Sources and Geodynamics of the Late Cenozoic ... - Springer Link

ISSN 08695911, Petrology, 2010, Vol. 18, No. 3, pp. 278–307. © Pleiades Publishing, Ltd., 2010. Original Russian Text © V.M. Savatenkov, V.V. Yarmolyuk, E.A. Kudryashova, A.M. Kozlovskii, 2010, published in Petrologiya, 2010, Vol. 18, No. 3, pp. 297–327.

Sources and Geodynamics of the Late Cenozoic Volcanism of Central Mongolia: Evidence from IsotopeGeochemical Studies V. M. Savatenkova, V. V. Yarmolyukb, E. A. Kudryashovab, and A. M. Kozlovskiib a

Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences, nab. Makarova 2, St. Petersburg, 199034 Russia email: [email protected] b Institute of Geology of Ore Deposits, Petrography, Mineralogy, and Geochemistry, Russian Academy of Sciences, Staromonetnyi per. 35, Moscow, 119017 Russia email: [email protected] Received July 10, 2009; in final form, October 20, 2009

Abstract—In the Late Cenozoic, the volcanism of the South Khangai Volcanic Region (SKhVR) spanned the Khangai Range and its framing. Geochronological, petrochemical, geochemical, and isotope studies were performed for volcanic rocks of this region, which are represented by highK basic and intermediate rocks of OIB affinity. Initial Sr, Nd, and Pb isotope ratios in the volcanic rocks of the SKhVR are close to those of the volcanic rocks of Pitcairn Island and form trends between PREMA, EMI, and EMII sources. The petrochemical, geochemical, and isotope zoning is unraveled in distribution of the Late Cenozoic asso ciations within SKhVR. Volcanic sequences of the Vodorazdel’nyi graben occupying the watershed part of the Khangai Range and adjacent valley lava flows are located in the central part of the area. The peripheral part is made up of the volcanic associations formed within the Lake Valley and Taryat grabens and the Orkhon– Selenga area. Compositional zoning is characterized by an increase in contents of alkalis, Ti, P, and some other lithophile elements, as well as systematic changes of isotope composition of the rocks from central part toward periphery. Taking into account gravimetric and seismotomographic data marking asthenospheric rise beneath Central Khangai, it was concluded that the studied volcanism is related to mantle plume activity. Revealed composi tional zoning of the volcanic region presumably reflects the plume heterogeneity. The volcanism of the water shed part of the Khangai Range was controlled by plume channel, which was presumably fed by PREMA type lower mantle. The isotopic enrichment of lavas in the peripheral parts of the volcanic region was not related to participation of lithospheric components, but reflects the distribution of compositionally different mantle sources in plume structure. The most probable source of enriched components in the Late Cenozoic rocks of SKhVR was Early Precambrian recycled crustal material, which was isolated from upper mantle con vection after subduction and transported by the ascending mantle jet to the lithosphere base only in the Late Cenozoic. DOI: 10.1134/S0869591110030057

INTRODUCTION Several autonomous volcanic areas united into the Central Asian volcanic province were formed within Central Asia in the Late Cenozoic (Yarmolyuk et al., 1995a). They originated in intracontinental setting far away from lithospheric plate boundaries, owing to withinplate processes accompanied by eruptions of mainly subalkaline and alkaline basic rocks. Geophys ical studies revealed a large rise in lithospheric mantle at the base of the volcanic province. The roof of this rise reached a depth less than 100 km from the Earth’s surface. In addition, local asthenospheric rises to depths less than 50 km were identified beneath some volcanic areas (Zorin et al., 1988; 2004, 2005). These rises were interpreted as mantle diapirs produced by mantle hot field. The similarity in source areas pro vided definite synchronism in the evolution of struc

turally separate volcanic fields (Yarmolyuk et al., 1995a). The geological characteristics of these areas also attest to their relation with deepseated mantle source. These areas are characterized by longterm evolution for more than few tens of million years, steady magmatism, and systematic migrations of vol canic centers over the area presumably marking the lithospheric motions above hot spots (Kovalenko et al., 1997; Kovalenko, 2009; Kudryashova et al., 2006), as well as by relation with uplifted areas in the presentday relief (Yarmolyuk et al., 2008). In the complete form, all these characteristics manifest themselves in the South Khangai volcanic region, whose evolution is traced for last ~160 Ma (Yarmolyuk et al., 1994; 1995a). First systematic data on global seismic tomography of the Earth obtained at the end of 20th century showed that the territory of East and Central Asia is

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underlain by “cold” mantle consisting mainly of oce anic slabs that were subducted on the Pacific side. This fact gave rise to some uncertainty in the understanding the nature of withinplate magmatism of Central Asia, and primarily, put in doubt its relation at least with deepseated mantle plumes. In some publications of that time, including petrological works, withinplate activity was regarded as related to either ascending mantle flows initiated by subduction on the Pacific (Zorin et al., 2006a, 2006b) or Hindustan sides (Demonterova et al., 2007), or to passive rifting and lithospheric delamination (Barry et al., 2003, 2007). The aim of this work was to justify the plume nature of the Central Asian withinplate activity on the basis of isotopegeochemical studies of the Late Cenozoic volcanic products from the South Khangai volcanic region, to estimate of composition of magma sources, and conditions of their generation in sublithospheric mantle of the region. GEOLOGICAL SETTING The South Khangai volcanic region (SKhVR) occupies a significant part of Central Mongolia, span ning the Khangai highland and its northeastern fram ing, as well as adjacent areas of Gobi Altai and Gobi Tien Shan in the south. The volcanic region was formed within territory representing a collage of geo logical structures of different age (Fig. 1): Hercynian structures to the south of the Gobi Altai ranges and Caledonian structures to the north, with intervening Main Mongolian lineament. In this part of Mongolia, the Caledonides include also Early Precambrian (Dzabkhan, Tarbagatai), Riphean (Khangai), Baikalian and late Baikalian (BayanKhongor) ter ranes. The SKhVR evolution is traceable from the termi nal Late Jurassic to the present time (Yarmolyuk et al., 1994, 1995a). During this period, the manifestations of volcanic activity sequentially migrated over South and Central Mongolia, thus forming volcanic area 800 × 400 km in size consisting of loop chain of over lapping volcanic fields of different age (Kudryashova et al., 2006; Yarmolyuk et al., 2007b). This chain is considered as trace of mantle hot spot formed during the passage of continental lithosphere over it (Kudr yashova et al., 2006; Yarmolyuk et al., 2007b). This concept is consistent with paleomagnetitc data, according to which the paleolatitudes of all magmatic areas regardless of their presentday position coincide within confidence intervals with position of modern volcanic area (Kovalenko et al., 1997; Kovalenko, 2009). In the Late Cenozoic, an area of active volcanism was localized within the Lake Valley. Since the end of the Early Miocene, volcanic activity was shifted to the Khangai highland and its northeastern framing (Yarmolyuk et al., 2007a), where produced two volca nic areas, the Khangai and Orkhon–Selenga ones, in PETROLOGY

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the Mongolian and Amur microplates, respectively (Fig. 1). These microplates are separated by boundary, which was presumably formed in the Late Cenozoic (Yarmolyuk et al., 2003) and split territories that were different in scales and evolution of volcanism (Yarmolyuk et al., 1995a). Differences in volcanism as well as spatial isolation of the areas led to the concept of their independent development within adjacent microplates. Our isotopegeochemical studies were focused on volcanism of both the areas for the last ~17 Ma of the SKhVR evolution. The Khangai area involves the Khangai range and its foots. The volcanic activity in this area was mainly confined to three grabens (or volcanic areas): Lake Valley, Vodorazdel’nyi, and Taryat, as well as to the young nearfault valleys of the Khangai range water shed (Fig. 1). Within the range of ~17–12 Ma, the volcanism occurred mainly in the Lake Valley graben (Kudryash ova et al., 2006), where it produced thin lava plateau consisting of trachybasaltic andesites, more rarely, tra chybasalts, basanites, and phonotephrites. This time also marks the onset of volcanic activity within the Vodorazdel’nyi graben (Yarmolyuk et al., 2008) with production of separate thin lava fields of mainly tra chybasaltic composition. In the Late Miocene and Pliocene (~10–4 Ma), volcanism produced thick lava plateau (up to 500 m thick) within the Vodorazdel’nyi graben and eastern part of the Taryat graben (Yarmolyuk et al., 2007a, 2008). The predominant volcanic products of this stage are trachybasaltic and trachybasaltic andesite lavas, with less abundant basalts, basanites, and foid ites (melanephelinites). Within the time range < 3 Ma (Yarmolyuk et al., 2008), volcanism was associated with the growth of the Khangai highland and led to the formation of long val ley flows corresponding in composition to trachyba saltic andesites, trachybasalts, and more rare basani tes. Within the watershed of the Khangai Range, these are the valley flows in the upper reaches of the Baid ragGol (2.7–2.1 Ma), Tuin Gol (1.25 ma), Ongiin Gol (0.25 Ma), and Orkhon Gol (0.22 Ma) rivers, as well as in the Taryat basin (0.75–0.36 Ma). The youngest valley eruptions of basanites (about 6 ky ago) were ejected by Khorgo volcano in the Taryat basin (Devyatkin, 1981). The Orkhon–Selenga area embraces the Orkhon Gol–Selenga interfluve. Within this area, the distribu tion of volcanic fields is controlled by the Ugiinur and Orkhon–Khanui grabens (Fig. 1). The predominant volcanic products are trachybasaltic andesites (shos honites, mugearites) and phonotephrites, with less common trachybasalts and basanites. The oldest volcanic events occurred between ~ 15.5 and 11 Ma (Yarmolyuk et al., 2007a). Within the range of ~ 15.5–14 Ma, a system of shield volcanoes formed in the incipient Ugiinur Graben produced lavas,

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Fig. 1. Scheme of distribution of the Late Cenozoic volcanic areas of the Khangai highland and its northeastern framing in the tectonic structures of the Central Asian fold belt. (1) lava fields; (2⎯6) tectonic structures: (2) Hercynides, (3) Caledonides, (4) terranes of the Riphean crust (Khangai), (5) ter ranes of the Early Precambrian crust (Dzabkhan, Tarbagatai), (6) BaikalianLate Baikalian BayanKhongor zone; (7) faults (a), Main Mongolian lineament (b); (8) inferred walls of the grabens; (9) watershed part of the Khangai Range; (10) inferred boundary between the Mongolian (Mon) and Amur (Am) microplates. Grabens of the South Khangai volcanic area: (I) Lake Valley, (II) Vodorazdel’nyi, (III) Taryat, (IV) OrkhonKhanui, (V) Ugiinur.

which covered the most part of its territory. At the boundary of ~11 Ma (Kudryashova et al., 2009), vol canic centers migrated northward, toward the area of the KhushirUnder settlement located between the Ugiinur and Orkhon–Khanui grabens (Fig. 1). From the Late Miocene [within ~7–4.6 Ma, (Yarmolyuk et al., 2007a)], the area of volcanic activity migrated further north, in the OrkhonGol River basin. It was related to several shield volcanoes, whose lavas erupted along the weakly rugged valleys of the OrkhonGol River tributary, forming a large lava cover, which is distinguished as the upper lava terrace in the presentday relief. The same period (Kudryash ova et al., 2009) was marked by the formation of small Altatyngol lava field extended along the northern slope of the Orkhon–Selenga watershed.

The resumption of the volcanic activity within the graben occurred in the Pliocene. Within the range of ~4–3 Ma (Yarmolyuk et al., 2007a), the valley lava flows were formed in the Orkhon–Khanui graben. The lava flows of the OrkhonGol River filled deep river valley, which at the moment of lava eruptions incised not only lavas of the upper lava terrace, but also their basement. The final stage of the volcanic activity occurred in the Pleistocene (Yarmolyuk et al., 2007a). At that time, the eruption centers formed in the boundary part of the Orkhon–Khanui graben yielded valley lava flows, which propagated beyond the graben along the Selenga river tributary. The comparison of volcanic events in the Khangai and Orkhon–Selenga areas revealed their significant PETROLOGY

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similarity, in spite of the multipulse and longer term development of volcanic processes in the Khangai area. The Middle Miocene lavas in the Ugiinur graben and area of the Khushig–Under settlement with ages between 16 and 11 Ma are correlated with17–12 Ma old eruptions within the Lake Valley and Vodor azdel’nyi grabens of the Khangai area. In the Late Miocene, these centers began to move to the north. The volcanic centers in the Khangai area are localized within the Vodorazdel’nyi and Taryat grabens, while those in the Orkhon Gol River basin are confined to the Orkhon–Selenga area. Between 7.6 and 4.5 Ma, the upper lava terraces consisting of compositionally close rocks (shoshonites and subordinate trachybasalts and basanites) were formed in the eastern part of the Taryat graben and within the Orkhon lava field (Yarmolyuk et al., 2007a, 2008). Almost synchronous pulses of compositionally close rocks occurred in both these areas in the Pleistocene (Yarmolyuk et al., 2007a, 2008). Abovementioned features of resemblance suggest that the Orkhon–Selenga and Khangai areas were formed in similar geodynamic setting during the Late Cenozoic stage of the evolution of the South Khangai volcanic region. METHODS The majorelement composition of the rocks was determined by Xray fluorescence on a CRM25 Quantometer at the Vinogradov Institute of Geochemistry and Analytical Chemistry, Siberian Branch of Russian Academy of Sciences (Irkutsk) using technique (Afonin et al., 1991). Multielement analysis of geological samples was conducted on an ICPMS PlasmaQuad3 VG Elemen tal mass spectrometer at the Institute of Analytical Instrument Making, Russian Academy of Sciences (St. Petersburg). The operation conditions were pub lished in the work (Kovalenko et al., 2003). To correct for drift in the relative sensitivity of the apparatus in sample series (no more than 5–10 samples), standard solutions of heavy metals (Ti, Cr, Ni, Cu, Pb) as well as BCR1 standard were analyzed. REE were cali brated using multielement standard REE solution of Matthew Johnson. The relative error in element deter mination was no worse than 5–10%. In addition, multielement analysis of some samples was carried out on an ICPMS Agilent 7500 ce qua drupole mass spectrometer at the Earth’s Crust Insti tute, Siberian Branch of Russian Academy of Sci ences, following technique (Panteeva et al., 2003). The Nd and Sr isotope analyses were conducted at the Institute of Precambrian Geology and Geochro nology of RAS (St. Petersburg). The preparation of samples for isotope analyses, including chemical decomposition and sequential extraction of elements by ionexchange chromatography, was described in detail in (Savatenkov et al., 2004). The isotope analysis PETROLOGY

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of Nd and Sr was conducted on a Triton multichannel solidstate mass spectrometer. The measurement reproducibility for Rb, Sr, Sm, and Nd calculated on the basis of replicate analyses of BCR1 standard cor responds to ±0.5%. The blank was 0.05 ng for Rb, 0.2 ng for Sr, 0.3 ng for Sm, and 0.8 ng for Nd. The results of the analysis of BCR1 standard (six determinations) were as follows: [Sr] = 336.7 μg/g, [Rb] = 47.46 μg/g, [Sm] = 6.47 μg/g, [Nd] = 28.13 μg/g, 87Rb/86Sr = 0.4062, 87Sr/86Sr = 0.705036 ± 22, 147Sm/144Nd = 0.1380, 143Nd/144Nd = 0.512642 ± 14. The reproducibility of isotope analyses was controlled against the La Jolla and SRM987 standards. During Sr measurements, the 87Sr/86Sr value in SRM987 standard was 0.710241 ± 15 (2σ, 10 measurements), while 143Nd/144Nd in the La Jolla standard was 0.511847 ± 8 (2σ, 12 determina tions). The Sr isotope composition was normalized to 88 Sr/86Sr = 8.37512, while Nd composition, to 146Nd/144Nd = 0.7219. The Nd isotope composition was then adjusted to the table La Jolla standard value 143Nd/144Nd = 0.511860. The chemical extraction of Pb and U from the rocks was conducted using standard technique. The Pb and U isotope analyses were performed on a multi channel Finnigan MAT261 mass spectrometer in a regime of simultaneous record of ion currents from studied elements with measurement error within 0.01% (2σ). For isotope analysis, the Pb and U sam ples were loaded on single rhenium filaments using the silica gel–H3PO4 technique. The total laboratory blanks for Pb and U were no more than 0.1 and 0.01 ng, respectively. The Pb isotope ratios were cor rected for fractionation by double isotope dilution using 204Pb/207Pb tracer (Mel’nikov, 2005). The mea surement errors of 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb isotope ratios determined from a series of parallel analysis of BCR1 standard were no worse than 0.03%, 0.03%, and 0.05%, respectively. COMPOSITION OF THE VOLCANIC ROCKS Classification of the Volcanic Rocks The volcanic rocks of both the areas are character ized by close petrographic and petrochemical compo sition. The representative compositions of their rocks are listed in table. In the TAS diagram (Fig. 2) (Le Bas et al., 1986), the volcanic rocks of both the areas vary in composition from rare foidites (olivine melanephe linites occurring only in the Taryat graben), basanites and basalts (typical of only Vodorazdel’nyi Graben) to predominant trachybasalts and trachybasaltic andes ites, as well as phonotephrites (typical mainly of the Middle Miocene rocks of the Ugiinur and Lake Valley grabens). All the rocks are characterized by a promi nent potassic affinity. Most rocks have K2O + 2% > Na2O and can be classed with potassic trachybasalts and shoshonites (Classification of Magmatic…, 1997). At a general compositional similarity, the lavas of the Orkhon–Selenga area differ in higher SiO2 contents

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Contents of major and trace elements and isotope characteristics of the South Khangai volcanic region

Components

SiO2, wt % TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 L.O.I. Total Mg# Sc, ppm V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U 143Nd/144Nd ±2σ* 87Sr/86Sr ±2σ* εNd(T) εSr(T) 206Pb/204Pb 207Pb/204Pb 208 Pb/204Pb

SKh13/2 48°53.372′N 102°04.521′E TBA 2 53.53 2.42 16.78 8.92 0.10 3.74 6.13 4.62 2.78 0.72 –0.07 99.73 0.45 7.11 87 29 20 38 27 648 15 176 39 383 20 42 5.88 26 5.46 1.89 5.22 0.71 3.40 0.58 1.36 0.16 0.91 0.13 4.39 2.23 2.51 1.91 0.22 0.512592 6 0.704192 9 –0.89 –4.38

SKh13/3 48°51.031′N 102°04.756′E TBA 2 52.68 2.63 15.56 9.44 0.10 4.89 6.09 4.49 2.86 0.81 0.08 99.54 0.51 8.70 80 76 27 81 29 781 15 207 54 405 23 57 6.96 33 6.91 2.21 5.98 0.76 3.52 0.65 1.14 0.13 0.82 0.12 4.73 3.01 2.34 1.77 0.45 0.512519 4 0.704097 5 –2.32 –5.73 17.96 15.47 37.77

Khanui area SKh13/4 SKh13/8 48°49.998′N 48°59.109′N 102°03.911′E 102°09.932′E TBA TBA 2 2 53.13 53.05 2.34 2.39 15.60 15.62 9.63 9.55 0.11 0.11 5.49 5.26 6.46 6.37 4.16 4.32 2.49 2.65 0.64 0.67 –0.28 –0.23 100.05 99.97 0.53 0.52 10.86 9.43 99 89 88 70 31 28 85 69 26 28 651 695 16 16 170 185 39 41 383 406 19 22 45 50 6.17 6.41 27 29 6.12 6.25 1.84 2.05 5.37 5.56 0.72 0.78 3.51 3.42 0.60 0.63 1.41 1.43 0.19 0.19 1.02 1.18 0.14 0.13 4.07 4.65 2.27 2.51 2.49 2.95 1.89 1.79 0.28 0.28 0.512587 0.512576 4 10 0.704139 0.704224 6 5 –0.99 –1.20 –5.13 –3.92 18.12 15.51 38.12

SKh13/11 48°50.350′N 101°23.987′E TBA 3 51.54 2.66 15.03 10.17 0.11 6.12 6.81 4.15 2.75 0.74 –0.30 100.09 0.54 11.86 118 139 34 69 28 767 16 211 50 432 22 53 6.85 31 6.74 2.17 5.93 0.79 3.55 0.60 1.32 0.16 0.87 0.13 4.72 2.67 2.61 2.15 0.66 0.512662 11 0.704139 5 0.50 –5.14

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SKh13/15 48°50.773′N 101°42.904′E TBA 3 52.36 2.04 15.96 10.24 0.12 5.48 7.98 3.83 1.49 0.39 0.00 99.89 0.51 12.30 101 64 31 58 16 458 15 128 24 265 14 32 3.86 20 4.73 1.68 4.60 0.66 3.34 0.65 1.50 0.20 1.19 0.18 3.35 1.48 1.60 0.93 0.41 0.512616 8 0.704260 11 –0.41 –3.42 18.03 15.50 37.89 No. 3

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Table. (Contd.)

Components

SiO2, wt % TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 L.O.I. Total Mg# Sc, ppm V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U 143Nd/144Nd ±2σ* 87Sr/86Sr ±2σ* εNd(T) εSr(T) 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb PETROLOGY

SKh12/7 48°48.686′N 103°09.698′E TBA 3 53.04 2.45 15.26 9.20 0.10 4.68 5.90 4.78 3.27 0.93 0.20 99.60 0.50 5.99 67 71 22 62 33 696 13 217 47 423 25 51 7.29 32 7.31 2.07 5.48 0.74 3.43 0.53 1.05 0.12 0.76 0.10 4.78 2.47 4.45 2.09 0.53 0.512501 9 0.704176 7 –2.64 –4.65

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SKh12/18 48°47.116′N 102°51.289′E TB 3 48.99 2.68 13.68 11.38 0.13 7.58 7.51 4.15 2.84 0.92 0.03 99.87 0.57 12.37 118 182 38 135 38 1064 19 184 64 681 33 67 10.42 46 9.16 2.95 7.63 1.09 4.06 0.73 1.74 0.24 1.34 0.13 4.99 3.88 3.47 3.34 0.34 0.512359 5 0.704226 4 –5.41 –3.92

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Orkhon area SKh12/20 SKh12/30 48°38.731′N 48°36.419′N 103°02.762′E 102°49.820′E TBA TBA 3 3 50.40 52.29 2.54 2.65 14.84 14.97 10.61 9.25 0.12 0.11 5.66 5.60 7.41 6.43 3.76 4.45 2.26 2.72 0.67 0.99 1.58 –0.05 98.27 99.45 0.51 0.55 12.09 7.48 130 79 102 75 37 26 89 64 24 24 1503 861 17 15 162 208 53 42 451 375 21 24 44 55 6.45 7.83 29 36 6.48 7.31 2.12 2.56 5.80 6.27 0.81 0.79 3.68 3.76 0.63 0.57 1.45 1.27 0.16 0.14 0.99 0.78 0.11 0.11 3.47 4.50 2.86 2.45 1.80 2.00 1.67 1.98 0.62 0.37 0.512459 0.512570 3 11 0.704116 0.704682 5 7 –3.48 –1.30 –5.43 2.58 17.99 15.48 37.81

SKh12/32 48°43.123′N 102°36.851′E TBA 3 52.41 2.50 15.16 9.47 0.11 5.01 5.96 4.76 3.15 1.02 0.21 99.54 0.51 7.14 82 68 25 49 35 924 14 216 52 447 27 58 8.18 37 7.82 2.56 6.49 0.81 3.64 0.56 1.09 0.10 0.69 0.08 4.57 2.28 1.69 2.24 0.31 0.512468 16 0.704380 34 –3.26 –1.72

SKh12/34 48°46.073′N 102°37.405′E TBA 3 53.56 3.19 15.41 8.82 0.10 4.80 6.15 4.55 2.25 1.19 –0.27 100.01 0.52 8.14 82 57 25 42 22 1146 17 268 70 352 23 58 9.30 45 10.46 3.34 7.62 1.03 4.49 0.65 1.44 0.15 0.96 0.12 5.75 3.83 1.80 1.73 0.18 0.512625 2 0.704410 3 –0.23 –1.25 18.05 15.49 37.86

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Table. (Contd.)

Components

SiO2, wt % TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 L.O.I. Total Mg# Sc, ppm V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U 143Nd/144Nd ±2σ* 87Sr/86Sr ±2σ* εNd(T) εSr(T) 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb

SKh11/1 47°51.475′N 102°59.558′E PhT 4 48.57 2.85 13.08 11.41 0.12 5.23 6.12 3.56 5.04 1.07 2.64 97.06 0.48 9.84 111 72 28 51 58 1044 18 272 74 728 36 82 10.96 49 9.41 2.64 7.16 0.96 3.85 0.62 1.37 0.18 1.02 0.11 5.52 3.44 1.41 2.64 0.62 0.512203 13 0.705134 6 –8.34 8.81

UgiiNur Lake area SKh11/2 47°53.156′N 103°03.322′E PhT 4 50.27 2.85 13.30 11.07 0.12 5.47 6.32 4.18 4.34 0.94 0.93 98.87 0.49 9.87 118 85 29 51 60 719 16 279 79 598 30 65 8.71 37 7.74 2.22 5.69 0.81 3.52 0.60 1.49 0.17 0.96 0.12 5.59 4.00 1.54 2.48 0.52 0.512279 6 0.704530 9 –6.78 0.01

SKh11/4 47°53.726′N 103°03.147′E TBA 4 49.53 3.10 12.97 11.93 0.13 6.88 6.56 3.86 3.48 0.88 0.51 99.32 0.53 11.16 125 123 38 139 75 635 17 220 71 540 24 57 8.38 37 7.72 2.20 5.89 0.85 3.76 0.61 1.25 0.16 1.09 0.15 4.82 3.92 4.99 1.91 0.57 0.512234 4 0.704372 6 –7.74 –2.62 17.94 15.48 37.80

TsM4/5 48°09.194′N 99°57.600′E Bsn 1 48.18 2.10 15.00 11.56 0.15 7.79 7.17 4.34 3.23 0.77 –0.22 100.29 0.57 15.46 156 157 43 142 40 870 19 263 53 500 44 89 9.65 38 7.03 2.36 6.37 0.94 4.51 0.77 1.83 0.24 1.31 0.16 5.84 3.09 5.55 4.26 1.29 0.512535 3 0.704706 5 –2.01 2.92 17.14 15.45 37.23

Taryat area SKh7/8 48°18.893′N 100°27.657′E TB 2 49.09 2.23 15.06 10.90 0.14 7.15 7.66 4.27 2.74 0.63 0.12 99.86 0.56 14.85 159 119 38 94 30 787 18 213 47 506 28 57 6.84 29 6.03 1.97 5.74 0.81 3.90 0.68 1.66 0.22 1.24 0.16 5.08 2.53 4.71 3.40 5.30 0.512637 10 0.704525 11 –0.02 0.34 17.51 15.51 37.57

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SKh7/9 48°02.312′N 99°59.430′E TBA 2 51.64 2.40 17.03 9.25 0.12 3.55 6.86 4.88 3.28 0.74 0.11 99.75 0.43 12.52 158 21 26 19 39 801 19 254 50 551 30 69 7.59 33 7.10 2.13 5.98 0.88 4.22 0.73 1.46 0.21 1.09 0.19 5.58 2.62 5.68 4.72 0.93 0.512655 9 0.704479 10 0.33 –0.31

No. 3

2010


285

Table. (Contd.)

Components


SKh7/10 48°02.833′N 100°03.485′E TBA 2 51.67 2.34 16.85 9.19 0.12 4.17 7.05 4.75 3.16 0.67 –0.04 99.97 0.47 13.69 151 39 29 29 48 811 19 250 46 552 31 62 7.87 32 6.11 1.93 5.95 0.88 4.14 0.81 1.79 0.25 1.35 0.19 5.67 2.82 6.10 4.10 1.05 0.512647 10 0.704518 10 0.18 0.25 17.83 15.50 37.80 Vol. 18

No. 3

SKh7/12 48°07.111′N 100°01.519′E TBA 2 51.99 2.40 16.52 8.75 0.11 4.03 7.20 4.93 3.17 0.68 0.19 99.79 0.48 12.40 148 81 27 24 34 804 18 232 54 496 28 62 7.99 32 6.48 2.18 5.87 0.86 4.21 0.72 1.57 0.22 1.05 0.17 5.30 3.07 5.05 3.19 0.73 0.512614 8 0.704349 13 –0.46 –2.16

2010

Taryat area SKh14/13 SKh7/3 48°06.395′N 48°19.334′N 100°03.978′E 100°28.060′E TBA Bsn 2 3 52.01 44.62 2.57 2.34 17.15 13.61 8.78 12.10 0.11 0.17 3.69 10.03 6.57 9.13 4.77 3.03 3.53 2.61 0.75 0.75 –0.09 1.33 99.94 98.38 0.45 0.62 8.48 18.09 102 207 15 151 20 48 12 143 32 44 708 847 15 20 199 182 51 60 461 644 24 35 50 74 6.56 8.63 28 37 5.65 7.17 2.02 2.31 5.20 6.51 0.67 0.91 3.05 4.56 0.59 0.83 1.21 1.60 0.14 0.23 0.95 1.41 0.15 0.20 4.82 4.28 2.94 2.90 3.35 2.97 2.65 4.01 0.55 0.98 0.512610 0.512756 5 9 0.704215 0.704040 4 9 –0.54 2.37 –4.03 –6.64 17.82 15.49 37.77

SKh7/5 48°19.392′N 100°27.890′E TBA 3 50.00 2.30 15.65 10.10 0.13 6.03 7.57 4.28 2.60 0.64 0.66 99.29 0.54 12.88 156 104 35 73 31 780 16 192 47 469 23 50 6.44 29 6.42 2.00 5.61 0.80 3.90 0.65 1.27 0.18 1.16 0.15 4.52 2.38 3.18 2.63 0.52 0.512702 9 0.704387 9 1.31 –1.66

SKh14/2 48°24.420′N 100°31.678′E TBA 3 51.11 2.33 15.76 9.85 0.12 5.82 7.01 4.73 2.86 0.65 –0.33 100.24 0.54 10.82 107 90 31 48 38 649 15 194 42 408 20 46 5.94 26 5.94 1.94 5.06 0.75 3.38 0.56 1.27 0.16 0.97 0.12 4.56 2.42 2.56 2.37 0.57 0.512706 6 0.704154 3 1.34 –5.02 17.84 15.49 37.80

286

SAVATENKOV et al.

Table. (Contd.)

Components


Taryat area SKh7/14 SKh14/5 48°24.287′N 48°10.453′N 100°31.566′E 100°23.058′E Ph Bsn 3 3 42.06 44.94 2.89 2.62 14.20 14.45 12.77 12.03 0.17 0.16 6.86 8.34 8.51 7.96 5.54 4.64 3.78 3.23 1.61 0.84 1.41 0.50 98.39 99.20 0.52 0.58 11.14 13.10 169 133 70 81 41 39 83 85 50 49 1527 859 26 19 348 266 100 64 853 644 78 35 155 70 17.52 8.67 69 36 13.21 7.57 3.85 2.39 10.79 6.50 1.51 0.97 6.55 4.33 1.11 0.77 2.22 1.98 0.25 0.24 1.32 1.49 0.15 0.16 7.84 6.15 5.65 4.76 9.52 4.50 6.93 4.90 1.85 1.63 0.512674 0.512723 11 4 0.704056 0.703891 9 4 0.74 1.70 –6.28 –8.75 17.64 15.49 37.60

SKh10/5 46°30.419′N 102°08.063′E Bsn 2 48.74 2.29 14.62 11.96 0.14 7.35 7.17 4.27 2.90 0.66 –0.33 100.09 0.55 11.60 112 106 34 98 39 646 14 150 33 458 22 47 6.21 28 5.38 1.75 5.01 0.70 3.08 0.56 1.18 0.20 1.00 0.10 3.29 1.95 2.67 2.43 0.64 0.512363 15 0.704469 7 –5.37 –0.44

Vodorazdel’nyi area SKh10/7 SKh6/12 46°55.813′N 47°17.127′N 102°29.923′E 99°59.016′E TBA B 2 3 51.40 47.03 1.67 1.91 15.51 16.52 11.52 11.09 0.14 0.15 6.34 8.88 6.91 8.98 4.21 3.52 2.00 1.39 0.46 0.49 –0.36 0.03 100.16 99.96 0.52 0.61 9.09 22.64 75 185 84 209 25 43 67 135 17 20 559 793 13 21 109 175 19 33 272 388 16 25 36 53 4.83 6.67 20 27 4.60 5.81 1.45 2.05 4.07 5.35 0.59 0.81 2.90 4.05 0.51 0.80 1.10 2.23 0.17 0.29 0.90 1.72 0.13 0.25 2.92 3.55 1.17 1.87 3.06 2.00 2.06 2.57 0.41 0.66 0.512329 0.512742 12 8 0.704637 0.704138 16 9 –6.03 2.07 1.94 –5.12 18.28 15.49 38.16 PETROLOGY

Vol. 18

SKh6/10 47°21.100′N 100°13.990′E TB 3 46.78 1.93 15.72 11.17 0.16 9.05 9.01 3.43 1.73 0.46 0.53 99.43 0.62 23.80 184 283 44 152 33 677 22 184 38 395 26 52 6.51 28 5.57 2.04 5.42 0.81 4.35 0.82 2.37 0.30 1.86 0.31 4.35 2.21 5.36 3.08 0.79 0.512663 12 0.703977 10 0.55 –7.51

No. 3

2010


287

Table. (Contd.)

Components


SKh6/14 47°20.017′N 100°06.442′E TB 3 49.87 2.07 15.74 10.40 0.13 6.61 8.15 4.09 2.17 0.52 0.35 99.75 0.56 16.58 141 163 35 78 28 685 19 170 33 459 24 54 6.49 29 6.08 2.03 5.72 0.90 3.88 0.75 2.04 0.22 1.34 0.23 3.88 2.01 3.88 2.57 0.63 0.512448 11 0.704380 9 –3.64 –1.76 17.98 15.51 37.96 Vol. 18

No. 3

SKh6/17 47°26.242′N 100°12.596′E TB 3 49.24 1.96 15.54 11.32 0.15 7.89 8.17 3.78 1.75 0.41 –0.31 100.20 0.58 20.80 161 215 42 111 24 633 21 164 30 376 21 46 5.48 25 5.50 1.87 5.12 0.84 4.19 0.80 2.06 0.27 1.58 0.22 3.68 1.64 3.02 2.47 0.57 0.512594 9 0.704310 8 –0.79 –2.74

2010

Vodorazdel’nyi area SKh6/19 SKh6/21 47°25.973′N 47°29.030′N 100°12.571′E 100°14.394′E TB TB 3 3 48.28 49.81 1.96 2.04 15.22 15.85 11.73 10.98 0.15 0.14 9.02 6.93 8.26 7.63 3.12 3.94 1.87 2.31 0.44 0.55 –0.09 –0.08 100.05 100.18 0.60 0.56 19.67 17.74 168 152 182 145 45 38 126 89 28 27 649 764 20 20 157 187 29 39 430 489 23 26 48 57 5.86 7.23 27 31 5.74 6.72 1.98 2.23 5.47 5.83 0.80 0.90 4.25 4.30 0.77 0.83 2.06 1.92 0.28 0.27 1.48 1.58 0.27 0.22 3.54 4.47 1.66 2.21 2.61 3.83 2.83 2.53 0.73 0.69 0.512486 0.512491 5 16 0.704239 0.704243 4 11 –2.94 –2.81 –3.78 –3.68 18.13 15.52 38.09

SKh6/7 47°18.517′N 100°13.068′E B 3 46.63 1.82 16.17 11.00 0.16 9.44 9.18 3.38 1.26 0.42 0.47 99.45 0.63 23.00 184 240 40 130 21 582 21 162 31 324 22 47 5.75 23 5.35 1.72 5.18 0.80 4.50 0.83 2.42 0.30 1.90 0.28 3.77 1.90 2.11 2.43 0.68 0.512744 11 0.703944 10 2.11 –7.91 18.33 15.52 38.29

SKh6/8 47°18.122′N 100°13.139′E TBA 3 49.88 2.25 17.32 10.29 0.13 4.84 7.57 4.57 2.29 0.58 0.13 99.71 0.48 14.73 149 45 31 32 22 725 20 188 40 419 24 52 6.61 28 6.26 2.20 5.81 0.83 4.29 0.79 1.90 0.21 1.26 0.20 4.09 2.23 1.90 2.28 0.46 0.512557 4 0.704116 5 –1.51 –5.46 17.90 15.50 37.85

288

SAVATENKOV et al.

Table. (Contd.)

Components


SKh6/3 47°20.679′N 100°13.989′E TB 3 50.23 2.04 16.19 10.77 0.14 6.66 7.76 4.11 2.12 0.50 –0.48 100.52 0.55 17.57 162 145 36 78 26 673 19 175 34 390 24 50 5.99 26 5.82 1.92 5.30 0.83 3.99 0.68 2.08 0.25 1.55 0.17 3.90 1.96 2.50 2.04 0.53 0.512525 5 0.704204 5 –2.13 –4.26 17.96 15.49 37.86

Vodorazdel’nyi area SKh6/4 SKh5/2 47°20.661′N 47°07.317′N 100°13.978′E 100°56.537′E TBA TBA 3 4 50.61 51.63 2.09 2.12 16.88 15.42 10.54 10.48 0.14 0.13 5.39 6.58 7.54 7.35 4.39 3.87 2.21 1.61 0.52 0.44 –0.29 0.16 100.30 99.61 0.50 0.55 16.67 16.89 160 156 81 135 33 37 46 93 27 23 676 531 20 19 181 156 35 32 400 338 22 18 48 38 6.01 4.93 26 22 5.68 5.61 1.99 1.80 5.35 5.07 0.73 0.80 4.39 4.56 0.74 0.73 1.81 1.72 0.26 0.26 1.50 1.58 0.19 0.18 4.49 4.07 1.97 1.94 2.86 2.10 2.18 2.09 0.60 0.30 0.512539 0.512538 10 11 0.704377 0.704106 10 11 –1.86 –1.83 –1.81 –5.73

SKh5/10 47°27.541′N 100°51.757′E TB 4 46.59 2.20 14.82 11.40 0.15 8.85 8.39 3.42 2.13 0.66 1.53 98.62 0.61 18.49 181 168 44 127 32 879 22 244 56 546 39 82 9.82 39 7.88 2.76 7.09 1.04 5.62 0.90 2.28 0.27 1.55 0.21 5.93 3.56 3.50 4.76 1.30 0.512615 10 0.704037 11 –0.32 –6.64 18.19 15.49 38.16

Lake Valley area TsM1/4 TsM3/5 46°31.291′N 45°37.573′N 98°57.133′E 101°22.341′E TB TB 4 4 49.48 49.69 2.31 2.18 15.09 14.10 10.82 11.51 0.15 0.12 6.60 7.07 8.33 6.93 4.13 3.50 1.90 2.79 0.50 0.63 0.58 1.44 99.30 98.51 0.55 0.55 16.45 12.17 157 157 167 157 39 38 84 102 27 26 602 945 20 16 206 164 42 34 347 449 23 25 45 60 5.79 8.48 26 38 6.28 7.73 1.96 2.35 5.66 6.43 0.81 0.88 4.11 4.24 0.77 0.66 1.92 1.60 0.23 0.20 1.32 1.20 0.17 0.14 4.67 4.65 2.53 2.15 2.40 3.45 2.71 2.39 0.79 0.65 0.512797 0.512525 6 4 0.704622 0.704446 15 4 3.21 –2.08 1.70 –0.76

PETROLOGY

Vol. 18

No. 3

2010


289

Table. (Contd.) SKh8/3 45°28.566′N 101°15.680′E TBA 4 51.22 2.07 15.33 10.94 0.12 5.78 5.94 4.16 3.32 0.70 0.37 99.57 0.51 8.58 88 98 26 59 21 752 13 141 32 381 21 46 6.41 29 5.84 1.98 4.89 0.67 3.03 0.51 1.20 0.14 0.94 0.10 3.26 2.15 1.52 1.84 0.59 0.512135 7 0.704640 11 –9.70 1.95

Components

SiO2, wt % TiO2 Al2O3 Fe2O3 MnO MgO CaO Na2O K2O P2O5 L.O.I. Total Mg# Sc, ppm V Cr Co Ni Rb Sr Y Zr Nb Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf Ta Pb Th U 143Nd/144Nd ±2σ* 87Sr/86Sr ±2σ* εNd(T) εSr(T) 206Pb/204Pb 207Pb/204Pb 208 Pb/204Pb

Lake Valley area SKh8/4 SKh8/7 45°29.951′N 45°35.594′N 101°15.775′E 101°11.240′E TBA PhT 4 4 51.73 49.27 1.93 2.62 15.69 14.83 10.54 10.16 0.11 0.13 5.22 6.23 5.53 7.09 4.40 4.42 3.67 3.41 0.72 0.81 0.43 1.00 99.54 98.97 0.49 0.55 8.39 9.06 82 69 82 72 26 27 59 62 24 31 799 837 13 19 171 239 38 71 405 487 24 32 55 71 7.40 8.77 32 40 6.30 8.64 2.07 2.45 5.52 7.02 0.70 1.05 3.16 4.75 0.50 0.77 1.11 2.01 0.17 0.23 0.79 1.28 0.10 0.19 4.01 6.63 2.41 4.87 1.97 3.62 2.55 3.89 0.53 0.90 0.512115 0.512462 5 4 0.704526 0.704276 5 4 –10.07 –3.31 0.36 –3.25 17.20 17.48 15.44 15.45 37.34 37.59

TsM2/5 45°55.807′N 101°14.126′E Bsn 4 47.61 2.39 13.91 11.83 0.14 7.25 7.34 4.32 3.17 0.94 1.08 98.89 0.55 12.09 150 117 38 64 21 1076 19 194 50 710 40 88 11.42 48 9.01 2.78 7.34 1.02 4.69 0.76 1.68 0.23 1.18 0.16 4.60 2.89 2.58 3.23 0.86 0.512197 4 0.704463 2 –8.44 –0.44 17.43 15.47 37.51

Note: Rocks: (B) basalt, (Bsn) basanite, (F) foidite (olivine melanephelinite), (TB) trachybasalt, (TBA) trachybasaltic andesite, (PhT) phonoteph rite. Volcanic stages: (1) Holocene, (2) Pleistocene, (3) Pliocene–Late Miocene, (4) Middle Miocene, * – absolute error in last digit. Geo graphic coordinates (WGS 84). PETROLOGY

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2010

290

SAVATENKOV et al. Na2O + K2O, wt % 11

9

F

7

1 2 3 4

PhT

Bsn TBA TB

5 B 3 42

44

46

48

50

52

54 56 SiO2, wt %

Fig. 2. Diagram Na2O + K2O⎯SiO2 (Le Bas et al., 1986) for the Late Cenozoic (< 17 Ma) volcanic rocks of the Khangai highland and its northeastern framing. (1⎯4) rocks: (1) OrkhonSelenga area; (2⎯4) Khangai area: (2) Taryat area, (3) Vodorazdel’nayay Group, (4) Lake Valley area. (B) basalts, (Bsn) basanites, (TB) trachybasalts, (TBA) trachybasaltic andesites, (F) foidites (olivine melanephelinites), (PhT) phonotephrites.

(mainly 48–53 wt %) and K2O + Na2O, and, corre spondingly, are mainly represented by trachybasaltic andesites and phonotephrites (Fig. 2). Petrographic Characteristics As a rule, lavas of both the areas are represented by dense gray to black rocks. The widely developed porous varieties with vesicular or amygdaloidal struc ture typically compose the upper parts of lava flows. The petrographic appearance of the lavas is fairly uni form. These are mainly porphyritic rocks with phe nocrysts represented by variable proportions of ubiq uitous olivine, as well as clinopyroxene, plagioclase, and more rare Kfeldspar and nepheline. The typo morphic petrographic features corresponding to the petrochemical composition of the rocks are not observed. The exception is foidite (olivine melanephe linite) occurring only in the Taryat graben. This rock, in addition to olivine, contains K–Na feldspar and rarer nepheline. Olivine phenocrysts are typically 0.5–1.5 mm across, sometimes more than 2 mm. They are typically partly replaced by iddingsite aggregates or pale green fibrous and scaled minerals of serpentine and chlorite groups. Clinopyroxene forms prismatic brown crystals corresponding to alkaline Tiaugite (up to 1–1.5 mm, rarely > 2 mm). Plagioclase is commonly observed as large (up to 1–2 mm and more) tabular crystals with zoned pattern. Scarce Kfeldspar composes weakly pelitized grains up to 1 mm in size. Nepheline occurs

as single rectangular grains up to 1.7 mm often replaced by colorless mica. The groundmass is usually well crystallized, mainly finegrained, and consists of microlites, laths or xenomorphic grains of plagioclase (up to 0.5 mm, occasionally up to 1 mm), fine grains of olivine and clinopyroxene, as well as elongated crystals of black, more rarely reddish brown (hema tite) ore mineral. In addition, alkali feldspar some times forms individual grains or thin rims around pla gioclase; there are also brown (alkali and more acid) or colorless volcanic glass replaced by chlorite or fine yel lowish and greenish droplike aggregates supposedly representing a mixture of clay minerals with sericite, chlorite and serpentine group minerals. Accessory mineral is apatite. Chemical Composition As was mentioned above, the Late Cenozoic volca nic rocks of both the areas are represented by highK rocks with moderate Mg number (Mg# from 0.43 to 0.65 at average 0.55). In spite of this general similarity, the rocks of the Khangai and Orkhon–Selenga areas systematically differ in terms of other parameters. In particular, the rocks of the Orkhon–Selenga area in general are characterized by the higher contents of SiO2, K2O, TiO2, P2O5, and lower contents of Al2O3 and MnO (Fig. 3), owing to which the rocks of both the areas define separate compositional trends, which were presumably controlled by fractional crystalliza PETROLOGY

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2010

SOURCES AND GEODYNAMICS OF THE LATE CENOZOIC VOLCANISM Al2O3, wt % 20

(a)

291

Mg# 0.8

18

0.7

16

0.6

14

0.5

12

0.4

(b)

0.3 TiO2, wt %

10 K2O, wt % 6

(c)

4

(d)

5 3

4 3 2

2

HighK

1

ModerateK

0

1

MnO, wt % 0.20

(e)

P2O5, wt % 1.8

(f)

1.5 0.16

1.2 0.9

0.12

0.6 0.3

0.08 40

42

44

46

48

50

52

56 54 SiO2, wt %

0 40

42

44

46

48

50

52

56 54 SiO2, wt %

Fig. 3. Variations of major elements and Mg# versus SiO2 in the Late Cenozoic rocks of the Khangai highland and its northeastern framing. Symbols are shown in Fig. 2.

tion of parental magmas and different melting degree of their protoliths. In addition to these differences, significant petro chemical variations are also observed between rocks of individual volcanic fields. The lavas of the Vodor azdel’nyi graben and lavas of valley flows of the Khan gai Range watershed are close in composition and dif fer from the rocks of other areas by the lowest contents of K2O, TiO2, and P2O5, the highest Аl2O3, and Mg number (Mg#av = 0.57). Therefore, further these PETROLOGY

Vol. 18

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2010

rocks are considered as a single compositional group (Vodorazdel’naya). The rocks of the Lake Valley gra ben differ from volcanic rocks of other grabens of the Khangai area in elevated contents of TiO2, P2O5, and lowered contents of Al2O3, MnO, being in terms of these parameters close to the rocks of the Orkhon– Selenga area (Fig. 3). The rocks of the Taryat graben show a wide range in chemical composition. Foidites (olivine melanephelinites) with low SiO2 content (41.9–43.8 wt %) and lowered Mg number (Mg#av =

292

SAVATENKOV et al. Zr, ppm 380 (а) 1 280 2

180

80 Th, ppm 8 (b)

1

6

4 2 2

0

30

60

90

120 Nb, ppm

Fig. 4. Variations of trace elements versus Nb in the Late Cenozoic rocks of the Khangai highland and its northeastern framing. Shown are compositional trends for the Khangai (1) and OrkhonSelenga (2) areas. Symbols are shown in Fig. 2.

0.52), and the highest values of K2O, MnO, TiO2, and P2O5 contents were found only in this graben among the Late Pliocene rocks (Fig. 3). TraceElement Composition The Late Cenozoic volcanic rocks of the areas show wide variations in trace elements. The incompat ible trace elements have positive correlations with

each other (Fig. 4). The trends of trace elements show regional differences. For instance, the rocks of the Orkhon–Selenga area have the lower Th contents, pointing to difference in source composition for indi vidual volcanic areas. At the same time, the presence of such correlations indicates a leading role of crystal lization differentiation and/or melting degree of initial protolith in their formation. However, the fact that most trace elements negatively correlate with SiO2 PETROLOGY

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293

Sr, ppm 1900

Ba, ppm 900 (a)

(b)

1600

700 1300 1000

500

700 300 400 100 Nb, ppm 120 (c) 100

100 Y, ppm 30

(d)

25

80

20

60 15

40

10

20 0 La, ppm 90

5 Yb, ppm 2.5 (f)

(e) 2.0

60

1.5 1.0

30 0.5 0 Th, ppm 8

0 U, ppm 2.5 (g)

(h) 2.0

6

1.5 4 1.0 2 0 40

0.5 42

44

46

48

50

52

54 56 SiO2, wt %

0 40

42

44

46

48

50

52

54 56 SiO2, wt %

Fig. 5. Variations of trace elements versus SiO2 in the Late Cenozoic rocks of the Khangai highland and its northeastern framing. Symbols are shown in Fig. 2. PETROLOGY

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294

SAVATENKOV et al. Rock/Primitive mantle 103 (a)

102 1 2 3 4 5 101

100 Rb Ba Th U NbTa K La Ce Pr Sr P Nd ZrHf SmEu Ti GdTbDy Y HoErTmYb Lu 102 (b)

101

100 La

Ce

Pr

Nd

Sm

Eu

Gd

Tb

Dy

Ho

Er

Tm

Yb

Lu

Fig. 6. Average traceelement distribution patterns for the Late Cenozoic rocks from different areas of SKhVR. (1⎯5) primitive mantlenormalized average compositions of the rocks (Sun and McDonough, 1989): (1) OrkhonSelenga area; (2⎯4) Khangai area: (2) Taryat area, (3) Vodorazdel’naya group; (4) Lake Valley area; (5) oceanisland basalts (Sun and McDon ough, 1989).

(Fig. 5) suggests that compositional variations of melts were mainly related to the differences in melting degree of protolith. Together with common regularities in trace ele ment distribution, the rocks of the considered volcanic areas have some differences. Figure 6 demonstrates the primitive mantlenormalized (Sun and McDon

ough, 1989) patterns for the average trace element compositions of the Late Cenozoic rocks from various areas of SKhVR. In general, all rocks are characterized by enrichment in highly incompatible trace elements relative to primitive mantle; however, their contents, on average, are lower than those in the oceanisland basalts (Sun and McDonough, 1989). The trace ele PETROLOGY

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295

50 Nb/Th

(а)

40

30

20

10

OIB UC

LC

La/Th

0 2

4

6

8

10

12

16

14

250 (b)

Nb/U 200

150

100

MORB OIB

50 UC 0

20

La/U

LC 40

60

80

100

120

Fig. 7. Diagrams Nb/Th⎯La/Th (a) and Nb/U⎯La/U (b). Reference compositions of magmatic sources (light gray fields) after (Taylor and McLennan, 1988; Sun and McDonough, 1989): OIB are oceanisland basalts, MORB is midocean ridge basalts, UC is upper crust, LC is lower crust. Symbols are shown in Fig. 2.

ment distribution patterns in the volcanics of both the areas are close to OIB type. However, relative to the latters, they are enriched in Ba, K, Sr, P, and depleted in Th, U, Hf, and REE (especially in HREE). Note also some differences in trace element distri bution between different volcanic areas. First of all, the volcanics of the Vodorazdel’naya Group have more gentle REE patterns and elevated HREE relative to rocks from other fields (Fig. 6). The rocks of the Orkhon–Selenga area are depleted in Th and U as PETROLOGY

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2010

compared to the rocks of the Khangai area. This spe cifics is best expressed in the variation diagrams for pair ratios containing Th and U, where compositions of the two areas define practically nonoverlapping fields (Fig. 7). Noteworthy is that the compositions of the volcanic rocks of SKhVR, especially of the Khan gai area, in Fig. 7a are plotted closely to the composi tion of oceanisland basalts (OIB) (Sun and McDon ough, 1989), which are derivatives of enriched mag mas, and do not form trends toward the compositions

296

SAVATENKOV et al. 30

εNd(T)

(a) 1 2 3 4 5 6 7 8

15 PREMA

(б) 0 EMII

6

3

εSr(T)

EMI

–15 –40

–20

0

20

40

60

εNd(T)

80 (b)

PREMA

0 EMII –3

–6

–9 εSr(T)

EMI –12 –15

–10

–5

0

5

10

15

20

Fig. 8. Diagram εND(T)⎯εSr(T). Mantle sources (Zindler and Hart, 1986): PREMA is moderately depleted mantle; EMI and EMII are enriched mantles. Light gray field in Fig. 8a denotes the composition of Late Cenozoic rocks of the SKhVR. Rocks: (1⎯2) OrkhonSelenga area: (1) our data, (2) data (Barry et al., 2003); (3⎯4) Taryat area; (3) our data, (4) data of (Barry et al., 2003); (5) Vodorazdel’naya Group; (6) Lake Valley Group. (7) lower crustal xenoliths of Mongolia (Barry et al., 2003); (8) mantle xenoliths from the Late Cenozoic vol canic rocks of Central Asia (Kovalenko et al., 1990; Ionov et al., 2005).

of upper (UC) and lower (LC) continental crust (Tay lor and McLennan, 1988). Isotope Composition Isotope data (Sr, Nd, Pb) on the Late Cenozoic volcanic rocks of the SKhVR are shown in table and Figs. 8–9. Data on volcanic rocks from (Barry et al., 2003) as well as on mantle (Glebovitsky et al., 2007;

Kovalenko et al., 1990; Ionov et al., 2005) and lower crustal (Barry et al., 2003) xenoliths found in the vol canic fields of Central Asia are shown in these plots for comparison. As was noted previously (Yarmolyuk et al., 1995b; Barry et al., 2003), the isotope compositions of the Late Cenozoic basalts are plotted between three com ponents (Fig. 8) corresponding to characteristics of PETROLOGY

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SOURCES AND GEODYNAMICS OF THE LATE CENOZOIC VOLCANISM 15.7

μ = 8.5

207

Pb/204Pb

(a) EMII

4.55

UC

Ga

μ = 8.4

15.6

297

Pitcairn μ = 8.3 PREMA

15.5 EMI μ = 8.1

DM O rk h Se l

15.4

15.3

μ = 8.0 LC

39.0 208Pb/204Pb

(b)

Pitcairn EMII EMI

38.5

LC

38.0

37.5

D

M

37.0

36.5 16

C U

17

206Pb/204Pb

18

19

Fig. 9. Diagrams 207Pb/204P b ⎯ 206Pb/204Pb (a) and 208Pb/204Pb ⎯ 206Pb/204Pb (b). Trend of the rocks of the OrkhonSelenga area (OrkhSel) corresponding to the isochron that intersects DM curve at age of 2.5 Ga. Model trends of Pb isotopic composition evolution: (DM) depleted mantle (Stracke et al., 2003a); UC, LC, are upper and lower crust (Zartman and Haines, 1988). 4.55Gaold Geochron is constructed, with (μ) values (238U/204Pb) for one stage model of Pb isotope evolution (Faure, 1989). Model mantle sources (Hofmann, 1997; Zindler and Hart, 1986; Wilson, 1989): PREMA is moderately depleted mantle; EMI and EMII are enriched mantle. Light gray field is the field of midocean ridge and oceanisland basalts according to (Hofmann, 1997). Symbols are shown in Fig. 8. Data on mantle xenoliths from the Late Cenozoic rocks of the Vitim lava plateau (Glebovitsky et al., 2007).

PREMA, EMI, and EMII (Zindler and Hart, 1986). Note that the field of isotope compositions of the vol canic rocks practically is not overlapped with field of mantle xenoliths (Fig. 8a), which characterize lithos pheric depleted mantle (Kovalenko et al., 1990; Ionov et al., 2005). The compositions of rocks from different areas form partly overlapping fields (Fig. 8b), thus PETROLOGY

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reflecting different contribution of the indicated com ponents. The volcanic rocks from various areas reveal signif icant differences in the initial Nd and Sr isotope ratios (Fig. 8b, table). In particular, the oldest lavas of the Lake Valley demonstrate a significant contribution of EMItype source. In the volcanic rocks of the Vodor

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azdel’naya group, the PREMAtype source is domi nant, with a lesser contribution from EMI source. The contribution of EMII source is deciphered in the vol canic rocks of the youngest Taryat graben. Two trends can be identified in the distribution of these composi tions. One trend is distinctly oriented toward EMII source, while other strikes to EMI, with significant contribution of EMII component. The Nd and Sr isotope compositions of the rocks of the Orkhon–Selenga volcanic area also reflect the sequential change of sources with time (Fig. 8b, table). The oldest volcanic rocks of the Ugiinur graben have the most significant contribution from EMI source. The younger rocks of the Orkhon–Khanui graben have lesser contribution of this component. At the same time, in addition to EMI, these rocks contain insignificant amount EMII. The PREMAtype source is the second mixing components for all the basalts. The comparison of Pb isotope compositions of the Late Cenozoic volcanic rocks of SKhVR with those of inferred mantle magma sources (Hofmann, 1997; Zindler and Hart, 1986; Wilson, 1989) showed that the studied rocks are plotted in the field of basalts of mid ocean ridges and islands (Hofmann, 1997). The data points of the Khangai area in the 207Pb/204Pb– 206Pb/204Pb diagram (Fig. 9a) define an array extended along depleted mantle evolution curve (Stracke et al., 2003) from area close to PREMA (Zindler and Hart, 1986; Wilson, 1989) toward enriched mantle source EMI (Hofmann, 1997; Zindler and Hart, 1986; Wil son, 1989). The rocks of the Orkhon–Selenga area in the diagram 207Pb/204Pb–206Pb/204Pb (Fig. 9a) form a linear trend, which is oriented obliquely to a general array of data points of the rocks of the Khangai area. The corresponding trend also takes origin in the PREMA area and extends toward the presentday iso tope composition of the ancient lower crust according to model of plumbotectonics (Zartman and Haines, 1988). In addition, this plot demonstrates that contri bution of EMII source practically did not affect the composition of the rocks of SKhVR. In the diagram 208Pb/204Pb–206Pb/204Pb (Fig. 9b), the data points of the considered rocks define a common linear trend, which slightly deviates from the depleted mantle evolution curve (Stracke et al., 2003a) toward component depleted in U and enriched in Th. Presented plots (Fig. 9) revealed differences in Pb isotope compositions between rocks of studied volca nic areas. These differences in general are consistent with Sr–Nd isotope data and determined by variable contribution of moderately depleted (PREMA) and enriched (EMI) components in the melt formation. It is seen in Fig. 9a that the rocks of the Vodorazdel’naya group of the Khangai area and Orkhon–Selenga area were least affected by the enriched component EMI. Unlike the rocks of the Vodorazdel’naya group and Orkhon–Selenga area, the Pb isotope compositions of the rocks of the Lake Valley and Taryat grabens in the diagram (Fig. 9a) are significantly shifted to the left

from isochron, which indicates notable contribution of enriched EMItype component. In the diagram 208Pb/204Pb–206Pb/204Pb (Fig. 9b), no differences are observed between the rocks of the Khangai and Orkhon–Selenga areas. The evolution curve of ancient lower crust according to the model of plumbotectonics (Zartman and Haines, 1988) lies sig nificantly higher; hence, this deviation is not related with lower crustal contamination. COMPOSITIONAL VATIATIONS OF THE VOLCANIC ROCKS IN TIME AND SPACE The compositional variations of the Late Cenozoic rocks of the SKhVR are determined by their geograph ical position. The rock complexes of different age within individual volcanic fields show lesser differ ences, while the strongest differences are observed (Figs. 2, 3) between the rocks of the Vodorazdel’naya group that occupy the central part of the Khangai Highland and Orkhon–Selenga areas. The rocks from other areas are characterized by intermediate charac teristics, though volcanics of the Lake Valley are closer to the rocks of the Orkhon–Selenga area than lavas of the Taryat graben. Thus, the Late Cenozoic volcanic rocks of SKhVR show zoned distribution. The Central part of the area is occupied by the rocks of the Vodor azdel’naya Group with the most primitive composi tion, whereas rock associations of more differentiated composition are developed in its framing. The differences between central and peripheral zones of the area can be presumably related to the het erogeneous composition of melting protolith, as well as to the conditions of its melting (depth, tempera ture). Conditions of melting, in particular, changes in modal composition of initial protolith depending on pressure could cause aforementioned differences in trace element patterns between rocks of the Vodor azdel’naya Group and other areas (Fig. 6). Thus, the spatial chemical and geochemical variations of the rocks are related to either different depth or tempera ture of mantle melting. Similar explanation was pro posed to explain the differences in chemical composi tions of the Late Cenozoic volcanic rocks of the Khub sugul area (Demonterova et al., 2007). Revealed zoning is not determined by tectonic structure of the region. In particular, the rocks of the Lake Valley are located within the Dzabkhan micro continent with preRiphean crust and in block of the Caledonian crust, which frames the microcontinent from the south and belongs to the Mongolian micro plate in the presentday structure of the region (Fig. 1). However, in terms of chemistry, the lavas of this group are closer to the rocks of the Orkhon– Selenga area, which were formed mainly in the block with Caledonian juvenile crust and, in the modern structure, belong to the Amur plate (Fig. 1). The lavas of the Vodorazdel’naya Group, as well as most lavas of PETROLOGY

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the Taryat Graben are situated within terrane with Preriphean crust (Fig. 1), but have different petro chemical characteristics. This leads us to conclude that spatial compositional variations of the Late Cen ozoic rocks of the SKhVR were not controlled by lithospheric structure, but determined by sublithos pheric sources. Lowered (relative to LREE) HREE contents in the rocks of SKhVR can be related to the buffer effect of garnet, which characterizes the greater depths of melt initiation. Melting within garnet stability field also can explain lowered Al2O3 content in the parental melt (Kogiso et al., 2003). Correspondingly, the lower Al2O3 and HREE contents in the rocks from periphery of the Khangai highland (Fig. 6) as compared to the rocks of the Vodorazdel’naya Group could be related either with greater depth of their formation or with lower melting degree, i.e., with lower mantle melting temperatures. Such a zoning in magma generation regimes could be interpreted in terms of relation of magmatism with hot spots (mantle plume, Yarmolyuk et al., 1995a), with zoned distribution of temperature regime from plume center to its periphery. It should be noted that link of the Late Cenozoic Khangai magmatism with plume activity is confirmed by geophysical data (Zorin et al., 1988, 2004), which demonstrated that asthenosphere beneath Central Asia is risen to a depth less than 100 km. It was also established that the areas of development of the youngest volcanism, in particular, the Khangai high land, are situated above local asthenospheric rises reaching depth less than 50 km from the Earth’s sur face (Zorin et al., 1988, 2004; 2005). These conclu sions were later confirmed by seismotomographic studies (Mordvinova et al., 2007, 2008), which revealed intense lowvelocity anomaly beneath the central part of Khangai. The anomaly is traced to a depth of 200 km, then subsides to the north, where, within depths of 450–600 km, it has the narrowest width and can be considered as plume channel. Presented geophysical data suggest that the Khan gai area is underlain by a mushroomshaped jet of low velocity mantle ascending at least from the upper– lower mantle boundary to the lithosphere bottom. This jet causes the deformation (melting) of the latter, with formation of a diapiric asthenospheric dome above the mantle channel and its spreading in the form of extended lens of lowdensity mantle, which imparts a mushroom shape to the entire mantle jet. The most uplifted part of the mantle jet is located beneath Khan gai and submerges away from it. In the framework of this concept on structure of feeding mantle column, the volcanic occurrences of the Vodorazdel’naya group are spatially associated with central parts of the mantle plume projection onto the Earth’s surface. In this case, the lower fraction of garnet in residue could be related to higher melting temperatures, which are typical of plume channel, or to melting at shallower depths, above the garnet stabil PETROLOGY

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ity field, in the head of ascending mantle plume. Shal low depths or high melting temperature are also con firmed by more mafic composition of the volcanic rocks of the Vodorazdel’naya group. Estimation of Source Compositions of the Late Cenozoic Volcanism of the SKhVR The Nd, Sr, and Pb isotope compositions of the Late Cenozoic rocks of the SKhVR indicate the con tribution of different mantle sources in their forma tion. Observed variations in isotope characteristics of the studied rocks are not related to contamination by lithospheric mantle. This conclusion can be made from analysis of diagrams εNd(T)–εSr(T) (Fig. 8a), 207 Pb/204Pb–206Pb/204Pb (Fig. 9a), 208Pb/204Pb– 206 Pb/204Pb (Fig. 9b), where compositional field of xenoliths (Glebovitsky et al., 2007; Kovalenko et al., 1990; Ionov et al., 2005) that represent depleted litho spheric mantle is clustered separately from volcanic rocks and shows no correlation dependences with them. The distribution of data points of the lower crustal xenoliths (Barry et al., 2003) in the diagram εNd(T)–εSr(T) (Fig. 8a) and 207Pb/204Pb–206Pb/204Pb (Fig. 9a) suggests the crustal contribution as reason for isotope heterogeneity of the studied volcanic rocks. However, as was mentioned above, the traceelement composition of the volcanic rocks is not correlated with crustal composition (Fig. 7). The differences between compositions of the volcanic rocks and crustal sources are also well observed in the diagrams of Sr and Pb isotope correlations (Fig. 10a), which indicates that variations of isotope composition of the volcanics are not correlated with composition of regional continental crust represented by crustal xeno liths. Thus, the Late Cenozoic rocks of the SKhVR were not contaminated by crustal or lithospheric man tle components. The observed compositional varia tions of magmas were primarily defined by composi tional heterogeneity of their deep mantle sources. As was shown above (Fig. 9), the isotope composi tions of the Late Cenozoic volcanic rocks of the Khan gai area were presumably derived from two main sources, which were close in their parameters to PREMA and EM1type mantle. The same conclusion follows from diagram 143Nd/144Nd–206Pb/204Pb (Fig. 10b), where compositions of the area are con fined to the fields of basalts of ocean islands such as Pitcairn and Kergulen oceanic plateau, which were formed with participation of the EMItype enriched mantle (Stracke et al., 2003b). In the diagram 87Sr/86Sr–206Pb/204Pb (Fig. 10a), they also coincide with the compositional field of the rocks of Pitcairn Island (Stracke et al., 2003b). Such a similarity sug gests that source of the rocks of the Khangai area was isotopically close to sources of withinplate magmas in the Pacific and Indian oceans, which contained EMI component (Stracke et al., 2003b).

300 0.709

SAVATENKOV et al. 87

Sr/86Sr

(a)

0.708

EMII

0.707 0.706

EMI Pitcairn

0.705 0.704 0.703

HIMU

MORB 0.702 0.701 0.5135

143Nd/144Nd

(b)

0.5133

MORB

0.5131 0.5129

Pitcairn

HIMU Kergulen

0.5127 0.5125

EMII

0.5123 EMI

0.5121 0.5119 16

17

18

206Pb/204Pb

19

20

21

22

Fig. 10. Diagrams 87Sr/86Sr206Pb/204Pb (a) and 143Nd/144Nd206Pb/204Pb (b). The light gray field is the field of midocean ridge and oceanisland basalts with end members according to (Stracke et al., 2003b): HIMU denotes basalts with ele vated 238U/204Pb ratio; MORB is the midocean ridge basalts; EMI and EMII are enriched mantle. Symbols are shown in Fig. 8.

What is the nature of EMItype source? Based on Nd and Sr isotope composition (Fig. 8), the most sig nificant contribution of this source was found in the rocks of the Lake Valley graben, whereas its influence in the rocks of the Vodorazdel’naya group was less sig nificant; the radiogenic Srrich source was docu mented for the rocks of the Taryat graben. Data on Pb isotope composition (Fig. 9a) indicate that EMItype source made the most significant contribution in the rocks of the Lake Valley and Taryat graben, data points of which are most close to its model composition. While characterizing the compositional parameters of this enriched component, we noted a negative cor relation of 206Pb/204Pb with content of incompatible

Zr, La, Pb, Th, U, as well as with εSr(Т) in the Late Cenozoic rocks of the SKhVR (Fig. 11). All these facts testify that enriched component in the source of SKhVR melts was characterized by elevated contents of incompatible elements and values of εSr(Т) ≥ 5, as well as εNd(T) ≤ –10, 207Pb/204Pb ≤ 15.45, 206Pb/204Pb ≤ 17, 208Pb/204Pb ≤ 37.1. The appearance of source with such isotope char acteristics can be explained in the frameworks of model of polystage evolution of Pb isotope composi tion. This model was proposed to explain the nature of EMI source in the lavas of Pitcairn island (Eisele et al., 2002). Its origin was considered to be related to the recycled lithospheric material, in which proportions of U and Pb contents changed at early stages of its exist ence (Fig. 12). We believe that this source of SKhVR magma was extracted as juvenile crust from primitive mantle. In accordance to the model, in order to plot in the continuation of the revealed trend of Pb isotope compositions for the rocks of the Khangai area, their source should be subjected to following modifications. At the initial stage, this source had the U/Pb (μ) ratio higher than that in primitive mantle (PM), as well as enriched geochemical characeristics, including ele vated contents of LREE, which is required to obtain presentday negative εNd(Т). During following stage, its subductionrelated subsidence in mantle at higher temperature and higher pressure levels was accompa nied by partial and nonproportional loss of U, Th, and Pb from its rocks (Bolhar et al., 2007). This led to the decrease of U/Pb ratio and increase of Th/U ratio, which retarded the evolution of 206Pb/204Pb and 207Pb/204Pb ratios and growth of 208Pb/204Pb ratio. Finally, this source was buried in mantle and by the Late Cenozoic could acquire required isotope charac teristics, which determined its position in the left con tinuation of the trend of isotope composition of SKhVR in the plot of Pb isotope compositions (Fig. 9). The time of U loss from system due to its subsid ence is difficult to determine because of requirement to make some assumptions about formation time of enriched source (separation from mantle), initial U/Pb and Th/U ratios in the system, loss of compo nents of the U–Th–Pb system, and content of Pb at the moment of magma generation. Estimates based on such concepts are very speculative and primarily imply the variations in U/Pb and Th/U ratios. As to volcanic rocks of SKhVR, the age estimates and nature of enriched component are based on the following premises. The low εNd(Т) suggests the ancient age of enriched source as independent geochemical system. The values of εNd(Т) ≤ –10 in the end member of the Lake Valley lavas suggest the model age no less than 2.5 Ga (Yarmolyuk et al., 1988). Addi tional information on the age of SKhVR source can be obtained from analysis of lead isotope system. In the plots (Fig. 9a), the data points of the rocks of the Orkhon–Selenga area define a linear trend, which PETROLOGY

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70 Zr, ppm

(а)

La, ppm

(b)

U, ppm

(d)

εSr(T)

(f)

60

300

50 250 40 200 30 150

20

100 6

10 2.0 Th, ppm

(c)

5

1.5

4 1.0 3 0.5 2 1

0

8

20 Pb, ppm

(e)

6

10

4

0

2

–10

0

–20 16.9

17.2

17.5

17.8

18.1

18.4 18.7 206Pb/204Pb

16.9

17.2

17.5

17.8

Fig. 11. Diagram of variations of trace elements (Zr, La, Pb, Th, U) and εSr(T) versus 206Pb/204Pb. Trends of compositional variations are shown. Symbols are shown in Fig. 8.

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18.4 18.7 206Pb/204Pb

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Pb/204Pb

15.4

Subducted lithosphere μ = 5.5, κ = 4.4

3

Presentday composition of enriched component in the magma source of SKhVR

2.5 Ga

15.0 2

Enriched crust μ = 12.5, κ = 3.4

14.6 2.9 Ga 1 14.2 12

Depleted mantle μ = 8.2, κ = 3.8 14

206Pb/204Pb

16

18

20

Fig. 12. Model of Pb isotope composition evolution in enriched magma source of SKhVR in the diagrams 207Pb/204P b⎯ 206Pb/204Pb. Arrows show isotope variations of source at different stages of its evolution. Gray dashed lines are variation trends for isotope res ervoirs, without disturbance of their composition. Numbers in circles are stages of evolution of source isotope composition of: (1) compositional evolution of depleted mantle (Stracke et al., 2003) within a range of 4.552.9 Ga; (2) formation of enriched source (crust) and evolution of its composition between 2.9 and 2.5 Ga; (3) conservation of subducted modified lithosphere in mantle and its compositional evolution within a range < 2.5 Ga. Symbols are shown in Fig. 8.

intersects geochron and can be considered as isochron with age of extraction of parental material from depleted mantle (μ = 8.2, κ = 3.8, Stracke et al., 2003a) at 2.5 Ga ago. The distribution of data points along isochron denotes that source of their melts remained heterogeneous with respect to distribution of U, Th, and Pb up to involvement in magma gener ation and, owing to this, acquired varying values of isotope ratios, which are generally determined by iso chron dependence recording the time of formation of these heterogeneities. In the diagrams of Pb isotope composition (Fig. 9), the data points of the Orkhon–Selenga area with indi cated model age of 2.5 Ga are plotted in the linear field of compositional variations of the Late Cenozoic lavas of SKhVR. As compared to them, the rocks of the Taryat area and Lake Valley graben differ in lower 206Pb/204Pb and 208Pb/204Pb ratios, occupying an inter mediate position between them and enriched EMI component. This makes it possible to consider them as products of mixing between sources of the rocks of the Orkhon–Selenga area and Vodorazde’naya Group, on the one hand, and EMI, on the other hand. This is also supported by the negative correlations between 206Pb/204Pb and contents of some incompatible ele ments, which link the compositions of the rocks of the

Taryat and Lake Valley grabens of the Khangai area in a single trend. For the rocks of the Orkhon–Selenga area, this trend is determined by heterogeneities of melting protolith with age of 2.5 Ga. This suggests that compositional heterogeneities of the same age are characteristic of the protolith of the Taryat and Lake Valley grabens, and, hence, the EMI component has an age of about 2.5 Ga. Corresponding solution of the considered model can be obtained taking μ = 12.5 and κ = 3.4 for the stage of formation of EMI component, which corre sponds to composition of the Archean enriched source (Stracke et al., 2003a). The following, subduction related stage was presumably responsible for the loss of Th, U, and Pb and change of their ratios to values of μ = 5.5 and κ = 4.4. It should be noted that these pro cesses were accompanied by insignificant changes in 232Th/238U (κ). In this case, the isolation of subducting slab occurred about 2.5 Ga ago. Since the duration of Wilson cycles that determine the lifetime of oceanic basins with juvenile crust is estimated by the range of 200–500 Ma, we can suggest that enriched source (juvenile crust) was formed from depleted mantle approximately 300–400 Ma prior to subduction, i.e. about 2.9 Ga ago. Historical boundaries correspond ing to these assumptions in the evolution of EMI com PETROLOGY

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ponent, as well as its μ and κ parameters for this stage are shown in model (Fig. 12). Considered approaches to estimation of age of sub ducted lithosphere lead us to conclude that lavas of the Khangai area were presumably derived from ancient (Archean) lithosphere with enriched Nd, Sr, and Pb isotope characteristics, which was involved in recy cling no later than 2.5 Ma ago. Note that enriched sources of the Taryat graben, on the one hand, and volcanics of the Lake Valley and Vodorazdel’naya Group, on the other hand, consider ably differ in terms of Nd and Sr isotope characteris tics. However, they have practically identical Pb iso tope characteristics. The curves of isotope mixing between depleted and enriched sources for different isotope systems represented in Fig. 8b indicate insig nificant contribution of EMII enriched source in the composition of lavas of the Taryat area. The enriched source of the rocks of the Taryat graben presumably contained component with elevated Rb/Sr ratios (ter rigenous constituent of the subducted lithosphere), which determined the higher content of radiogenic Sr in magma source (Fig. 8b). As was shown above (Fig. 11), this source was also characterized by ele vated content of incompatible lithophile elements, for instance, La, Th, and Pb. Thus, the enriched sources of the rocks of the Taryat and Lake Valley areas could have some differences in geochemical composition, but experienced similar evolution (close time of the formation of enriched sources and their burial in man tle). The latter follows from similar Pb isotopic char acteristics in enriched components of the lavas (Fig. 9). Model of Isotope Structure of the Mantle Sources of Magmatism The complex of geological, geophysical, and iso topegeochemical data indicates the following: (1) Late Cenozoic volcanism of SKhVR has a within plate nature and forms local volcanic area, which is spatially separated from other occurrences of the Late Cenozoic volcanism in Central Asia; (2) this volcanic area is confined to the local rise of asthenosphere and heated mantle, whose roots are traceable to the lower boundary of upper mantle; (3) geochemical character istics of volcanic rocks show zoned distribution away from central parts of the area and, correspondingly, from the central parts of the asthenospheric rise; (4) the isotope compositions of the volcanic rocks cor respond to sublithospheric sources with characteristics of recycled lithosphere, which was isolated for a long time from melting processes. All these parameters make it possible to relate the volcanism of the area with mantle plume activity. Its development area at the Earth’s surface in the Late Cenozoic was complicated by boundaries between microplates (Fig. 1). The body of mantle plume is mainly localized beneath the Mon golian microplate and only its northeastern protuber PETROLOGY

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ance is traceable beneath the Amur microplate, where it was responsible for the formation of the Orkhon– Selenga volcanic area. Such a plume structure is also consistent with results of geophysical studies, which identified the asthenospheric rise beneath the central part of the Khangai highland (Zorin et al., 1988; 2004). Figure 13 demonstrates the model structure of this plume, which explains the distribution of isotopically different magma source within it. In accordance to the model, the plume has a mushroom shape, most uplifted part of which coincides with asthenospheric rise identified by geophysical methods. The zoned pat tern in distribution of depths and temperature of magma generation areas estimated on the basis of chemical and REE composition of volcanic rocks from various areas of the Khangai highland agrees with such structure of thermal mantle anomaly (plume), which has concentric pattern and causes the melting of mantle rocks. The shallowest depth of protolith melting corresponds to the central part of the mag matic area—watershed of the Khangai Range located above the mantle rise. Magmatic sources that con trolled the volcanic activity along periphery of the Khangai Highland were formed at greater depths and lower melting degree. The most primitive isotope composition of mag matic sources is noted in the volcanic products of the Vodorazdel’naya Group. In the Nd–Sr–Pb isotope diagrams, this composition resembles PREMA type, indicating the contribution of lower mantle material. The pathway of this mantle to the surface (or mantle plume channel) was recorded by teleseismic tomogra phy at least to boundary with lower mantle (Mordvi nova et al., 2007, 2008). Based on Nd and Pb isotope characteristics, the rocks of the Taryat and Lake Valley grabens and the Orkhon–Selenga area contain ancient recycled com ponent of Archean age, which was subsided in mantle in the Early Proterozoic time. This indicates that this component was involved in convective processes that spanned also lower mantle. This conclusion is well consistent with above mentioned similarity of the rocks of the Vodorazdel’naya group to the melting products of the lower mantle protolith. The isotopic variations in the rocks of the peripheral zone of the SKhVR suggest a compositional heterogeneity of the recycled material and the predominance of terrige nous component in its portion, which corresponds to the Khangai area. Thus, the model of formation of SKhVR in the Late Cenozoic testifies that this area was related to mantle plume fed by PREMAtype lower mantle material and recycled Early Precambrian lithospheric material. Note also a peculiar position of the plume beneath the volcanic area. It occurred within Central Asia ter ritory underlain by the large field of highvelocity “cold” mantle (Maruyama, 1994). The existence of this field is maintained by longlived system of numer

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SAVATENKOV et al. South Khangai volcanic region Khangai area Khangai Range OrkhonSelenga area LITHOSPHERE

UPPER MANTLE

ЕМII

ЕМI

LOWER MANTLE

ЕМI

PREMA

Fig. 13. Schematic section of SKhVR illustrating the model of plumeinduced withinplate magmatic activity in the Late Ceno zoic. Dark gray lenses correspond to fragments of entrapped recycled lithosphere with characteristics of enriched EMI and EMII sources. Dashed lenses show the position of magma generation areas.

ous zones of subducting oceanic lithosphere and litho spheric slabs sinking to the core, which is considered also as cold superplume (or slab graveyard) that cools mantle of this region. Such a deep structure of the mantle should prohibit the formation of hot mantle plumes. However, they exist. Moreover, as was shown earlier (Yarmolyuk et al., 1995; 1998; 2007b), some of them evolved within last ~ 160 Ma, the time of numer ous rearrangements of subduction systems at the east ern margin of the continent. This fact indicates that mantle plumes of Central Asia were not triggered by subduction processes, but controlled by deeper seated Earth’s shells, in particular, lower mantle. This con clusion is consistent with data on participation of such typical lower mantle sources as PREMA and Precam brian recycled lithosphere in the production of mag matic rocks of SKhVR. Correspondingly, it is in con flict with concepts of upper mantle nature of sources of withinplate magmatism in Central Asia and their derivation from young (Mesozoic) lithosphere sub ducted on the Pacific side. At the same time, we can not preclude that convective processes in the lower mantle that compensated the deep subsidence of large volumes of cold lithospheric material could also affect the activity of the central Asian plumes.

CONCLUSIONS The evolution of the SKhVR in the Late Cenozoic involved formation of numerous volcanic fields at the territory of the modern Khangai highland and its northeastern framing. The volcanic activity mani fested itself mainly in the form of plateau eruptions and valley ejections of the highly mobile mainly high K basic and intermediate lavas: foidites, basanites, basalts, trachybasalts, trachybasaltic andesites, and phonotephrites. The formation of volcanic area was controlled by activity of mantle plume, whose location at the base of the area was determined from gravimetric and seismo tomographic studies (Mordvinova et al., 2007, 2008) and finds support in signatures of the volcanic rocks. The composition and structure of the plume are deci phered from compositional zoning of the volcanic area, which is consistent with orographic zoning of the region. Center of this zoned structure is occupied by the volcanic sequences of the Vodorazdel’nyi graben and valley lava flows, which descent from the water shed part of the Khangai Range. The peripheral area is represented by volcanic associations formed within the Lake Valley and Taryat grabens and Orkhon– Selenga area in the framing of the Khangai Range. The compositional zoning is expressed in increasing alka PETROLOGY

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lis, Ti, P, and some other lithophile elements, as well as in systematic variations of isotope composition with distance from central part of the area. This zoning is caused by differences in P⎯T parameters of melting, which determine the relation of geochemical charac teristics of parental melts with changes in mineral composition of melting mantle protolith depending on depth of its location. Correspondingly, mantle plume that controls this zoning has a mushroom shape, with uplifted part reconstructed beneath the watershed of the Khangai Range, whose rocks (Vodor azdel’naya Group) are characterized by the highest melting degree of asthenospheric mantle uplifted to the shallowest depths. A zoned distribution of isotope characteristics of the rocks mirrors the distribution of mantle sources in the structure of mantle plume. Its channel that deter mined volcanism in the axial part of the Khangai Range (Vodorazdel’naya Group) was presumably fed in the lower mantle reservoirs with PREMAtype characteristics (Zindler and Hart, 1986; Wilson, 1989). An increase of enriched isotope components in lavas of peripheral volcanic areas together with corre sponding changes in chemical composition are not related with lithospherinc contribution in their forma tion, but reflect the distribution of compositionally different source in mantle plume. The most probable source of enriched components of the Late Cenozoic rocks of the SKhVR was Early Precambrian recycled crustal material, which after subduction was isolated from processes of upper mantle convection and only in the Late Cenozoic was transported by ascending man tle jet to the lithosphere bottom. Obtained data indicating the contribution of PREMAtype and Early Precambrian recycled lithos pheric components in the magmatism of SKhVR are inconsistent with concepts of derivation of these sources from a young (Phanerozoic) lithosphere, which could be supplied into upper mantle of the region due to subduction on the Pacific side. ACKOWLEDGMENTS The work was supported by the Russian Founda tion for Basic Research (project no. 070500876), Program No. 16 of the Presidium of Russian Academy of Sciences, and Program no. 10 of the Earth’s Sci ence Division of the Russian Academy of Sciences. REFERENCES 1. M. J. Le Bas and R. W. Le Maitre, A. Streckeisen, and B. Zanettin, “A Chemical Classification of Volcanic Rocks Based on the Total AlkaliSilica Diagram,” J. Petrol. 27, 745–750 (1986). 2. V. P. Afonin, N. I. Komyak, V. P. Nikolaev, and R. I. Plotnikov, XRay Fluorescence Analysis (Nauka, Novosibirsk, 1991) [in Russian]. PETROLOGY

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