the age and origin of volcanics in the riphean section of the

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The age and origin of volcanics in the Riphean section of the Siberian craton (western Baikal area)

D.P. Gladkochub a, *, A.M. Mazukabzov a, T.V. Donskaya a,B. De Waele b, A.M. Stanevich a, S.A. Pisarevskii c

a Institute of the Earth’s Crust, Siberian Branch of the RAS, 128 ul. Lermontova, Irkutsk, 664033, Russiab British Geological Survey, Kingsley Dunham Centre, Keyworth, NG5 5GG, UK

cUniversity of Edinburgh, Grant Institute, The King’s Buildings, West Mains Road, Edinburgh EH9 3JW, UK

Received 25 June 2007; in revised form 6 February 2008; accepted 31 March 2008Available online xx September 2008

Abstract

In the western Baikal area, the structural position, composition, and age of volcanic rocks in the section of the Riphean margin of theSiberian craton were studied. The age of these rocks, earlier assigned to the Khoto Formation, is estimated at 274±3 Ma (concordia constructedover 11 zircon grains, SHRIMP-II). The geochemical and isotope compositions of volcanics evidence that they resulted from the melting ofmantle source of EM-I type contaminated by crustal material. The intrusion of volcanics into the upper crustal horizons might have beencaused by the evolution of the Permian active margin of the Siberian continent, which took place on the background of the closure of theMongolo-Okhotsk ocean. Based on the results of studies, a new subvolcanic complex of Early Permian age has been recognized in the region,which includes the above-mentioned volcanics and earlier described porphyrite dikes of close age in the Sharyzhalgai uplift. The data obtaineddisprove the concept that the studied volcanics are of Riphean age; therefore, the available stratigraphic charts of the Siberian Precambrianmust be revised.© 2008, IGM SB RAS. Published by Elsevier B.V. All rights reserved.

Keywords: Active margin; Permian; zircon; Siberian craton; Mongolo-Okhotsk ocean

Introduction

On the southern flank of the Siberian craton, volcanic rocksoccur mainly in the Paleoproterozoic North Baikal volcanoplu-tonic belt (Fig. 1). Local exposures of volcanics extend in thesouthwestern direction from northern Baikal area to centralBaikal shore (Bugul’deika River basin) (Fig. 2). Up to now,the age of these volcanics and the formation to which theybelong have been unclear.

With lack of isotope-geochronological data, these volcanicswere usually assigned to the Khoto Formation (Fig. 3), whichwas dated at the Paleoproterozoic (Aleksandrov, 1990), Mid-dle Riphean (Maslov, 1983; Maslov and Kichko, 1985;Ryabykh and Ryabykh, 1979), or Late Riphean (Gladkochubet al., 2007; Postnikov, 2001; Stanevich et al., 2007). Onpaleogeodynamic reconstructions, the volcanics, along withdiscordantly overlying sedimentary rocks of the BaikalianGroup, were interpreted as a section of the Riphean passive

margin of the Siberian craton, which resulted from the breakupof Rodinia (Gladkochub et al., 2001, 2006a; Mazukabzov etal., 2001; Postnikov, 2001; Sklyarov, 2001a; Stanevich et al.,2007). In accordance with correlation schemes (Khomen-tovsky, 2002; Krasnov, 1983), effusive rocks of the KaragasGroup in the Sayan foretrough (adjacent to the western Baikalarea) are regarded as analogs of the described volcanics. TheLate Riphean age of these rocks (741±4 Ma) was confirmedby 39Ar/40Ar dating (Gladkochub et al., 2006b).

Since the age and genesis of the above volcanics areambiguous, we studied these rocks in the area of theirabundance (Fig. 3). The results obtained and their interpreta-tion are reported in this paper.

Geologic occurrence of volcanics

Exposures of volcanics of unclear age in the western Baikalarea are traceable for ∼60 km from the mouth of theBugul’deika River to the upper reaches of the Anga River,along the zone of tectonic contact of sedimentary strata of the

Russian Geology and Geophysics 49 (2008) 749–758

* Corresponding author.E-mail address: [email protected] (D.P. Gladkochub)

Late Riphean Baikal Group (Stanevich et al., 2007) with EarlyProterozoic granitoids of the Primorsky complex. In accord-ance with the legends of geological maps (Geological Map...,1964) and regional stratigraphic charts (Krasnov, 1983), thevolcanics belong to the Riphean Khoto Formation. The rocksof this formation are confined to the concave (northwestern)part of the Anga uplift of the Siberian Platform basement(Fig. 2).

For petrological and geochemical studies and isotope datingof the Khoto Formation volcanics, we carried out investiga-tions in the basin of the Barlog River (right tributary of theBugul’deika) (Fig. 2), where volcanics concordantly overliegranites of the Primorsky complex (Geological Map..., 1964).The contact of volcanic rocks with deposits of the lowerhorizons of the Baikal Group (Goloustnaya and Uluntui

Formations) is not exposed. Earlier data (Krasnov, 1983;Maslov, 1983; Maslov and Kichko, 1985; Ryabykh andRyabykh, 1979) lack information on the stratotype andreference section of the Khoto Formation.

The western slope of the exposure, where volcanic rocksfor dating were sampled (zone of contact of volcanics withthe Goloustnaya Formation deposits) (Fig. 3), is soded andforested. Near the contact zone, the Goloustnaya Formationdolomites occur only as scarce debris and in trenches locatedfar from the volcanic bedrocks. Volcanics come in contactwith granitoids of the Primorsky complex along the assumedfault of NE strike, which is inferred from forested depression.As the study area is denuded rather poorly, the volcanicsdirectly overlying granites were not observed in the Barlogbasin.

We tried to study contacts of volcanics with granitoids andsedimentary rocks of the Baikal Group in three profiles on theleft bank of the Bugul’deika (Fig. 3). But since the area isdensely forested and its topography is poorly dissected, wefailed to reveal immediate contacts of volcanics with Paleo-proterozoic granitoids and rocks of the Baikal Group in theseprofiles. The absence of contacts of volcanics with under- andoverlying deposits was noted earlier during the compilation of

Fig. 1. Schematic geologic structure of the Baikal area. 1–6 — complexes of theSiberian Platform: 1 — Phanerozoic sedimentary cover; 2 — nonmetamor-phosed sedimentary rocks of Riphean passive margin; 3 — metamorphosedsedimentary rocks of Riphean passive margin; 4 — Early Proterozoic grani-toids; 5 — North Baikal volcanoplutonic belt (Early Proterozoic); 6 — uplifts ofEarly Precambrian basement; 7–9 — Central Asian Fold Belt; 7 — areas ofEarly Paleozoic consolidation (Donskaya et al., 2000); 8 — area of activecontinental margin in the Devonian-Carboniferous (Zorin, 1999); 9 — area ofactive continental margin in the Permian-Early Jurassic (Zorin, 1999); 10 —main faults: a — observed, b — concealed; 11 — study area (see Fig. 2).

Fig. 2. Schematic geologic structure of the Bugul’deika and Anga basins (west-ern Baikal area). 1–4 — deposits of plate complex of the Siberian Platform: 1 —Neogene, 2 — Lower Cambrian, 3 — Vendian, 4 — Late Riphean (BaikalGroup); 5 — Khoto Formation; 6 — basement (Paleoproterozoic granite-meta-morphic deposits); 7 — Ol’khon terrane — segment of the Sayan-Baikal foldedarea (Caledonides); 8 — main faults; 9 — study area.

750 D.P.Gladkochub et al. / Russian Geology and Geophysics 49 (2008) 749–758

State geological maps of different scales (Geological Map...,1964) (Fig. 3). In the studied profiles, the Khoto Formationvolcanics are massive dark-green and greenish-gray green-stone-altered. No evidence for the presence of fluids, layering,and internal differentiation of strata was discovered in thestudied exposures. We failed to determine reliably the dipazimuth of volcanics because of the absence of visible contactsand the internal homogeneity of the strata, but weak signs ofbanding suggest the subsidence of rocks to the southeast.

Analytical techniques

We analyzed about 30 samples of volcanic rocks taken inthe Bugul’deika basin (Figs. 2 and 3). All of them wereanalyzed for major oxides, and twelve most representativesamples, for trace and rare-earth elements as well (Table 1).Major oxides were determined by chemical analysis at theInstitute of the Earth’s Crust, Irkutsk (analysts G.V. Bondarevaand N.Yu. Tsareva) and by the X-ray fluorescence method atthe Analytical Center of the Institute of Geology and Miner-alogy, Novosibirsk (analyst N.M. Glukhova). The contents of

Co, Ni, Sc, V, and Cr were determined by spectral analysisat the Institute of the Earth’s Crust, Irkutsk (analysts V.V.Shcherban’ and A.V. Naumova), and the content of Zr, by theX-ray fluorescence method at the Geological Institute, UlanUde (analyses were supervised by B.Zh. Zhalsaraev). Thecontents of other trace and rare-earth elements were deter-mined by ICP MS (VG Plasmsquad PQ-2) at the AnalyticalDepartment of the Irkutsk Scientific Center, Irkutsk (analystsS.V. Panteeva and V.V. Markova), following the technique ofGarbe-Schonberg (1993). The analytical error was no morethan 5%.

The isotopic composition of Nd was measured at HokkaidoUniversity, Japan. Powdered 100 mg samples were decom-posed in the HCl + HF + HNO3 mixture in Teflon bottles at110°C. The completeness of decomposition was checked underbinocular. Rare-earth elements were separated by standardcation exchange chromatography, and Nd, by extractionchromatography. The isotopic compositions of Nd weremeasured on a Finnigan MAT-261 multichannel mass spec-trometer in static regime. The measured 143Nd/144Nd ratioswere normalized to 148Nd/144Nd = 0.241578, which corre-sponds to 146Nd/144Nd = 0.7219, and then were reduced to143Nd/144Nd = 0.511860 in the La Jolla Nd standard. Theweighted average 143Nd/144Nd value in this standard was0.511839±0.000007 (n = 12). The blank sample contained0.3 ng Nd. The 147Sm/144Nd ratios were calculated from theconcentrations of Sm and Nd determined by ICP MS.

The εNd(T) and TNd(DM) values were calculated usingmodern CHUR (143Nd/144Nd = 0.512638, 147Sm/144Nd =0.1967) (Jacobsen and Wasserburg, 1984) and DM values(143Nd/144Nd = 0.513151, 147Sm/144Nd = 0.2136) (Goldsteinand Jacobsen, 1988).

The U-Pb zircon dating of volcanics (Curtin SHRIMP-II,standard samples CZ3, 564 Ma (Pidgeon et al., 1994), andBR266, 559 Ma (Stern, 2001)) was carried out at the MassSpectrometry Center of Techonological University (Perth,Australia) following the standard technique (Williams, 1998).Each analysis included six scanning cycles. The beam was∼30 µm. Corrections for terrestrial lean were introduced inaccordance with the model values (Stacey and Kramers, 1975).On the calculation of ages, the generally accepted values ofuranium decay constants (Steiger and Ja

..ger, 1977) were used.

The data obtained were processed using the SQUID program(Ludwig, 2001a), and the plots with concordia were con-structed using the ISOPLOT program (Ludwig, 2001b). Dia-gram with concordia was constructed in the coordinates ofTera and Wasserburg (1972).

Petrography and geochemistry of volcanics

For petrographic and geochemical studies, we sampledmore than 40 samples of volcanic rocks. Petrographic studyshowed that these are altered rocks of basic composition,sometimes of porphyritic texture. Some samples bear relics ofamygdaloidal structures; the amygdule cavities are filled withgranular epidote with traces of chlorite and greenish-brown

Fig. 3. Schematic geologic structure of the left bank of the Bugul’deika (Barlogmouth area). Compiled using Ryabykh’s Geological Map, scale 1 : 50,000(1973), and our data. 1 — Quaternary deposits; 2, 3 — deposits of the BaikalGroup (Riphean): 2 — Ulungui and 3 — Goloustnaya Formations; 4 — volcanicrocks of the Khoto Formation; 5 — granitoids of the Primorsky complex (EarlyProterozoic); 6 — dislocations: a — thrusts, b — faults; 7 — structural ele-ments: a — dip, b — granitoid fracturing; 8 — studied volcanic profiles andtheir numbers; 9 — localities of sampling for isotope dating.

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biotite. In most of the studied thin sections, the primarystructures cannot be identified because of intense secondaryalterations.

The main minerals in the studied volcanics are plagioclase,epidote, actinolite, and, sometimes, chlorite, which are presentin varying proportions. Secondary minerals are quartz andbiotite. The only relict mineral is plagioclase (laths) recrystal-lized to fine-grained saussuritized aggregate. Plagioclase relicscorrespond to andesine. Primary plagioclase has not beenpreserved in the volcanics. It was almost completely replacedby actinolite, chlorite, and epidote. Volcanic glass of thegroundmass was crystallized and replaced by thin acicularsegregations of actinolite, earthy aggregate of epidote, andscales of greenish chlorite.

Since the volcanics have undergone greenstone alterations,we made a geochemical typification of rocks and conclusionson their compositions and geodynamic nature based on studiesof the contents of elements (Ti, Zr, Y, Nb, Th, REE) whoseconcentrations are accepted as constant on low-temperaturetransformations (Sklyarov, 2001b).

The studied volcanics have SiO2 = 47.08–55.54%, Al2O3

= 13.61–15.46%, TiO2 = 0.56–0.83%, and P2O5 = 0.08–0.12(Table 1). According to Jensen’s (1976) classification, theserocks can be referred to as high-Mg tholeiitic basalts andcalc-alkalic basalts (Fig. 4). On the Nb/Y–Zr/TiO2 diagram ofWinchester and Floyd (1976), the composition points of thevolcanics lie in the fields of subalkaline basalts and an-desites/basalts (Fig. 5) and can be assigned to the samebasalt–andesite-basalt series.

The differentiation index (Mg#) of the studied volcanicsvaries from 69 to 43 (Table 1). On Harker’s diagrams (Fig. 6),Mg# shows a weak negative correlation with SiO2 (Fig. 6, a),

no correlations with TiO2 and P2O5 (Fig. 6, b, c), and a distinctcorrelation with Cr and Ni (Table 1, Fig. 6, d).

The volcanic rocks are characterized by moderately frac-tionated chondrite-normalized (Sun and McDonough, 1989)REE patterns ((La/Yb)n = 5–9) with a negative or absent Euanomaly (Eu/Eu* = 0.71–1.02) (Fig. 7).

The primitive-mantle-normalized (Sun and McDonough,1989) multielemental patterns of the studied volcanics (Fig.8) show distinct negative Nb-Ta, P, and Ti anomalies. Chaoticdistribution of contents is observed for Rb, Ba, K, and Sr,which might be due to the mobility of these elements onsecondary alterations of rocks. At the same time, LREE,HREE, and HFSE show parallel patterns, which evidences thattheir primary concentrations did not change and that thevolcanics belong to the same basalt–andesite-basalt series.

All analyzed rocks are enriched in Th, U, and LREE andare depleted in Ti, P, and HREE. This is the differencebetween the studied volcanics and typical N- and E-MORB(Fig. 8).

Isotope geochemistry of Nd

The Nd isotope studies were carried out for five samplesof volcanic rocks (Table 2). All the samples show extremelylow εNd(T) values (to –22.0) and TDM = 2.67–2.83 Gacorresponding to the Neo- and Mesoarchean. These valuesevidence a certain contribution of Archean substrate to thesource of the studied volcanics. Note that similar TDM values(2.62–2.72 Ga) were established for gneisses of the Goloust-naya marginal uplift of the Siberian craton (Gladkochub et al.,2008a,b), where the studied volcanics are localized, and forrapakivi granitoids of the Primorsky complex (Fig. 3) in thestudy area (Donskaya et al., 2005).

Isotope dating of volcanics

For isotope dating, we sampled volcanics (sample 05150)from the bedrock on the right bank of the Barlog River

Fig. 4. Al2O3–(FeO* + TiO2)–MgO diagram (Jensen, 1976) for the studiedvolcanics from the western Baikal area. BK — basaltic komatiites, CA —calc-alkalic andesites, CB — calc-alkalic basalts, CD — calc-alkalic dacites,CR — calc-alkalic rhyolites, PC — picrites, HFT — high-Fe tholeiites, HMT —high-Mg tholeiites, TA — tholeiitic andesites, TD — tholeiitic dacites, TR —tholeiitic rhyolites.

Fig. 5. Nb/Y–Zr/TiO2 diagram (Winchester and Floyd, 1977) for the studiedvolcanics.


Table 1Chemical compositions of representative volcanic rocks

Component 1 2 3

05150 05151 0638 0639 0640 01122 01124 01126 01127 01129 0661 0663 06426

SiO2, wt.% 51.68 51.55 48.49 51.89 48.52 48.59 47.08 51.15 51.94 50.40 55.54 53.58 52.52

TiO2 0.57 0.73 0.74 0.74 0.61 0.70 0.75 0.74 0.76 0.67 0.56 0.83 0.66

Al2O3 13.95 14.25 14.76 14.28 15.46 13.61 14.33 15.45 13.66 13.79 14.15 13.85 14.8

FeO 3.6 4.83 N.d. N.d. N.d. N.d. N.d. N.d. N.d. N.d. 6.65 5.53 5.47

Fe2O3 5.69 4.87 N.d. N.d. N.d. N.d. N.d. N.d. N.d. N.d. 2.8 4.62 3.96

Fe2O3* N.d. N.d. 11.20 10.16 10.66 10.74 11.36 9.90 6.97 9.56 N.d. N.d. N.d.

MnO 0.16 0.17 0.18 0.18 0.18 0.18 0.19 0.15 0.28 0.17 0.14 0.17 0.12

MgO 8.93 7.00 8.14 7.49 8.25 10.06 9.53 6.49 7.55 9.21 3.70 6.33 5.22

CaO 8.20 8.73 11.73 10.04 13.04 12.74 9.86 9.30 8.75 10.15 13.68 9.14 11.48

Na2O 2.30 4.34 2.06 2.79 1.61 1.50 3.55 2.49 3.49 3.50 0.56 1.91 3.65

K2O 0.78 0.14 0.09 0.31 0.06 0.21 0.23 0.38 0.81 0.54 0.27 0.36 0.12

P2O5 0.08 0.11 0.11 0.10 0.08 0.08 0.10 0.11 0.10 0.08 0.08 0.12 0.09

H2O– N.f. N.f. N.d. N.d. N.d. N.d. N.d. N.d. N.d. N.d. N.d. 0.18 0.1

LOI 3.72 2.96 2.22 2.17 1.76 2.05 2.95 3.33 4.77 1.69 1.98 3.32 2.26

CO2 N.f. N.f. N.d. N.d. N.d. N.d. N.d. N.d. N.d. N.d. 0.06 0.11 N.f.

Total 99.66 99.68 99.71 100.15 100.23 100.46 99.93 99.49 99.08 99.75 100.17 100.05 100.45

Mg# 64 58 59 59 61 65 62 56 68 66 43 54 51

Rb, ppm 21 8 1 10 1 2 2 8 48 13 8 9 N.d.

Sr 117 137 372 215 293 270 165 301 453 151 566 206 N.d.

Y 17 18 26 25 19 20 19 19 19 24 16 22 N.d.

Zr 69 70 100 100 80 77 73 77 81 86 78 106 N.d.

Nb 3 6 6 6 4 5 7 8 5 6 4 6 N.d.

Ba 359 200 108 163 49 57 78 238 297 283 138 191 N.d.

Co 16 37 42 46 50 70 36 51 52 41 30 43 N.d.

Ni 63 110 81 100 150 160 93 96 150 140 80 66 N.d.

V 110 220 280 290 240 250 230 340 430 230 240 250 N.d.

Cr 160 290 120 170 280 390 260 270 380 260 140 91 N.d.

La, / 16.93 13.47 25.80 20.54 14.40 17.44 13.96 15.48 16.02 20.61 11.91 18.37 N.d.

Ce 31.35 27.89 46.55 41.64 29.78 34.84 28.60 30.56 31.13 39.35 23.24 36.56 N.d.

Pr 3.84 3.42 4.83 4.43 33.4 4.37 3.11 3.92 3.48 5.18 2.63 4.22 N.d.

Nd 14.16 13.78 21.10 19.28 14.30 16.59 12.95 1573 14.06 18.67 11.90 18.47 N.d.

Sm 2.52 2.86 4.00 3.81 2.93 3.21 2.95 3.15 2.97 3.74 2.34 3.67 N.d.

Eu 0.58 0.90 1.09 1.03 093 0.78 0.93 0.91 0.90 0.89 0.84 1.13 N.d.

Gd 2.52 3.41 4.40 4.31 3.23 3.24 2.86 3.39 2.88 3.84 2.76 3.92 N.d.

Tb 0.39 0.54 0.65 0.66 0.54 0.55 0.52 0.55 0.51 0.63 0.42 0.65 N.d.

Dy 2.78 3.31 4.15 3.90 2.89 3.40 3.13 3.32 2.85 4.14 2.51 3.76 N.d.

Ho 0.55 0.72 0.92 0.81 0.67 0.75 0.63 0.76 0.61 0.92 0.54 0.80 N.d.

Er 2.12 2.26 2.45 2.30 1.78 2.01 2.00 1.89 1.83 2.49 1.50 2.03 N.d.

Tm 0.36 0.32 0.40 0.42 0.32 0.38 0.29 0.31 0.26 0.38 0.26 0.31 N.d.

Yb 1.35 1.74 1.87 1.84 1.54 1.88 1.65 1.61 1.44 2.24 1.26 1.80 N.d.

Lu 0.22 0.25 0.28 0.27 0.23 0.28 0.22 0.23 0.20 0.29 0.17 0.24 N.d.

Hf 1.77 1.06 1.24 1.19 1.23 1.29 1.07 1.27 1.18 1.30 0.99 1.45 N.d.

Ta 0.16 0.44 0.34 0.33 0.41 0.48 0.76 0.87 0.47 0.52 0.21 0.30 N.d.

Th 3.21 2.64 5.17 5.36 3.77 4.33 2.06 2.57 2.66 4.94 1.62 2.64 N.d.

U 0.94 0.88 0.79 0.75 0.56 0.61 0.40 0.39 0.95 0.66 0.22 0.40 N.d.

(La/Yb)n 8.4 5.2 9.2 7.4 6.3 6.2 5.6 6.4 7.5 6.1 6.3 6.8 —

Eu/Eu* 0.71 0.88 0.80 0.78 0.93 0.74 0.99 0.85 0.95 0.72 1.02 0.92 —

Note. 1, 2, 3, — numbers of profiles in Fig. 3. Major oxides in samples 05150, 05151, 0661, 0663, and 06426 were determined by chemical analysis, and theother components, by the X-ray fluorescence method. Mg# = 100⋅Mg2+/(Mg2+ + Fe*2+), where Mg2+ = MgO/40.31, Fe*2+ = FeO*/71.85; Eu/Eu* =Eun/√(Smn⋅Gdn) ; n — chondrite-normalized (Sun and McDonough, 1989); N.d. — not determined; N.f. — not found.


(52°44.135′ N, 106°02.420′ E) (Fig. 2) and extracted accessoryzircon. It has colorless to yellowish subeuhedral and euhedralcrystals 50 to 150 µm in size (elongation 1:1 and 2:1). Resultsof analyses of 13 zircon grains on a SHRIMP-II ion micro-probe are presented in Table 3 and Fig. 10. Three points (2,13, and 17) were excluded after the first scanning cyclebecause of the extremely high intensities of 204Pb on theimages. Cathodoluminescent (CL) photomicrographs of 11zircon grains show a distinct oscillation zoning (Fig. 9, a).The contents of 206Pbc vary from 0 to 1.32%, and the contentsof U and Th are 209–811 and 182–1085 ppm, respectively.All analyzed zircons have high Th/U ratios, 0.75–1.8, typicalof magmatic zircons.

On the 207Pb/206Pb–238U/206Pb diagram (Tera and Wasser-burg, 1972), the coordinates of 11 points of the studied zirconsform a concordant cluster with an age of 274±3 Ma (MSWD =1.07) (Fig. 10).

Besides the main population of zircons (with oscillationzoning), the heavy fraction includes two homogeneous zircons(grains 1 and 7 in Table 3 and Fig. 9, b). In addition tomorphological difference, these grains differ from the maingeneration of zircons in having lower contents of U and Th(69–203 and 75–159 ppm, respectively) but have similar Th/Uratios (0.81–1.13) (Table 3). Zircons 1 and 7 have lower238U/206Pb ratios and form a concordant pair with an age of321±9 Ma (Fig. 10). Probably, these zircons are trapped.

With regard to the morphology and geochemistry of zirconsof the main population evidencing their magmatic genesis, theobtained age of 274±3 Ma (Early Permian) might be inter-preted as the age of the zircon crystallization and, correspond-ingly, the age of the studied volcanics.

Discussion

The results obtained permit us to refine the age of thestudied volcanics and thus elucidate whether it is valid toassign them to the Khoto Formation and to separate the latteras an individual stratigraphic subunit. In the existing strati-graphic charts (Krasnov, 1983; Maslov, 1983; Maslov andKichko, 1985), the Khoto and Nugan Formations are regardedas age analogs and are attributed to the basement of theRiphean sections of the Baikal Group. A specific feature ofthe Khoto Formation sections is the presence of volcanics of

Fig. 6. Harker’s diagram for the studied volcanics from the western Baikal area.

Fig. 7. REE patterns of the studied volcanics. Normalized after Sun andMcDonough (1989).


unclear age. Taking into account the determined age of thestudied volcanics (Early Permian) and the absence of anyreliably documented correlations between them and sedimen-tary rocks in the Khoto Formation area (Bugul’deika basin),we separated the volcanic rocks of this formation as asubvolcanic complex of Permian age. This makes using theterm “Khoto Formation” as an individual stratigraphic subunitunnecessary, because the remaining (terrigenous) part of theformation section will correspond in structure to the NuganFormation. The latter is widespread in the western Baikal area,and, as things stand now, its Riphean age is beyond question.

The specific geochemistry of the studied volcanics, inparticular, the presence of negative anomalies of Nb-Ta andTi on their multielemental patterns, suggests two possiblesources of such rocks: (1) lithospheric-mantle source enrichedduring subduction (Fitton et al., 1988; Hawkesworth et al.,1993) and (2) mantle source contaminated by crustal material

(Halama et al., 2004; Wilson, 1989). But the Nd isotopecomposition of the volcanics, namely, their extremely lowεNd(T) values (Table 2), supports the second model, as it pointsto a certain contribution of crustal material into the magmageneration source.

To prove the possible formation of volcanics contaminatedby crustal material, we calculate the model mixing of twocomponents (Jahn et al., 2000):

Xm = [(εc − εmc)⋅Ndc]/[εmc⋅(Ndm − Ndc) − (εm⋅Ndm − εc⋅Ndc)],

where Xm is the portion of mantle component, Ndm and Ndcare the concentrations of Nd in mantle and crustal components,respectively, and εmc, εm, and εc are the values of εNd for the

obtained mantle-crustal mixture and mantle and crustal com-ponents.

Fig. 8. Multielemental patterns of the studied volcanics. The compositions of primitive mantle, N-MORB, E-MORB, and OIB were borrowed from Sun andMcDonough (1989), and the composition of LCC (lower continental crust), from Rudnick and Fountain (1995).

Table 2Sm-Nd isotope data for the studied volcanics from the western Baikal area

Sample Age, Ma Content, ppm 147Sm/144Nd 143Nd/144Nd(±2σ meas.)

εNd(T) TNd(DM), Ma

Sm Nd

0638 274 4.00 21.10 0.1142 0.511400±12 –21.3 2671

0639 274 3.81 19.28 0.1191 0.511465±11 –20.2 2703

0640 274 2.93 14.30 0.1235 0.511470±12 –20.2 2826

0661 274 2.34 11.90 0.1183 0.511370±10 –22.0 2831

0663 274 3.67 18.74 0.1178 0.511376±9 –21.9 2806


As a crustal component, we use the Nd isotope compositionof Paleoproterozoic granites of the Primorsky complex (local-ized in the immediate vicinity of the studied volcanics, Fig. 3)calculated for the time of the volcanics formation, 274 Ma (εc = −25)(Donskaya et al., 2005). We take the concentration of Nd asits average content in the Primorsky complex granites (50ppm) (Donskaya et al., 2005). As a mantle component, weuse the Nd isotope composition of the source EM-I (Zindlerand Hart, 1986) calculated for the age of 274 Ma(εm = −5.8). The Late Paleozoic subduction of the crust of theMongolo-Okhotsk ocean beneath the Siberian continent (Vander Woo et al., 1999; Zorin, 1999) permits the mantle sourceEM-I typical of active continental margins (Sklyarov, 2001b)

to be used in our calculations. We take the content of Ndequal to 9 ppm, as in E-MORB (Sun and McDonough, 1989).The average εNd value for the analyzed volcanics is about –21(Table 3). Substituting these values into the above formula,we obtain the following proportion for the initial melt: mantlecomponent — 60% and crustal component — 40%. Thus, thecalculation data suggest that the studied volcanics resultedfrom the melting of the mantle source EM-I contaminated bycrustal component.

Analysis of the spatial occurrence of the volcanics and newpetrological and geochronological data suggest that the intru-sion of these rocks might have been caused by the intenseevolution of the Permian active margin of the Siberian

Fig. 9. Cathodoluminescent (CL) images of zircons from the studied volcanics: a — main population, b — trapped grains. Numerals are the grain numbers.

Table 3Results of U-Pb dating of zircons from volcanics of the Khoto Formation (sample 05150)

Sample,crystal

206Pbc, % U, ppm Th,ppm

232Th/238U Isotope ratios Age, Ma

(1) 238U/206Pb* ±1σ (1)207Pb*/206Pb*

±1σ (1) 206Pb/238U (1) 207Pb/206Pb

05150-1 0.154 69 76 1.13 19.54171 0.42599 0.05385 0.00444 322±7 365±186

05150-2 0.101 811 1085 1.38 21.02667 0.33624 0.05199 0.00078 300±5 285±34

05150-3 0.681 222 212 0.99 22.68187 0.40512 0.04723 0.00339 278±5 61±171

05150-4 — 562 573 1.05 23.46471 0.38231 0.05249 0.00064 269±4 307±28

05150-5 0.132 212 307 1.50 23.62178 0.42130 0.05158 0.00154 267±5 267±69

05150-6 0.162 209 201 0.99 22.93281 0.41016 0.05177 0.00301 275±5 275±133

05150-7 0.024 203 159 0.81 19.56112 0.34518 0.05154 0.00114 321±6 265±51

05150-8 — 455 413 0.94 22.77771 0.40878 0.05274 0.00096 277±5 318±41

05150-9 1.317 336 306 0.94 22.37975 0.38632 0.05306 0.00341 282±5 331±146

05150-10 0.414 469 674 1.49 28.46016 0.47224 0.04914 0.00178 223±4 155±85

05150-11 0.062 640 1116 1.80 29.26653 0.47767 0.04944 0.00087 217±3 169±41

05150-12 0.167 421 535 1.31 22.88523 0.37951 0.05201 0.00138 276±4 286±60

05150-14 0.776 266 297 1.15 22.84724 0.40718 0.05473 0.00350 276±5 401±143

05150-15 0.504 356 372 1.08 22.77893 0.38473 0.05009 0.00232 277±5 199±108

05150-16 0.636 366 266 0.75 23.47889 0.39737 0.04896 0.00265 269±4 146±127

05150-18 0.223 243 182 0.78 24.12520 0.42506 0.05210 0.00168 262±5 290±74

Note. The calculation errors are at the 1σ level. Pbc and Pb* are terrestrial and radiogenic lead, respectively. The error of calibration of the standard BR266

sample was 0.48% (2σ). (1) — correction for terrestrial lead was made by measured 204Pb.


continent (Fig. 1) (Zorin, 1999). The southern flank of thecraton, along with the terranes of different geodynamicsaccreted in the Early Paleozoic (Donskaya et al., 2000), mightbe interpreted as part of the Late Paleozoic volcanic belt,which developed above the subduction zone subsided beneaththe Siberian continent during the closure of the Mongolo-Ok-hotsk ocean (Van der Woo et al., 1999; Zorin, 1999). ThePermian magmatic activity within the belt is evidenced notonly by the studied volcanics but also by the nearly coevaldike swarms of the Sharyzhalgai uplift (275±4 Ma (Pisarevskyet al., 2006)) and southern Baikal area (275–300 Ma (Shadaevet al., 2005)). Up to now, Early Permian volcanic rocks havenot been revealed in the southern Siberian craton (Luchitskii,2001). The performed studies, along with the earlier data(Pisarevsky et al., 2006), show that it is necessary to recognizea new subvolcanic complex (Bugul’deika) of Early Permianage in the western Baikal area.

On the southern margin of the Ol’khon terrane (Fig. 2)accreted to the Siberian craton in the Early Paleozoic (Gladko-chub et al., 2008), volcanic rocks of unclear age werediscovered not far from the area of occurrence of the studiedvolcanics. Some of their varieties are similar in petrochemicalcomposition to Early Permian volcanic rocks (Zamaraev,1979). Perhaps these geochemical analogs might also bereferred to as indicators of extension processes on theperiphery of the Siberian continent (Zorin, 1999) in the LatePaleozoic. We consider the reported concept of the genesisand age of the studied West Baikal volcanics to be the mostsubstantiated, though we do not rule out other interpretation.

Conclusions

Based on the results of study of volcanics of the westernBaikal area, we have drawn the following conclusions:

The volcanic rocks that were earlier assigned to theRiphean Khoto Formation have an age of 274±3 Ma (concor-dia constructed over 11 zircon grains). This date indicates thepresence of such volcanics on the Riphean margin of theSiberian craton, which requires the revision of availablestratigraphic charts, including the Regional Stratigraphic Chartof Upper Proterozoic deposits of southern Siberia (Krasnov,1983).

The geochemical and isotope compositions of the studiedvolcanics suggest that they resulted from the melting of themantle source EM-1 contaminated by crustal material.

The intrusion of these volcanic rocks might have beencaused by the evolution of the Permian active margin of theSiberian continent. Their age analogs (evidence for magmaticand volcanic activity on the active margin during the closureof the Mongolo-Okhotsk ocean) are widespread in theTransbaikalian areas adjacent to the southern part of the craton(Shadaev et al., 2005; Zorin, 1999).

The performed studies confirmed the necessity of recogni-tion of a new Early Permian subvolcanic complex in thewestern Baikal area. This complex includes the studiedvolcanics and earlier described nearly coeval porphyrite dikesof the Sharyzhalgai uplift (Pisarevsky et al., 2006).

We thank V.A. Vernikovsky and A.B. Kuz’michev forvaluable remarks and useful recommendations.

This work was supported by Russian President’s grantsMD-242.2007.5 and NSh 3082.2008.5, grants 08-05-00245,08-05-00177, and 07-05-00339 from the Russian Foundationfor Basic Research, the Russian Science Assistance Founda-tion, Integration Programs ONZ-10.1 and ONZ-6.5 of theSiberian Branch of the RAS, and grant RNP.2.2.1.1.7334NOTs “Baikal” from the Departmental Integral Program“Development of the Scientific Potential of the Higher School(2006–09)”.

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