Mg rocks in the Bereya volcanic center.

Figure 4. Relations between the 14-13 Ma and younger (12.6-0.6 Ma) volcano-sedimentary units in the Bereya volcanic center. The drill-hole lines of the sections BA, BC, and DE are shown in Fig. 3.

Figure 5. Schematic cross-section that demonstrates relations between the 16-14 and 1413 Ma volcano-sedimentary units in the Bereya volcanic center. Location of the quarry that exposed sediments with fragments of high-Mg volcanic rocks is shown in Fig. 3.

Figure 6. Coviriation diagrams of K2O vs. K2O/Na2O (a), K2O vs. SiO2 (b), total alkalis vs. silica (TAS) (c), and MgO vs. SiO2 (d) diagrams for mid-Miocene rocks from the Bereya volcanic center. Among volcanic rocks from the 16-14 Ma layers (Fig. 5), we define three groups: 1) high-Mg basalts and trachybasalts of K-Na series, 2) basanites and basalts of transitional (K-Na-K) compositions, and 3) picrobasalts and basalts of K series. Among volcanic rocks of the 14-13 Ma unit (Fig. 4), we recognize two groups: 1) basalts and basaltic andesites of transitional (Na-K-Na) compositions and 2) basalts and trachybasalts of K-Na series. In the diagram (a), a field of K series of moderate-Mg lavas from the Ugii-Nur volcanic field, Central Mongolia is shown after Rasskazov et al. (2012). In the diagram (b), the dividing lines of the series in terms of the potassium contents are used after Rollinson (1993). Fields in the diagram (c): P - picrite, PB - picrobasalt, BSN, TPH - basanite, tephrite, PHT -

phonotephrite, B - basalt, TB - trachybasalt, BA - basaltic andesite, BTA - basaltic trachyandesite, TA - trachyandesite, F - foidite. The dividing lines in (c) are after (Le Bas, Streckeisen, 1991). Contents of major oxides are recalculated to 100 % without loss on ignition.

Figure 7. Diagrams of CIPW normative Ne-(Hy+Qtz) vs. 100xAn/(An+Ab) (a) and Ne-(Hy+Qtz) vs. MgO (b). Symbols are as in Fig. 6.

Figure 8. Pyrolite-normalized concentrations of incompatible elements (a) and chondrite-normalized contents of REE (b). The pyrolite and chondrite compositions are after McDonough and Sun (1995).

Figure 9. Binary diagrams of Th/Yb vs. Ta/Yb (a) and initial 87Sr/86Sr vs. 1000/Sr (b). Due to young ages of rocks and low Rb/Sr, the initial Sr isotopic ratios do not differ from the measured values. E-MORB and OIB are from McDonough and Sun (1995). Symbols are as in Fig. 6.

Figure 10. Ni, Th, La, K2O, Rb, Sr, and Ba vs. MgO (a-g) and Zr/Y vs. Zr/Nb (h). Symbols are as in Fig. 6.

Figure 11. Initial 143Nd/144Nd vs. initial 87Sr/86Sr (a) and 143Nd/144Nd vs. 147Sm/144Nd (c). Insets (b) and (d) show data on xenoliths from the 16-14 Ma high-Mg rocks (Ionov et al., 2005). Symbols are as in Fig. 6.

Figure 12. Initial 143Nd/144Nd vs. 100/Nd. Symbols are as in Fig. 6.

Figure 13. (a) (La/Yb)N vs. (Yb)N and (b) La/Yb vs. MgO diagrams. For simplification the diagram (a), data points of bulk and groundmass compositions for only two samples of the contaminated basalts that yield the highest (La/Yb)N are plotted. Abbreviations: Grt - garnet, Cpx - clinopyroxene, Amph - amphibole, Phl - phlogopite, Ap - apatite. Ratios of minerals in sources are based on trace-element modeling (see the text). Mineral/melt distribution coefficients for La and Yb are shown in Table 3. The primitive mantle composition is adopted after McDonough and Sun (1995). In diagram (b), data obtained are compared with data fields of Miocene lavas from the Hannuoba Province of China: the groups of evolved alkali olivine basalts (EV-AOB), primitive alkali olivine basalts (PR-AOB), transitional basalts (TRB), and tholeiitic basalts (TLB) after Zhi et al. (1990).

Figure 14. Primary eruption and mantle potential temperatures as a function of the MgO contents in primary magmas. Primary eruption temperatures are anhydrous olivine liquidus temperatures at 1 bar calculated for the primary magmas using the simplified equation Tpe (°C)=935+33xMgO - 0.37xMgO2. The bracket shows the uncertainty at the 2o level. Potential temperatures are calculated from the equation TP(°C) = 1463+12.7xMgO - 2924/MgO (Herzberg et al., 2007). Arrows display the effects of H2O content on liquidus temperature. Partial crystallization of primary magmas will yield actual eruption temperatures that are lower than


1111 U"U

iiu^c uciscms

, OJP - Ontong

Java Plateau, I (ET) - Island (early Tertiary), I (PD) - Island (present day). Highlighted in three temperature ranges are: (1) Tp = 1550 °C and above - plumes rising from the deeper parts of the mantle (Gorgona, Hawaii), (2) Tp = 1430-1510 °C - upper-mantle melting anomalies extending from the mantle transition zone (Ontong Java Plateau, Iceland, Transbaikalian domain), and (3) Tp = 1280-1400 °C - melting anomalies in the shallow part of the upper mantle (mid-ocean ridges, Sayan-Mongolian domain).

Figure 15. Comparisons of trace element (a, c) and REE (b, d) patterns of rocks (lines with markers) and calculated model liquids (black lines without markers). Symbols are as in Fig. 6. (a, b) - high-Mg basanites-basalts and picrobasalts-basalts of 16-14 Ma, (c, d) - moderate-Mg basalts, basaltic andesites, and trachybasalts of 14-13 Ma. Normalization values are after (McDonough and Sun, 1995). The calculated degrees of melting (melt fractions) and garnet contents in sources are indicated (see discussion in the text).

Figure 16. Change of sources for rocks of the Bereya volcanic center in terms of the mid-Miocene thermal impact on the lithosphere by a sub-lithospheric hot convective material from the Transbaikalian melting domain. Tp - potential temperature, F - degree of partial melting.

I dUlc 1

Major oxides (wt.%), trace elements (ppm), and Sr, Nd isotope ratios in representative compositions of the rock groups from the Bereya volcanic center

Age (Ma) 16-14 14- -13

High-Mg Basalts-

Group Contaminated rocks, K-Na series basanites-basalts of transitional (K-Na-K) Picrobasalts -basalts, K series basaltic andesites of transitional (Na-K-Na) Basalts-trachybasalts, K-Na series

composition composition

Sample VT-13-9 VT-13- 9gm VT-13-20 VT-13-19 V9 V18

SiO2 56.20 47.90 42.55 42.51 52.89 50.71

TiO2 1.13 1.89 1.85 2.09 1.85 2.28

AI2O3 15.30 13.03 9.79 10.27 14.36 16.13

Fe2O3 3.95 5.84 4.20 5.46 11.18* 10.75*

FeO 2.96 5.29 7.54 6.61 N.d. N.d.

MnO 0.11 0.16 0.16 /<0.15 0.18 0.17

MgO 5.52 10.12 17.23 13.87 7.23 5.61

CaO 5.22 7.11 8.06 8.99 8.87 8.69

Na2O 3.80 2.90 2.22 1.28 3.08 3.88

K2O 3.31 2.11 1.60 1.61 0.63 1.70

P2O5 0.36 0.62 0.64 0.73 0.23 0.48

LOI 2.04 2.98 4.07 6.61 0.00 0.00

CyMMa 99.90 99.95 99.91 100.18 100.50 100.40

Sc 15.9 22.2 12.3 11.5 22.7 19.0

V 105 163 147 156 169 183

Co 30 47 63 57 44 34

Cu 5 35 40 35 55 40

Ni 214 326 744 458 131 54

Cr 238 550 1301 651 193 85

Rb 73 45 26 22 9 24

Sr 632 769 791 4042 389 713

Y 22.4 26.1 18.9 20.9 23.0 24.0

Zr 88 156 201 227 123 198

Nb 21.9 38.2 49.1 57.0 14.5 32.7

Cs 1.26 0.62 0.19 0.25 N.d. N.d.

Ba 1103 845 651 1048 194 470

La 48.3 56.7 44.4 54.3 10.8 25.8

Ce 95.6 114.5 89.6 107.0 23.4 52.1

Pr 10.9 13.1 10.9 12.7 3.2 6.6

Nd 42.2 50.9 41.3 49.4 14.2 26.5

Sm 7.23 9.87 8.64 9.71 4.08 6.16

Eu 2.37 2.92 2.85 3.34 1.52 2.09

Gd 6.12 9.68 6.59 7.71 3.89 5.83

Tb 0.77 1.24 0.98 1.13 0.73 0.87

Dy 4.69 5.73 4.54 5.08 4.15 4.59

Ho 0.91 0.97 0.72 0.79 0.78 0.84

Er 2.26 2.35 1.56 1.68 2.00 2.08

Yb 1.72 1.80 0.86 0.97 1.57 1.63

Lu 0.21 0.26 0.09 0.11 0.22 0.22

Hi 1. 96 4.3 3 4. /0 5.4 7 3.12 4.4 7

Ta 1.36 2.33 3.18 3.62 0.88 2.03

Pb 12.4 9.7 4.5 5.1 N.d. N.d.

Th 5.39 5.77 5.51 6.51 1.18 2.88

U 0.81 1.11 1.17 2.44 0.32 0.81

87Sr/86Sr 0.706556 N.d. 0.704069 0.704243 0.70374 0.70431

2o 0.000036 0.000050 0.000075 0.00001 0.00001

144Nd/143Nd 0.512444 N.d. 0.512852 0.512809 0.512934 0.512788

2o 0.000049 0.000024 0.000044 0.000009 0.000009

The group of high-Mg contaminated rocks is exhibited with bulk and groundmass compositions VT-13-9 and VT-13-9gm, respectively. *Defined only total iron content in the form of Fe2O3. N.d. - no data. Data on samples from the 16-14 Ma unit were obtained by authors, on those from the 14-13 Ma unit by Harris (1998).

Table 2

Trace element concentrations (ppm) in the pyrolite and mean concentration in minerals from xenoliths used for calculations of the model magmatic sources beneath Vitim Plateau

Element Pyrolite Cpx Grt Phl Amph Ap

Rb 0.6 0.45 N.d. 263 6 15

Ba 6.6 11 0.03 3364 178 187

Th 0.0795 0.07 N.d. 0.005 1.50 97

U 0.0203 0.01 N.d. N.d. 0.50 23

K 240 83 N.d. 73390 15994 N.d.

Nb 0.658 0.79 0.14 19.1 42.8 0.66

Ta 0.037 0.028 N.d. 0.38 3.23 h. g.

La 0.648 2.33 0.01 0.69 6.01 2300*

Ce 1.675 9.42 0.07 0.9 21.3 4042

Pr 0.254 2* 0.02 0.27 4* 377

Sr 19.9 104 0.18 212 467 13295

Zr 10.5 95.5 24.3 18 119 24

Hf 0.283 2.76 0.24 0.31 4.94 0.71*

Sm 0.406 3.25 0.48 0.14 5.5 170*

Ti 1260 7008 1112 47098 25120 N.d.

Y 4.3 8.29 27.7 0.78 11.4 173

Yb 0.441 0.40 2.92 0.06 0.52 7.67

Ni 1960 468 47 N.d. 868 N.d.

MgO, % 37.80 14.68 20.72 19.40 16.76 N.d.

Data sources: Ashchepkov, Andre, 2002; Ashchepkov et al., 2003; 2011; Ionov et al., 1997; Glaser et al., 1999; Litasov et al., 2000; O'Reilly, Griffin, 2000. Th and U in Amph is after Zheng et al. (2006). The pyrolite composition is after McDonough, Sun (1995). *The values obtained after interpolation from normalized spectra. N.d. - no data.

Table 3

The mineral/basaltic melt distribution coefficients used for the partial melting and AFC modeling

Ele ment Ol Opx Cpx Grt Phl Amph Ap

Rb 0.00002 0.0001 0.0004 0.00002 5.18 0.2 0.01

Ba 0.000005 0.00006 0.0007 0.00007 3.48 0.16 1.555

Th 0.000007 0.0002 0.013 0.0021 0.0014 0.0039 2.131

U 0.000009 0.00004 0.006 0.011 0.0011 0.0041 2.131*

K 0.00002 0.0001 0.07 0.013 3.67 0.58 N.d.

Nb 0.00005 0.0013 0.0027 0.02 0.0853 0.159 0.001

Ta 0.00005 0.0025 0.012 0.05 0.2255 0.16 0.001*

La 0.0002 0.003 0.054 0.0007 0.0004 0.055 7.39

Ce 0.00007 0.0021 0.086 0.0026 0.0006 0.096 5.965

Pr 0.0003 0.0022 0.154 0.01 > 0.0009 0.17 4.64*

Sr 0.00004 0.0015 0.096 0.0007 0.183 0.298 3.055

Zr 0.001 0.012 0.123 0.2 0.0232 0.18 0.012

Hf 0.0029 0.019 0.256 0.4 0.048 0.29 0.01

Sm 0.0009 0.0037 0.291 0.1 0.0008 0.33 4.603

Ti 0.015 0.12 0.384 0.32 1.768 1.29 N.d.

Y 0.0082 0.02 0.467 2.5 0.007 0.52 3.43

Yb 0.024 0.032 0.43 6.4 0.023 0.47 0.93

Ni 7 3.1 1.4 0.19 5 3.4 N.d.

Mg 5 4.7 5.4 N.d. N.d. 3.8 N.d.

Data sources: olivine (Ol): Ni, Mg (Wang, Gaetani, 2008), other elements (Halliday et al., 1995); orthopyroxene (Opx): Ta (Green et al., 2000), Ti (Girnis et al., 2006), Yb (Kennedy et al., 1993), Ni (O'Reilly et al., 1991), Mg (Henderson, 1982), other elements (Halliday et al., 1995); clinopyroxene (Cpx): Rb (Halliday et al., 1995), Ba, K, REE, Zr, Hf, Ti, Y (Hart, Dunn, 1993), Th and U (Hauri et al., 1994), Nb and Sr (Foley et al., 1996), Ta (Green et al., 1989), Ni (O'Reilly et al., 1991), Mg (Villemant et al., 1981); garnet (Grt): Rb, Ba, Th, U, K, REE, Sr, Y (Halliday et al., 1995), Nb, Ta, Zr, Hf, Ti (Girnis et al., 2006), Ni (O'Reilly, Griffin, 1995); phlogopite (Phl): Rb, Ba, Nb, Sr, Zr, Y (Foley et al., 1996), Th, U, K, Ti (LaTourrette et al., 1995), Ta, La (Gregoire et al., 2000), Ce, Pr, Hf, Sm, Yb (Ionov et al., 1997), Ni (O'Reilly et al., 1991); amph (Amph): Ta (Zack et al., 1997), Zr, Hf, Sm (Ionov et al., 1997), Ni (O'Reilly et al., 1991), Mg (Henderson, 1982), other elements (LaTourrette et al.,

1995); apatite (Ap): Rb, Nb, Ta, Hf (Ionov et al., 1997), other elements (Chazot et al.,

1996). *Values after interpolation. N.d. - no data.


DOMAIN (depth 50-200 km) /


UD (14.0) \

ES (9 8) (9>

•V KH (10.0)' "



VT (20.4)

VL (9.4)*".,

S: -v;

4.5 4.4 4.3 4.2

E4-1 7

Vs, km-s


\ 30 N

150° E

It Late Cenozoic volcanic field * Location of high-Mg rocks (maximal MgO = 14.0-20.4 wt.%)

Location of moderate-Mg rocks (maximal MgO = 7.2-10.1 wt.%)

^ Paleovalleys formed between 16 and 14 Ma Northwestern boundary of the area

CJ Volcanic rocks burjed beneath basa|ts not buried ocupied by Mesozoic trachybasalts

BJ*« Volcanic * center

| Area of study


W4454, wellhead-1139 m.. 110,219,234,240 m

W4585, wellhead-1125 m

W4633, wellhead -1140 m 95,101,106,190 m

W4590, wellhead -1135 m

W4657, wellhead -1015 m 26,46,57,95, 103 m

W4655, wellhead - 1003 m 8,23,37 m

14-13 Ma unit Lavas:

K-Na series transitional between K-Na and Na series


Younger units of the time interval 12.6-0.6 Ma I Lavas

Pyroclastic material


Basement 114.31 K-Ar age (Ma)

a Selected drill hole, wellhead altitude, depth of sampling

16-14 Ma

Columnar basalts, lenses of sediments

Gravel layer 3, blocks of picrobasalts (PB)

Gravel layer 2,

blocks of high-Mg basanites(BS) Gravel layer 1,

blocks of Mg-basalts-trahybasalts (B) Basement

14-13 Ma

High-Mg rocks, 16-14 Ma basalts-trachybasalts groundmass bulk composition (a) basanites-basalts picrobasalts-basalts Moderate-Mg rocks, 14-13 Ma ^ basalts-basaltic andesites VT139 P ' / Î / > f'Ugii-Nur volcanic field, ^ Central Mongolia # (15.5-14.2 Ma) n

basalts-trachybasalts i0y j* N

CO <D :

<D (O l____ .J

K-Na series K series a

0.8 1.2 K,0/Na,0

i o * 2

I I I Ultrahigh-K

i series

Absarokite I I



{ 'Jife^ > I

I___--___ Medium-K

^—" / 1 \ series xl * \

r^^Low^K b


Si02 (wt.%)


TPH BSN / 1° / /TB\ B7/4 \ VT-13-9

/V i AiA (

PB 1, » a B BA

MgO>18 %p[y c

Si02 (wt.%)

Si02 (wt.%)

60 80 100*An/(An+Ab)

MgO (wt.%)

16-14 Ma trachyandesite-like rock

16-14 Ma picrobasalt

-O- VT-13-9 —A—VT-13-14 —•—VT-13-19

16-14 Ma high-Mg basanite

Cs Rb Ba Th U K Nb Ta La Ce Pb Pr Sr P Zr Hf Sm Ti Y Yb

c o .c

16-14 Ma picrobasalt

14-13 Ma K-Na series

14-13 Ma transitional from Na to K-Na series

16-14 Ma trachybasalt-like rock

16-14 Ma high-Mg basanite

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu

Crustal contamination


P 0.706

1 0.705

0.704 ■


increasing degree of melting

1 1.5 2 2.5

1000/Sr (1/ppm)

§. 4000

MgO (wt.%)

\ A VT-13-9 /v-1-3 > c ^ VT-13-14

/ - oCSy.

Q. 3 re CO

MgO (wt.%)

10 15 20

MgO (wt.%)

10 15 20 MgO (wt.%)

VT-13-21 0

f 3 \VA


/ ___^ A X/J

^VT-13-9 --s' -

MgO (wt.%)

10 15 20

MgO (wt.%)

+ Garnet periditite O Spinel periditite A Vein minerals

Initial 87Sr/86Sr

01 48 % , Opx 25 %

0 5 10 15 20 25

MgO (wt.%)

Gorgona picrites

Hawaii Gorgona komatiites

Archean P Al-undepleted komatiites

DaSgr^OtP Cfeorgona picrites "

UPD) / Hawaii v A Gorgona komatiites

jang Kamar


i Vitim Dariganga O *Kamar Udokan

Transbaikalian domain

Sayan-Mongolian domain ^^

MgO (wt.%)

^ 100 o

32 o o

--Melt 0.04, Gt 6 %

--- Melt 0.04, Gt 12 %

- Melt 0.06, Gt 6 %

Rb Ba Th U K Nb Ta La Ce Pr Sr Zr Hf Sm Ti Y Yb


o o CC

--Melt 0.04, Gt 6 %

---Melt 0.04, Gt 12 %

- Melt 0.06, Gt 6 %

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu

- Melt 0.08, Gt 3.5% --Melt 0.04, Gt 3.5%

Rb Ba Th U K Nb Ta La Ce Pr Sr Zr Hf Sm Ti Y Yb


La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu

Initiation of rifting, uplift and erosion

16-14 Ma, episode 1

Minor contaminated high-Mg basalts-trachybasalts


Grt = 6 %

Contamination of

mantle liquids by crustal material

Partial arrest of ascending liquids in shallow mantle

Solid state diffusion in the lithospheric mantle due to

elevated temperature

Grt = 6-12 =

16-14 Ma, episode 2

16-14 Ma, episode 3

Minor high-Mg transitional (K-Na-K) basanites-basalts

Minor picrobasalts-basalts of K series

Voluminous rift-related moderate-Mg basalts-basaltic andesites of transitional (Na-K-Na) compositions

14-13 Ma, episode 1

Melting at shallow

mantle source » i

Grt- ,'

F = 0.15-0.20

Motion of the lithosphere, rate 2-3 cm yr'1

Uniform partial melts (F = 0.042-0.045)

Elevated degree of melting due to adiabatic heat transfer (F ~ 0.06-0.07)

Decreasing volume of moderate-Mg basalts-trachybasalts of K-Na series

Grt— ~rt+"

Magma generation at deeper mantle level


F = 0.06-0.10

Ambient mantle



Heat transfer in the lithosphere by magmatic conduits

Sub-lithospheric mantle anomaly

Region of decreasing temperature beneath the lithosphere

Adiabatic heat transfer

Melt segregation and separation


• High- and moderate-Mg volcanics erupted in the Bereya center at 16-14 and 14-13 Ma, respectively.

• The initial minor high-Mg rocks were produced by a sub-lithospheric source at depth 115-150 km.

• The subsequent voluminous moderate-Mg rocks were derived from a shallow mantle source.

• No depleted lithospheric material, evidenced by mantle xenoliths, was involved in melting.

• The studied transition indicates a thermal impact on the lithosphere with Tp as high as 1510°C.