High-magnesium picrite–basalt associations of the Tunka terrane (Baikal–Hövsgöl region) as an indicator of the back-arc basin spreading

www.elsevier.com/locate/rgg

High-magnesium picrite–basalt associations of the Tunka terrane (Baikal–Hövsgöl region) as an indicator of the back-arc basin spreading

S.I. Shkol’nik *, V.G. Belichenko, L.Z. ReznitskiiInstitute of the Earth’s Crust, Siberian Branch of the Russian Academy of Sciences, ul. Lermontova 128, Irkutsk, 664033, Russia

Received 13 February 2012; accepted 23 March 2012

Abstract

High-Mg metabasalts and metapicrites discovered within the Urtagol Formation in the central zone of the Tunka bald mountains (EastSayan) are studied. In geochemistry the high-Mg metavolcanics are similar to subductional rocks. We have established that the Nb-rich recycledmaterial of oceanic crust (RSC) was a source of elements for high-Ti metabasalts, and subductional fluid rich in LREE and Th relative toNb was the source of these elements for high-Ti metapicrites. The enrichment of low-Ti metavolcanics, formed, probably, at the early stagesof the basin opening, was due to the contamination of melt with continental-crust material. A comparison of the metavolcanics withnonmetamorphosed analogs is made, and some genesis aspects are considered. The results obtained led to the conclusion that the metavolcanicsmark the paleospreading of the back-arc basin.© 2013, V.S. Sobolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved.

Keywords: metavolcanics; picrites; basalts; geochemistry; back-arc spreading; Tunka terrane

Introduction

Formation of high-Mg rocks (picrites, picrite basalts) is arather rare phenomenon in island-arc settings. Such rocks aremost typical of large igneous provinces (LIPs), hot spots, riftsystems, and mid-ocean ridges (Herzberg and O’Hara, 1998;Kerr et al., 1996; Larsen et al., 2003; Tsikouras et al., 2008).This is due to the conditions necessary for the generation ofhigh-Mg melts, first of all, high temperature and high degreeof melting. The formation of picritic volcanics in island arcs(Lesser Antilles, Japanese, Solomon) is associated with youngsubduction zones, hot oceanic crust, and anomalous geotherms(Schuth et al., 2004). Formation of picrite–basalt associationsduring subduction processes is also reconstructed for ancientsettings (Izokh et al., 2010; Sharkov and Smolkin, 1997). Insubduction settings, such associations were discovered notonly in island arcs but also among volcanics of back-arcbasins. In the latter case they are regarded as markers ofback-arc spreading (Junichi et al., 2004; Perfit et al., 1996;Sharkov and Smolkin, 1997). Therefore, it is of special interestto study the primitive high-Mg metavolcanics of the Tunkaterrane, where no spreading settings were earlier established.

The geologic position and composition of the Tunka terrane

The Tunka terrane is part of the complex tectonic ensembleof the Baikal–Hövsgöl segment of the Central Asian mobilebelt, which also includes the Khamar-Daban metamorphic andDzhida island-arc terranes and the Tuva–Mongolian massif(paleomicrocontinent) (Fig. 1). Most of the Tunka terrane isthrust over the Tuva–Mongolian massif and rests upon it as acomplex allochthone (Belichenko et al., 2002). The terrane isseparated from the Siberian craton by the Belaya–Kitoi(Kitoikin) metamorphic terrane. According to earlier ambigu-ous paleontological data, the terrane deposits accumulatedfrom Cambrian to Silurian (Boos, 1991). The recently deter-mined age of the granitoids of the Munku–Sardyk massif (481± 2 Ma; accessory-zircon dating (Reznitskii et al., 2007))marks the upper bound of the time of the terrane volcanism,sedimentation, and metamorphism, whereas the lower boundis still debatable.

The chemogenic-volcanogenic-terrigenous terrain unit 2.0–2.2 km in total thickness is subdivided into two formations:Urtagol and overlapping Tolta (Boos, 1991) (Fig. 2). Theydiffer in the quantitative proportion of rock lithotypes: TheUrtagol Formation is composed mainly of metaclastic rocksand metavolcanics, with scarce carbonate rocks, whereas theTolta Formation is, on the contrary, made up predominantly

Russian Geology and Geophysics 54 (2013) 153–165

* Corresponding author.E-mail address: [email protected] (S.I. Shkol’nik)

Available online at www.sciencedirect.com

ed.

+1068-7971/$ - see front matter D 201 IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reservV S.. S bolevo,

http://dx.doi.org/10.1016/j.rgg.201 .0 .003 1 33

of carbonate rocks and metavolcanics, with subordinate meta-terrigenous lithotypes. The Tunka terrane, together with theKhamar-Daban terrane, northern part of the Dzhida terrane,and adjacent area of the cover of the Tuva–Mongolian massif,underwent polyfacies metamorphism of a wide temperaturerange (Belichenko and Boos, 1988), which varies fromlow-temperature greenschist to moderate-temperature amphi-bolite facies and is of disthene–sillimanite type. Paragenesesin rocks of all lithotypes vary depending on the degree ofmetamorphism. Clastic rocks of low-grade metamorphism arechlorite–sericite schists, often carbonate-containing. As thedegree of metamorphism grows, biotite, garnet, staurolite, anddisthene successively appear in the rocks, which are changedby fibrolite-containing garnet–two-mica gneisses in the high-temperature zone. The gneisses often preserve relict stauroliteand disthene. Among the carbonate rocks, there are calcareous,magnesian, and siliceous varieties. In the siliceous-dolomiticrocks, quartz-dolomite paragenesis is changed by tremolite-and diopside- to forsterite-containing marbles.

Metavolcanics of the Tunka terrane. Lithotypes falling intothe fields of igneous rocks (orthometamorphites) on most ofidentification diagrams were referred to as metavolcanics. Weused diagrams permitting identification of most of major com-ponents, in particular, the well-known diagrams of N.A. Do-maratskii, N. de La Roshe and M. Roubault, A.A. Predovskii,and A.N. Neelov. Here, we present only Neelov’s diagram(Neelov, 1980) showing earlier published and new data on theterrain metavolcanics (Fig. 3).

Most metavolcanics lack structural or mineralogical fea-tures of initial rocks and can be identified only by chemicalcomposition. However, Boos (1988, 1991) discovered differ-ently preserved relics of primary magmatic structures andtextures in metavolcanics near the least metamorphosednorthern margin of the terrane, in the Tunka bald mountains(below the isograde of biotite in metapelites and actinolite inbasic rocks). In particular, laths of saussuritized plagioclasehave a spilitic texture (Fig. 4a, b), sometimes porphyritic, withan intersertal groundmass. Mesostasis is completely replacedby an aggregate of chlorite, epidote–clinozoicite, albite, car-

bonate, ore mineral, and quartz. In places, rocks of amygdaloi-dal structure, with amygdales similar in composition tomesostasis, were found. Also, transitions of rocks with relicsof primary structures to greenschists are observed. The schistsof low-grade metamorphism are of essentially chloritic (withepidote) composition. As the degree of metamorphism grows,this paragenesis is changed by chlorite–epidote–actinoliteone and then by epidote–actinolite (chlorite-free) paragenesis;also, biotite can appear. In the sillimanite–fibrolite zone,actinolite gives way to hornblende; sometimes, garnet isproduced in amphibolites. Correspondingly, the granular com-position changes from fine- to small- and medium-crystalline(Fig. 4c, d).

Along with effusive rocks, pyroclastic rocks (tuffs andtuffites—mixed terrigenous-carbonate-pyroclastic rocks withnumerous preserved fragments of pyroclase laths) have beendiscovered in the area of initial metamorphism (Fig. 5). Bythe type and size of fragments, Boos recognized crystal andlithic (fragments of spilites or altered glass) clastics withpsammitic or silt-pelitic particles among them. A specificfeature of pyroclastic and mixed rocks is a high Ca/(Ca + Al)ratio and layering of different grades (fine alternation withterrigenous, carbonate, and calc-silicate rocks). These featuresare preserved during metamorphism and help to distinguishbetween tuffaceous rocks and metaeffusive ones in the entiremetamorphic section of the terrane, though these rocks havean identical qualitative mineral composition.

In general, the relict structures of orthometamorphites andtheir permanent association with metapyroclastic rocks indi-cate that they belong to metavolcanics, most likely, to eruptedeffusive rocks rather than sills and dikes. Only ultrabasic rockspresent as serpentinite lenses can be considered intrusive. Inthe rock unit, metavolcanics form concordant sheet-like bodies10–20 to 100 m thick; their amount is variable and can locallyreach 30 vol.%. The diagram in Fig. 3 shows their compositionrange. In basicity the metavolcanics vary from andesites andandesite-basalts to basanites, with a strong predominance ofbasalts. The Mg# value also varies over a wide range (from2.2 to 16.3 wt.% MgO). Variations in chemical and mineralcompositions are expressed not as a change of paragenesesdepending on the degree of metamorphism but as differentquantitative proportions of salic and femic minerals anddifferent Fe/(Fe + Mg) ratios of the latter.

Earlier we considered the petrochemistry of metavolcanicsof the Tolta Formation on the western flank of the Tunka basin(Fig. 2) (sections along the Irkut River and its right tributar-ies—Belyi Irkut, Srednii Irkut, and Aerkhan Rivers) (Shkol’-nik et al., 2009). These metavolcanics were reconstructed astholeiitic basalts and, partly, andesite-basalts and andesites,with geochemical parameters specific for basaltoids of back-arc basins. They were characterized by low and medium Mg#values (MgO = 3.7–7.7 wt.%). Despite the detailed samplingand representative number of samples, no high-Mg varietieswere found. Later we sampled high-Mg metavolcanics alongthe Bogo-Khongoldoi River (right tributary of the Kitoi River)in the central zone of the Tunka bald mountains (East Sayan),

Fig. 1. Schematic occurrence of terranes in the Baikal–Hövsgöl region. Terra-nes: Dzh, Dzhida, KhD, Khamar-Daban, Tn, Tunka, Kt, Belaya–Kitoi (Ki-toikin); TM, Tuva–Mongolian microcontinent; SC, Sharyshalgai uplift of theSiberian craton. The area from Fig. 2 is outlined.

154 S.I. Shkol’nik et al. / Russian Geology and Geophysics 54 (2013) 153–165

Fig. 2. Schematic occurrence of metavolcanics of the Tunka terrane. 1, Cenozoic deposits; 2, granitoids; 3, syenites; 4, basic rocks; 5, Sagansair Formation; Tunkaterrane: 6–8, Tunka Group: 6, Tolta Formation, 7, Urtagol Formation, 8, horizons containing metavolcanics; 9–11, Khamar-Daban terrane: 9–10, Khangarul Group:9, Bezymyanskaya Formation, 10, Kharagol Formation, 11, Slyudyanka Group; 12, Tuva–Mongolian massif; 13, thrust boundary between the Tunka terrane and theTuva–Mongolian massif; 14, predicted boundary between the Tunka and Khamar-Daban terranes; 15, sampling locality: 1, sections along the Irkut River etc.,2, section along the Bogo-Khongoldoi River.

Fig. 3. Neelov’s (1981) a = Al/Si–b = Fe3+ + Fe2+ + Mn + Ca + Mg classification diagram for high-Mg metavolcanics of the Tunka terrane. 1, high-Ti metavolcanics,2, low-Ti metavolcanics; 3, 4, data from Boos (1991) and Shkol’nik et el. (2009): 3, high-Ti metavolcanics, 4, low-Ti metavolcanics. Composition fields of:1, ultrafelsic liparites; 2, liparite-dacites; 3, dacites; 4, andesite-dacites; 5, andesites; 6, andesite-basalts; 7, basalts; 8, basanite-basalts; 9, basanites; 10, picrites;11, meiymechites. Dash-dotted line marks the boundaries between supergroups and groups, after Neelov (1980).

S.I. Shkol’nik et al. / Russian Geology and Geophysics 54 (2013) 153–165 155

in the area of the Urtagol Formation rocks (Shkol’nik et al.,2011).

High-Mg metavolcanics and associated metabasalts in the Tunka bald mountains

In contrast to the volcanics of the Tolta Formation,concordant sheet bodies of high-Mg metavolcanics and tholei-itic metabasalts alternate not with carbonate but mainly withmetaterrigenous aluminosilicate and calc-silicate rocks. An-other peculiarity is the presence of lenticular serpentinitebodies in the same rock unit. The rocks of some metavolcanicbodies vary in chemical composition, but we did not find rocksstrongly different in Mg# (i.e., high- and low-Mg) within thesame body. Parageneses of high-Mg metavolcanics, like thoseof tholeiitic metabasalts, include chlorite, actinolite, epidote,albite, and magnetite in different proportions. No primarymagmatic structures have been preserved.

The limits of using the term “picrites” in the paper.Restriction on using this term is necessary because theclassification of the so-called high-Mg volcanics was repeat-edly revised and not all researchers accept the recommenda-tions of corresponding commissions and committees, inparticular, ones concerned with the limits of using the term“picrite”. This term was introduced in the early 19th centuryfor olivine-rich effusive rocks, which were considered to bevolcanic analogs of peridotites, olivine-enriched gabbro, andteschenites. Later, the nomenclature of picrites was discussedfor a long time. In accordance with the recommendations ofthe Subcommission on Systematics of Igneous Rocks of theInternational Union of Geosciences (IUG), the TerminologicalCommission of the Petrographic Committee of the USSR’sAcademy of Sciences (Bogatikov et al., 1981) recommendedto recognize a family of picrites, including picrites, meyme-chites, and komatiites. The recommended limiting contents ofcomponents for picrites are as follows (wt.%): MgO = 20–32,SiO2 = 39–43.5, Na2O + K2O ≤ 1, CaO ≤ 7.5, and Al2O3 ≤

Fig. 4. Textures of metavolcanic rocks. a, b , Differently preserved metavolcanics with a relict spilitic texture; c, fine-grained epidote–chlorite schist; d, fine-grainedepidote–actinolite schist. Thin section, ×20 (Boos, 1988).


8–8.5. The petrochemical coefficients of picrites lie in a verynarrow range: a = Al2O3 + CaO + Na2O + K2O = 9–14, S =SiO2–(Fe2O3 + FeO + MgO + MnO + TiO2) = –4 to 1.5.Picrite-basalts and picrite-dolerites (MgO = 12–24%) wereregarded as basic rocks, though part of them was assigned toultrabasic rocks according to the limiting contents of silica(SiO2 = 42–46%). In the next publications (Bogatikov, 1983;Laz’ko and Sharkova, 1988), these limiting contents werepreserved, but the authors sometimes ignored them in thedescription of particular rocks, e.g., rocks with CaO ≤10–14%, Al2O3 ≤ 10–11%, and MgO > 17% were sometimesreferred to as picrites.

Considerable changes in the limits of using the term“picrite” were made by the Subcommission on Systematics ofIgneous Rocks of the IUG, whose recommendations wereapproved at the International Geological Congress in 1989 andpublished in Russia in 1997 (Efremov et al., 1997). Thelimiting contents of the main parameters of picrites, Na2O +K2O and SiO2 (TAS), were risen to 53% SiO2 (i.e., expandedto the entire range of values for basic rocks) and to 1–2%(Na2O + K2O) (instead of former <1%). The limiting contentof MgO was reduced to 18%. No restrictions on contents ofCaO and Al2O3 were made. For picrite-basalts, the narrowranges of SiO2 contents, the same as for basanites (41–43%),were left unchanged. The recommended classification ofhigh-Mg rocks was not accepted as satisfactory and there wasrevised by the International Working Group of the Subcom-mission of the IUG (Le Bas, 2000). The SiO2 contentboundary between the picrite and boninite families was shiftedfrom 53 to 52 wt.%. The range of MgO contents for picriteswas established as 12–18 wt.% (lower contents for picrite-ba-salts and higher ones for meymechites and komatiites). Thelimiting (Na2O + K2O) content was risen to 3 wt.% forpicrite–basalts (to 2 wt.% for meymechites and komatiites).The limiting contents of CaO and Al2O3 are still notestablished. These recommendations were adopted by someresearchers. For example, following the recommendationsmade in 2000, picrites of the Solomon Isles (Schuth et al.,2004), Mino accretionary complex, Japan (Ichiyama andIshiwatari, 2005), Pelagonian continental margin (Tsikouras etal., 2007), Tarim basin (Tian et al., 2010), etc. were recog-nized. In this paper we also accept the latest classification,i.e., we regard picrites as volcanics with MgO = 12–18, SiO2≤ 52, and (Na2O + K2O) ≤ 3 wt.%, sometimes with minordeviations from these limiting values. Deviations are allowedfor some samples that get into clusters or trends of correspond-ing group on various composition diagrams.

The geochemistry of metavolcanic associations. The con-tents of major components in 40 metavolcanic samples weredetermined by the “wet chemistry” method and X-ray spec-troscopy. The concentrations of trace elements were estab-lished by X-ray fluorescent analysis (Table 1), and those ofREE, Ta, Th, U, and Cs were measured by ICP MS on aPlasma Quad PQ2 mass spectrometer, using International andRussian standard samples. All analyses were carried out in thelaboratories of the Institute of Geochemistry and Institute ofthe Earth’s Crust, Irkutsk. Isotope studies were performed on

a Finnigan MAT-262 mass spectrometer at the Center ofCommon Use of the Irkutsk Scientific Center.

The common problem of preservation of the primary major-and trace-element compositions of metavolcanics was earlierconsidered for the metaeffusive rocks of the Tunka terrane(Shkol’nik et al., 2009). The distinct correlation between thecontents of major and trace elements, specific for igneousrocks, led to the conclusion that most of HFSE, REE, andtransitional trace elements are well preserved. In some sam-ples, the contents of K, Sr, and Rb showed deviations fromthe correlation trends.

Since the high-Mg metavolcanics are similar in petrochemi-cal composition (high contents of MgO and SiO2) to boniniticrocks but differ either in higher contents of TiO2 (>0.5) or inlower contents of SiO2, all studied metavolcanics were dividedinto two groups according to TiO2 content: low-Ti (<0.5 wt.%TiO2) and high-Ti (>0.5 wt.% TiO2), i.e., their boundary valueis the accepted boundary between boninites and picrites. Thus,the group of low-Ti metavolcanics includes rocks containing(wt.%): SiO2 = 46.2–51.7, Al2O3 = 11.5–14.8, and MgO =9.4–17.1. The group of high-Ti metavolcanics shows a widerrange of the contents of major components (wt.%): SiO2 =43.1–52.3, MgO = 6.7–17.3, Al2O3 = 8.8–15.4, and TiO2 =0.51–2.7. Each group has varieties that fall in the fields ofpicrite basanites and medium-alkali picrites on the petrochemi-cal diagram (Neelov, 1980) (Fig. 3). According to theclassification by Le Bas (2000), these varieties are referred toas picrites (Fig. 6). Let us consider each group of thesemetavolcanics.

Low-Ti metavolcanics of the Urtagol Group. The group oflow-Ti metavolcanics (TiO2 = 0.07–0.44 wt.%) includes rocks

Fig. 5. Metatuffite with preserved fragments of plagioclase laths. Thin section,×20 (Boos, 1988).


Tab

le 1

. C

onte

nts

of m

ajor

(w

t.%)

and

trac

e (p

pm)

com

pone

nts

in r

epre

sent

ativ

e sa

mpl

es o

f m

etav

olca

nics

of

the

Tun

ka t

erra

ne

Com

pone

ntH

igh-

Ti

met

apic

rite

sH

igh-

Ti

met

abas

alts

43/8

686

6/11

55b/

8660

a/86

865/

886

5/19

29/8

653

a/86

56a/

8686

11/2

865/

586

5/16

865/

1786

5/18

865/

2086

6/14

867/

386

9/7

17b/

8625

/86

27/8

6

SiO

248

.19

52.2

651

.56

46.6

350

.96

51.9

948

.11

50.9

354

.56

47.8

150

.04

51.3

546

.53

48.6

650

.88

46.9

547

.46

49.7

149

.26

46.5

343

.14

TiO

20.

560.

660.

810.

510.

770.

730.

791.

014

0.79

0.73

0.96

1.05

1.19

1.18

1.55

2.79

0.99

1.6

1.02

1.19

2.35

Al 2

O3

8.78

10.9

99.

459.

9010

.03

9.82

15.3

914

.82

11.0

013

.19

14.8

513

.38

13.4

14.0

713

.11

12.2

314

.77

16.8

12.1

513

.412

.58

FeO

9.22

9.55

10.3

08.

897.

837.

407.

99.

569.

647.

529.

919.

9410

.32

11.9

310

.53

14.4

88.

419.

328.

710

.32

11.6

Fe2O

32.

242.

811.

473.

853.

62.

193.

181.

910.

392.

553.

13.

275.

442.

73.

922.

682.

981.

663.

535.

445.

38M

nO0.

240.

260.

230.

220.

260.

160.

240.

190.

190.

160.

240.

20.

230.

210.

160.

250.

210.

230.

220.

230.

25M

gO16

.32

12.2

713

.22

16.2

311

.61

15.1

99.

007.

4010

.13

11.0

68.

656.

998.

478.

346.

757.

6110

.52

7.81

11.2

88.

478.

39C

aO9.

715.

9910

.74

8.75

10.0

88.

979.

948.

838.

6212

.12

10.4

69.

2212

.34

8.56

8.96

8.92

11.1

19.

187.

8012

.34

12.4

4N

a 2O

0.83

0.89

1.18

1.13

2.35

1.03

2.98

2.76

0.97

1.78

1.41

2.89

1.17

2.44

2.23

1.98

1.82

2.06

2.18

1.17

1.26

K2O

1.76

0.61

0.40

1.50

0.93

0.08

0.56

1.00

1.07

0.24

0.4

0.38

0.24

0.18

0.42

0.27

0.19

0.18

1.73

0.24

0.38

P 2O

50.

150.

200.

170.

020.

210.

210.

050.

100.

310.

070.

110.

090.

090.

10.

140.

270.

070.

430.

260.

090.

18L

OI

0.97

2.37

0.47

1.33

0.48

1.53

0.76

0.35

1.97

2.04

0.56

0.25

1.13

0.34

0.17

0.06

0.84

1.61

0.81

1.13

0.75

Tot

al98

.97

98.8

610

0.0

98.9

699

.11

99.3

098

.91

98.8

699

.64

99.2

710

0.7

99.0

110

0.55

98.7

098

.82

98.5

999

.29

100.

698

.94

100.

5598

.70

Sc30

5044

6548

3948

3144

4842

4772

5272

5066

5078

7279

V15

018

022

017

018

018

019

024

023

020

026

029

030

032

228

033

029

028

023

030

046

0C

r19

0012

0093

016

0077

017

0017

017

064

011

0040

025

073

110

150

9336

026

065

373

160

Ni

620

420

210

600

190

500

140

8113

028

014

088

7483

2914

017

021

011

074

72C

o64

4552

6542

5746

4942

5648

5062

4842

6552

4349

6253

Sr7

714

1018

036

350

940

200

150

9833

056

063

054

084

160

460

130

560

390

Rb

31.7

10.8

810

.413

.76

13.1

9N

.f.

N.f

.14

N.f

.3.

36N

.f.

2.55

4.69

N.f

.N

.f.

3.48

N.f

.N

.f.

22N

.f.

6.92

Y15

.55

30.7

316

.98

13.0

118

.24

18.5

311

1518

18.3

19.7

1920

.46

1418

31.3

523

23.4

122

1023

.86

Zr

8185

5132

5885

3173

7637

4069

2846

9813

051

8110

028

42N

b3.

587.

253.

041.

434.

245.

24–

––

2.2

5.05

2.97

2.95

––

10.9

5–

10.7

0–

–4.

37C

s3.

070.

42N

.f.

1.63

0.53

0.00

2–

––

0.68

N.f

.0.

030.

01–

–0.

09–

––

–0.

11B

a74

021

024

066

093

027

150

190

200

5035

120

6134

230

9613

5583

061

170

La

8.03

10.6

310

.34

2.29

9.58

16.1

9–

––

2.03

4.14

3.06

8.57

––

15.0

5–

22.9

3–

–6.

94C

e20

.00

24.1

823

.59

6.78

19.0

732

.78

––

–4.

9810

.32

8.29

19.4

6–

–34

.44

–44

.57

––

16.3

4Pr

2.61

2.78

2.67

1.14

2.38

4.01

––

–0.

821.

411.

232.

57–

–4.

58–

5.8

––

2.41

Nd

11.3

212

.01

11.3

75.

7010

.48

17.1

2–

––

4.44

6.72

5.84

11.3

4–

–19

.05

–25

.33

––

11.7

6Sm

3.02

3.78

3.01

1.71

3.11

4.12

––

–1.

682.

181.

873.

25–

–5.

06–

6.02

––

3.37

Eu

0.87

1.56

0.71

0.59

1.04

1.02

––

–0.

810.

420.

690.

83–

–16

9–

1.61

––

1.04

Gd

2.63

3.83

2.45

1.74

2.80

3.13

––

–2.

131.

942.

643.

34–

–5.

15–

4.20

––

3.61

Tb

0.44

0.72

0.37

0.31

0.48

0.48

––

–0.

400.

380.

520.

55–

–0.

80–

0.56

––

0.62

Dy

2.94

5.38

2.34

2.29

2.98

3.30

––

–3.

062.

673.

743.

3–

–5.

07–

3.38

––

3.99

Ho

0.64

1.31

0.50

0.51

0.61

0.73

––

–0.

720.

600.

850.

70–

–1.

05–

0.70

––

0.89

Er

1.79

3.52

1.43

1.54

1.69

2.08

––

–2.

131.

712.

311.

93–

–2.

85–

1.89

––

2.42

Tm

0.26

0.49

0.22

0.23

0.26

0.29

––

–0.

310.

270.

350.

30–

–0.

41–

0.29

––

0.36

Yb

1.50

2.93

1.34

1.32

1.54

1.67

––

–1.

891.

752.

051.

86–

–2.

40–

1.75

––

2.17

Lu

0.26

0.51

0.23

0.21

0.27

0.26

––

–0.

330.

300.

330.

32–

–0.

41–

0.30

––

0.36

Hf

2.39

2.10

1.54

0.88

1.04

1.78

––

–1.

061.

091.

432.

40–

–2.

70–

2.64

––

1.21

Ta

0.27

0.40

N.f

.0.

110.

200.

33–

––

0.16

N.f

.1.

010.

20–

–0.

79–

0.05

7–

–2.

06T

h2.

566.

651.

49N

.f.

2.93

6.83

––

–0.

060.

480.

10.

93–

–0.

30–

0.13

––

0.53

U0.

490.

85N

.f.

0.05

1.24

0.85

––

–0.

002

N.f

.0.

071.

38–

–0.

25–

––

–0.

38

(con

tinu

ed o

n ne

xt p

age)


Tab

le 1

(co

ntin

ued)

Com

po-

nent

Low

-Ti

met

avol

cani

cs

2/86

2a/8

631

b/86

32/8

634

a/86

37/8

637

v/86

38/8

638

a/86

38b/

8639

a/86

40/8

686

7/2

867/

986

7/10

869/

286

9/6

869/

8-1

869/

1686

9/18

SiO

250

.08

51.3

350

.68

51.0

951

.42

45.5

151

.72

46.2

049

.08

48.4

450

.12

49.9

848

.20

50.2

449

.49

51.0

749

.27

46.3

847

.71

49.6

0T

iO2

0.36

0.39

0.24

0.17

0.31

0.35

0.44

0.39

0.28

0.24

0.23

0.21

0.14

0.33

0.24

0.07

0.21

0.08

0.38

0.24

Al 2

O3

12.3

312

.77

12.6

713

.212

.62

12.1

613

.413

.30

12.4

214

.20

13.4

111

.51

11.5

714

.09

13.4

410

.41

12.7

514

.87

14.7

015

.20

FeO

7.26

7.25

7.05

7.88

7.43

6.48

7.9

8.31

7.16

8.79

7.50

7.51

5.71

6.67

5.91

5.23

6.28

5.94

7.58

6.97

Fe2O

32.

642.

452.

541.

512.

244.

492.

673.

542.

921.

172.

362.

143.

283.

123.

731.

604.

633.

341.

531.

79M

nO0.

180.

160.

160.

160.

150.

160.

150.

190.

160.

160.

170.

190.

230.

160.

150.

190.

190.

220.

210.

15M

gO12

.37

11.3

111

.43

10.3

311

.50

9.73

10.2

212

.45

13.7

810

.84

12.0

515

.27

13.6

410

.50

9.35

17.1

010

.58

13.2

713

.80

11.2

7C

aO10

.41

10.6

011

.59

11.9

911

.46

16.4

610

.74

12.4

19.

6112

.09

10.8

68.

8114

.61

11.4

114

.50

12.0

613

.04

12.7

811

.63

12.1

0N

a 2O

1.13

2.23

1.63

1.36

1.39

1.02

2.11

1.03

1.56

1.72

1.17

1.16

0.54

2.02

0.67

0.70

1.15

1.3

1.74

2.07

K2O

0.46

0.16

0.23

0.18

0.13

0.18

0.16

0.20

1.42

0.18

0.21

0.14

0.08

0.07

0.20

0.06

0.18

0.14

0.14

0.10

P 2O

50.

030.

040.

030.

040.

030.

050.

050.

040.

010.

040.

020.

020.

030.

030.

030.

010.

040.

020.

020.

03L

OI

1.86

0.37

0.89

1.57

0.39

2.47

0.82

1.48

0.86

1.64

1.08

2.41

1.45

0.82

1.9

0.90

2.76

1.07

1.04

0.56

Tot

al99

.12

99.0

699

.14

99.4

899

.08

99.0

610

0.38

99.5

499

.25

99.5

199

.18

99.3

599

.48

99.4

699

.61

99.4

099

.40

99.4

210

0.48

100.

08Sc

8475

3948

4351

4242

3839

3536

3578

6744

4467

4462

V19

021

020

022

020

022

022

021

018

020

019

020

013

019

019

011

011

013

020

021

0C

r63

063

092

749

072

068

058

078

084

088

069

012

0011

0065

079

019

0019

0061

041

036

0N

i26

024

029

023

024

025

020

029

025

030

024

036

029

018

023

046

046

055

027

023

0C

o53

5656

5756

5352

6949

5453

5853

5255

5252

6347

55Sr

4911

072

140

7213

013

015

077

9397

3722

014

012

036

3639

011

019

0R

b5.

86N

.f.

N.f

.N

.f.

N.f

.N

.f.

N.f

.1.

04N

.f.

8N

.f.

3.68

N.f

.2.

9N

.f.

N.f

.N

.f.

N.f

.N

.f.

N.f

.Y

11.4

86

15.3

78

13.4

15.2

36

66

6.2

5.54

10.5

65

57

116

Zr

817

516

1110

922

57

55

1214

119

914

4211

Nb

1.08

––

N.f

.–

–0.

072.

97–

––

0.96

0.22

1.06

––

––

N.f

.–

Cs

2.05

––

N.f

.–

–N

.f.

0.08

––

–1.

31N

.f.

0.06

––

––

N.f

.–

Ba

2316

2122

1212

1212

4614

1212

1234

1212

1221

3034

La

3.03

––

0.68

––

0.68

1.77

––

–0.

830.

563.

15–

––

–2.

13–

Ce

5.79

––

2.06

––

2.23

4.45

––

–2.

061.

275.

68–

––

–4.

70–

Pr0.

70–

–0.

32–

–0.

280.

64–

––

0.32

0.19

0.65

––

––

0.56

–N

d2.

72–

–1.

88–

–1.

833.

22–

––

1.51

0.94

2.56

––

––

2.52

–Sm

0.67

––

0.94

––

0.75

1.03

––

–0.

480.

320.

62–

––

–0.

71–

Eu

0.20

––

0.31

––

0.27

0.31

––

–0.

410.

150.

19–

––

–0.

81–

Gd

1.11

––

1.03

––

0.87

1.55

––

–0.

600.

471.

07–

––

–0.

64–

Tb

0.23

––

0.22

––

0.19

0.33

––

–0.

110.

100.

22–

––

–0.

10–

Dy

1.57

––

1.76

––

1.48

2.35

––

–0.

950.

811.

45–

––

–0.

75–

Ho

0.38

––

0.43

––

0.37

0.58

––

–0.

250.

190.

36–

––

–0.

17–

Er

1.09

––

1.33

––

1.16

1.79

––

–0.

820.

621.

02–

––

–0.

48–

Tm

0.17

––

0.23

––

0.19

0.27

––

–0.

130.

110.

16–

––

–0.

08–

Yb

1.09

––

1.54

––

1.31

1.67

––

–0.

770.

710.

97–

––

–0.

56–

Lu

0.18

––

0.29

––

0.26

0.28

––

–0.

140.

120.

16–

––

–0.

10–

Hf

0.34

––

2.25

––

1.28

1.03

––

–0.

02–

0.57

––

––

0.48

–T

a0.

02–

–N

.f.

––

N.f

.0.

21–

––

0.12

0.02

0.82

––

––

N.f

.–

Th

0.65

––

N.f

.–

–N

.f.

1.28

––

–N

.f.

N.f

.0.

57–

––

–0.

004

–U

0.21

––

N.f

.–

–N

.f.

0.35

––

–N

.f.

N.f

.0.

31–

––

–N

.f.

–

Not

e. N

.f.,

Con

tent

of

elem

ent

is b

elow

its

det

ectio

n lim

it. D

ash,

Not

det

erm

ined

.


with a wide variation in Mg# (66–84) and contents of FeO*(6.7–11.4 wt.%), and CaO (9–16 wt.%). There are alsohigh-Mg varieties (>12 wt.% MgO), which can be referred toas picrites (Fig. 4), but we do not consider them because alllow-Ti metavolcanics form common trends or fields notseparated into clusters on diagrams. The increase in FeO*,TiO2, and P2O5 contents with a Mg# decrease (Fig. 7) andthe direct correlation between the Ni and MgO contents(Fig. 8) point to fractionation involving olivine. The higher

CaO/Al2O3 ratios (0.8–1.2) relative to that of island-archigh-Mg rocks might point to a higher content of clinopy-roxene as compared with orthopyroxene in the rock protolith(Schuth et al., 2004).

The low-Ti metavolcanics form slightly depleted (relativeto N-MORB) multielemental patterns similar in contents ofmost incompatible trace elements to the primitive-mantle onesbut differing in the presence of distinct positive anomalies ofSr (Fig. 9). The low-Ti metavolcanics show positive or nega-tive anomalies of Zr but no anomalies of Ti; their (Nb/La)PMvalues vary from 0.32 to 1.6. The contents of Th in mostsamples are below its detection limit; therefore, it is difficultto judge the existence of a Nb anomaly. These metavolcanicsare characterized by both LREE-depleted ((La/Yb)N = 0.32–0.77) and weakly LREE-enriched ((La/Yb)N = 1.9–2.3) frac-tionation patterns similar to those of boninitic rocks (Fig. 10).There are also weak to distinct positive and negative Euanomalies.

High-Ti metavolcanics of the Urtagol Formation. Meta-basalts and metapicrites of this group are considered sepa-rately, because they often form different, not overlapping oronly partly overlapping clusters on elemental diagrams.

The rocks with MgO = 6.8–11.2 wt.%, TiO2 = 0.96–2.8 wt.%, and SiO2 = 43.1–54.5 wt.% are referred to asmetabasalts. As the MgO content decreases, the contents ofFeO* and TiO2 increase, which corresponds to olivine frac-tionation. The high-Ti metabasalts show a wide range of

Fig. 7. Variations in TiO2, FeO*, Na2O, and P2O5 contents depending on Mg# in metavolcanics of the Tunka terrane. Designations follow Fig. 3.

Fig. 6. SiO2–MgO (Le Bas, 2000) classification diagram for metavolcanics ofthe Tunka terrane. Metavolcanics: 1, high-Ti; 2, low-Ti.


contents of incompatible trace elements, whose fractionationpatterns lie between those of N-MORB and OIB (Fig. 9). Inhigh contents of LREE, Rb, and Ba their patterns are similarto the OIB ones but show both positive and negative anomaliesof Zr and Ti. The wide variations in (Nb/La)PM (0.3–1.2) and(Nb/Th)PM (0.4–4.3) are due to positive or negative anomaliesof Nb.

The high-Ti metabasalts are characterized by fractionatedLREE ((La/Yb)N = 0.8–9.4) and gentle HREE patterns andboth positive and negative anomalies of Eu.

High-Ti metapicrites include metavolcanics with 11.6–16.3 wt.% MgO, 46.6–52.2 wt.% SiO2, and 1.1–3.2 wt.%(Na2O + K2O) (Le Bas, 2000). As Mg# decreases, the contentsof TiO2 and FeO* increase. This points to differentiation withthe participation of olivine (Fig. 7).

The composition diagrams of all recognized groups ofmetavolcanics show no correlations between the contents ofcomponents (Fig. 8), which suggests the participation ofclinopyroxene in fractionation. The trend bend, most likely,indicates a change of fractionating phases during the Mg#decrease.

The rocks show fractionation patterns of incompatible traceelements symbatic with the OIB ones (Fig. 9), with distinctnegative anomalies of Ti and Nb and both negative andpositive anomalies of Zr as well as LREE, Ba, Rb, and Thenrichment. But in contrast to the OIB patterns, they show adistinct negative anomaly of Sr. The (Nb/La)PM and

(Nb/Th)PM values are lower than those of the primitive mantle(0.28–0.65 and 0.1–0.24, respectively). The high-Ti metapi-crites are characterized by fractionated LREE ((La/Sm)N =0.9–2.5) and flat HREE ((Gd/Yb)N = 1.1–1.5) patterns.

Discussion

Despite the described difference, the above groups ofmetavolcanics show some geochemical similarity. Each groupincludes high-Mg varieties as well as metavolcanics enrichedin LREE and depleted in Nb relative to La and Th, whichsuggests the suprasubductional nature of their parental melts.

The diagrams in Figs. 9, 10, and 11 suggest that the Tunkametavolcanics were, most likely, generated from a melt similarto the parental melt for N-MORB or more depleted, with adifferent contribution of enriched material to the magmageneration area.

The Th–Nb–Ce (La) correlations are the most appropriateto elucidate the mantle sources of volcanics, since they reflectits composition and the presence of the material of oceanicand continental crust. The contents of Th and Nb in the low-Timetavolcanics were determined in the fourth and sixth of eightsamples, and in the rest samples they were below theirdetection limits. The contents of Th in three samples of low-Timetavolcanics are close to those of IAB or are somewhathigher (Dorendorf et al., 2000), and in one sample it is

Fig. 8. Contents of CaO, Al2O3, Fe2O3, and Ni vs. contents of MgO in metavolcanics of the Tunka terrane. 1, high-Ti metabasalts; 2, high-Ti metapicrites; 3, low-Timetavolcanics. Composition of clinopyroxene is after Mashima (2005).


Fig. 9. Primitive-mantle-normalized (Sun and McDonough, 1989) pattern of incompatible elements for metavolcanics of the Tunka terrane. Average compositions ofN-MORB and OIB are after Sun and McDonough (1989); ARC, average composition of high-Mg basalt from the Klyuchevskoi Volcano, after Dorendorf et al.(2000).

Fig. 10. Chondrite-normalized (Sun and McDonough, 1989) REE patterns for metavolcanics of the Tunka terrane. Average compositions of N-MORB and OIB areafter Sun and McDonough (1989); ARC, average composition of high-Mg basalt from the Klyuchevskoi Volcano, after Dorendorf et al. (2000).


extremely low. The contents of Nb in four samples are closeto those in IAB, and in two samples they are lower. The(Nb/Th)PM values in four samples are <1 (0.20–0.31), i.e.,these rocks are enriched in Th relative to Nb, which is typicalof contaminated or subducted volcanics. The (Nb/La)PM valuein five of seven samples is low, 0.11–0.38. Higher (Nb/La)PMvalues (>1) (in two samples) accompanied by high (Th/La)PMvalues (La-depleted varieties) are, most likely, the result ofthe higher degrees of partial melting. As seen from thediagram in Fig. 11, the low-Ti metavolcanics are actuallycharacterized by higher degrees of melting, which decreasefor the high-Ti metabasalts. For comparison, the EPR picrites(Perfit et al., 1996) are shown, whose degree of partial meltingreaches 15–18%. The low Nb/U values of the low-Ti meta-volcanics (3–9), as compared with those of N-MORB (40–60),also suggest the presence of crustal material in the mantlesource.

The high-Ti metabasalts are characterized by Th/Nb, Th/Zr,and Ce/Nb ratios similar to or somewhat higher than those ofN-MORB (Fig. 12) but significantly different Nb/Zr values.Their composition points lie near the fields of picrites (Huangand Frey, 2005) and picritoids (Izokh et al., 2005) enrichedin the recycled material of oceanic crust. They have low(Th/La)PM (0.05–0.87), high (except for one sample)(Nb/Th)PM, and varying (Nb/La)PM (0.3–1.1) values. All thispoints to the Nb enrichment of the rocks and suggests thepresence of the recycled material of oceanic crust (subjectedto melting in the subduction zone) in the source of metavol-canics.

High-Ti metapicrites differ strongly in composition fromall other metavolcanics, which can hardly be explained onlyby differentiation processes or different degrees of melt-ing (Fig. 11). Being similar in Mg# to the low-Ti metavol-canics, they have higher LILE and LREE contents and showdistinct negative anomalies of Nb ((Nb/Th)PM = 0.09–0.24,(Nb/La)PM = 0.28–0.66) and, partly, Zr, which is typical ofrocks formed in subduction zone. The presence of a subductioncomponent is evidenced both by high Ce/Nb and Th/Nb ratiosand by high Th/Zr values (Fig. 12) pointing to the presenceof enriched fluid in the source of high-Ti metapicrites.

The obtained isotope ratios, together with the availablegeochemical data, agree with the drawn conclusions. The Ndisotope composition of the metavolcanics is lowly radiogenic(143Nd/144Nd = 0.512001–0.512298), probably because of thepresence of some enriched component in the mantle source.The above isotope ratios are close to those of the enrichedmagma (EM) characterized by a low ratio of Nd isotopes.Formation of this mantle component is commonly associated

Fig. 11. MgO and Zr/Y vs. TiO2/Al2O3 in metavolcanics of the Tunka terrane. Designations follow Fig. 8. Fields of picrites from the East Pacific Rise (EPR) are afterPerfit et al. (1996).

Fig. 12. Th/Zr–Nb/Zr and Th/Nb–Ce/Nb diagrams for metavolcanics of theTunka terrane. Designations follow Fig. 8. Fields of the Emeishan picritoids(Izokh et al., 2005) and Koolau lavas (Huang and Frey, 2005) and back-arcbasin basalts (Wang et al., 2007) are shown. DMM, component of depletedmantle source MORB, RSC, residual component of recycled oceanic slab, SDC,subduction component of island-arc magmatism, after Saunders et al. (1988).


with the addition of crustal material as subducted sediments.It is difficult to establish the composition of enriched mantlesource (EM1 or EM2) from the available isotope data.However, it is worth noting that the low negative values ofεNd (from –4.8 to –9.6) (Table 2) are reached when ancientcrustal material is added to the magma generation zone.

There are different possible reasons for the presence ofgeochemically different types of metavolcanics. One is theirmelting from different mantle sources. Probably, a Nb-richrecycled oceanic crustal (slab) material was the enrichedsource of the high-Ti metabasalts, and a LILE- and LREE-richsubduction fluid, of the high-Ti metapicrites. The enrichmentof the low-Ti metavolcanics is, most likely, related to theassimilation of continental crustal material.

Let us dwell in more detail on the metavolcanics with thehighest Mg# values (68–84) and high contents of Cr and Ni,which point to their compositional similarity to primary melts.These rocks are characterized by low (Sm/Yb)N valuesindicating the high degrees of melting and/or absence of garnetfrom restite. A specific feature of the high-Mg metavolcanicsis a high content of silica. In modern island-arc systems,high-Mg high-Si melts formed at shallow depths (~50 km)with high temperatures (>1300 ºC) and high degrees ofmelting of the mantle source (Falloon and Danyushevsky,2000). High degrees of melting are reached either underextreme conditions in subduction zones, which favor decom-pression melting (oblique subduction, jump of subductionzone, transverse extension, etc.), or under heating of thesuprasubductional mantle by plume (Izokh et al., 2010).

Subduction settings with volcanics moderately enriched inLREE and weak positive or negative anomalies of Nb, P, andTi are referred to as back-arc ones (Manikyamba et al., 2009).The high-Fe picrites, picrite basalts, and picrites of the EarlyProterozoic Pechenga–Varguza belt, Kola Peninsula (Sharkovand Smolkin, 1997), and modern picrites of Kumejima Islandand Ryukyu Island (Junichi, 2004) are directly related to theback-arc spreading. The association with carbonate and ter-rigenous rocks, the coexistence with serpentinites in the rockunit, and the geochemical features of high-Mg metavolcanicsshow that the latter can mask the ancient spreading zone ofthe marginal basin within the Tunka terrane. Probably,depleted volcanics (low-Ti metavolcanics) formed at some-what earlier stages, during the maximum opening of theback-arc basin with active spreading. According to Ohara(2006), the low-Ti metavolcanics characterized by higherdegrees of partial melting (low contents of Zr and reducedZr/Y values) can actually correspond to the higher rates(7–8 cm/year) of the initial spreading of the back-arc basin.

The enriched volcanic rocks might have formed at the laterstages of spreading, when the extension setting was changedby a subduction one, marked as high-Ti metapicrites with highcontents of LILE and LREE and distinct negative Nb and Tianomalies.

Thus, taking into account one of the latest interpretationsof the structure of the Sayan–Baikal Fold Area (Zorin et al.,2009), we can regard the Tunka terrane as part of the back-arcbasin of the active island-arc margin and the primitivehigh-Mg metavolcanics as markers of back-arc paleospread-ing.

Conclusions

The compositional similarity of high-Mg metavolcanics inthe central zone of the Tunka bald mountains to rocks ofsubduction nature, the presence of basaltoids both with distinctnegative and with positive anomalies of Nb, and the coexis-tence of these rocks with serpentinites in the rock unit markthe ancient spreading zone of the back-arc basin within theTunka terrane.

Recycled Nb-rich oceanic crustal material (RSC) was asource of Nb for the high-Ti metabasalts, and a subductionfluid enriched in LREE and Th relative to Nb was a sourceof these elements for the high-Ti metapicrites. The enrichmentof the low-Ti metavolcanics, formed probably at the earlystages of the basin opening, was due to the contamination ofmelt by continental crustal material.

The different positions of low- and high-Mg basalts in thesection of the Tunka terrane, their association with andesites(Tolta Formation) and ultrabasic rocks (Urtagol Formation),respectively, and the increase in the degree of melt differen-tiation upsection reflects the evolution of volcanism.

References

Belichenko, V.G., Boos, R.G., 1988. Evolution of the Earth’s crust on thesouthern framing of the Siberian Platform in the Paleozoic, in: Evolutionof the Earth’s Crust in the Precambrian and Paleozoic (Sayan–BaikalMountainous Area) [in Russian]. Nauka, Novosibirsk, pp. 91–115.

Belichenko, V.G., Reznitskii, L.Z., Geletii, N.K., Barash, I.G., 2002. TheTuva–Mongolian massif and its eastern framing, in: Geology, Geochem-istry, and Geophysics at the Boundary of the 20th and 21st Centuries.Proc. All-Russ. Sci. Conf. [in Russian]. Irkutsk, pp. 19–21.

Bogatikov, O.A. (Ed.), 1983. Igneous Rocks. Classification. Nomenclature.Petrography [in Russian]. Nauka, Moscow, Part 1.

Bogatikov, O.A., Mikhailov, N.P., Gon’shakova, V.I. (Eds.), 1981. Classifi-cation and Nomenclature of Igneous Rocks [in Russian]. Nedra, Moscow.

Table 2. Sm-Nd isotope data for metavolcanics of the Tunka terrane

Sample T, Ma Content, ppm 147Sm/144Nd 143Nd/144Nd ± 2σ εNd(T)

Sm Nd

43/86 600 3.02 11.32 0.16058 0.512006 ± 5 –9.6

869/7 600 6.02 25.33 0.14305 0.512001 ± 5 –8.3

865/17 600 3.25 11.34 0.17250 0.512298 ± 13 –4.8


Boos, R.G., 1988. The Paleozoic Strata of the Tunka Bald Mountains, EastSayan. PhD Thesis [in Russian]. IZK SO AN SSSR, Irkutsk.

Boos, R.G., 1991. The Paleozoic Strata of the Tunka Bald Mountains, EastSayan [in Russian]. Nauka, Novosibirsk.

Efremov, S.V. (Ed.), 1997. Classification of Igneous Rocks and Dictionaryof Terms. Recommendations of the Subcommission on Systematics ofIgneous Rocks of the International Union of Geosciences [in Russian].Nedra, Moscow.

Dorendorf, F., Wiechert, U., Wörner, G., 2000. Hydrated sub-arc mantle: asource for the Kluchevskoy volcano, Kamchatka/ Russia. Earth Planet.Sci. Lett. 175, 69–86.

Falloon, T.J., Danyushevsky, L.V., 2000. Melting of refractory mantle of 1.5,2 and 2.5 GPa under anhydrous and H2O-undersaturated conditions:implications for the petrogenesis of high-Ca boninites and influence ofsubduction components of mantle melting. J. Petrol. 41, 257–283.

Herzberg, C., O’Hara, M.J., 1998. Phase equilibrium constraints on the originof basalts, picrites, and komatiites. Earth Sci. Rev. 44, 39–79.

Huang, S., Frey, F.A., 2005. Recycled oceanic crust in the Hawaiian Plume:evidence from temporal geochemical variations within the Koolau Shield.Contrib. Mineral. Petrol. 149, 556–575.

Ichiyama, Y., Ishiwatari, A., 2005. HFSE-rich picritic rocks southwesternJapan. Contrib. Mineral. Petrol. 149, 373–387.

Izokh, A.E., Polyakov, G.V., Tran Trong Hoa, Balykin, P.A., Ngo Thi Phuong,2005. Permian–Triassic ultramafic-mafic magmatism of northern Vietnamand southern China as expression of plume magmatism. Russian Geologyand Geophysics (Geologiya i Geofizika) 46 (9), 922–931 (942–951).

Izokh, A.E., Vishnevskii, A.V., Polyakov, G.V., Kalugin, V.M., Oyunchi-meg, T., Shelepaev, R.A., Egorova, V.V., 2010. The Ureg Nuur Pt-bearingvolcanoplutonic picrite–basalt association in the Mongolian Altay asevidence for a Cambrian–Ordovician Large Igneous Province. RussianGeology and Geophysics (Geologiya i Geofizika) 51 (5), 521–533(665–681).

Junichi, I., Hiroaki, S., Keiichi, S., 2004. Petrogenesis of picrite fromKumejima Island, Ryukyu Island, implications for the opening of theOkinawa Trough. Japan Earth Planet., Sci. Joint Meeting. Abstr. (K 073P 010).

Kerr, A.C., Marriner, G.F., Arndt, N.T., Tarney, J., Nivia, A., Saunders, A.D.,Duncan, R.A., 1996. The petrogenesis of Gorgona komatiites, picrites andbasalts: new field, petrographic and geochemical constraints. Li-thos 37, 245–260.

Larsen, L.M., Pedersen, A.K., Sundvoll, B., Frei, R., 2003. Alkali picritesformed by melting of old metasomatized lithospheric mantle: ManîtdlatMember, Vaigat Formation, Paleocene of West Greenland. J. Pet-rol. 44, 3–38.

Laz’ko, E.V., Sharkova, E.V. (Eds.), 1988. Igneous Rocks. Ultrabasic Rocks[in Russian]. Nauka, Moscow.

Le Bas, M.J., 2000. IUGS reclassification of the high-Mg and picritic volcanicrocks. J. Petrol. 41, 1467–1470.

Manikyamba, C., Kerrich, R., Khanna, T.C., 2009. Enriched and depleted arcbasalts, with Mg-andesites and adakites: a potential paired arc–back-arcof the 2.6 Ga Hutti greenstone terrane, India. Geochim. Cosmochim.Acta 73, 1711–1736.

Mashima, H., 2005. Partial melting controls on the northwest Kyushu basaltsfrom Saga-Futagoyama. Island Arc 14, 165–177.

Neelov, A.N., 1980. Petrochemical Classification of Metamorphosed Sedi-mentary and Volcanic Rocks [in Russian]. Nauka, Leningrad.

Ohara, Y., 2006. Mantle process beneath Philippine Sea back-arc spreadingridges: a synthesis of peridotite petrology and tectonics. IslandArc 15, 119–129.

Perfit, M.R., Fornari, D.J., Ridley, W.I., Kirk, P.D., Casey, J., Kastens, K.A.,Reynolds, J.R., Edwards, M., Desonie, D., Shuster, R., Paradis, S., 1996.Recent volcanism in the Siqueiros transform fault: picritic basalts andimplications for MORB magma genesis. Earth Planet. Sci. Lett. 141, 91–108.

Reznitskii, L.Z., Sal’nikova, E.B., Barash, I.G., Belichenko, V.G., Glebovit-skii, V.A., Kotov, A.B., Kovach, V.P., Yakovleva, S.Z., Fedose-enko, A.M., 2007. Upper age boundary of the accretion of terranes in thenorthwestern part of the eastern segment of the Central Asian Foldbelt.Dokl. Earth Sci. 414 (4), 548–551.

Saunders, A.D., Norry, M.J., Tarney, J., 1988. Origin of MORB andchemically-depleted mantle reservoirs: trace element constraints. J. Pet-rol. 415–445.

Schuth, S., Rohrbach, A., Münker, C., Ballhaus, C., Garbe-Schönberg, D.,Qopoto, C., 2004. Geochemical constraints on the petrogenesis of arcpicrites and basalts, New Georgia Group, Solomon Islands. Contrib.Mineral. Petrol. 148 (3), 288–304.

Sharkov, E.V., Smolkin, V.F., 1997. The early Proterozoic Pechenga–VarzugaBelt: a case of Precambrian back-arc spreading. Prec. Res. 82, 133–151.

Shkol’nik, S.I., Reznitskii, L.Z., Belichenko, V.G., Barash, I.G., 2009.Geochemistry, petrogenesis, and geodynamic typification of metavolcanicsof the Tunka terrane (Baikal–Hövsgöl region). Russian Geology andGeophysics (Geologiya i Geofizika) 50 (9), 779–788 (1013–1024).

Shkol’nik, S.I., Belichenko, V.G., Reznitskii, L.Z., Barash, I.G., 2011. Afragment of back-arc paleospreading in the Tunka terrane. Dokl. EarthSci. 436 (1), 61–65.

Sun, S., McDonough, W.F., 1989. Chemical and isotopic systematics ofoceanic basalts: implications for mantle composition and processes, in:Saunders, A.D., Norry, M.J. (Eds.), Magmatism in Oceanic Basins. Geol.Soc. London. Spec. Pub. 42, 313–345.

Tian, W., Campbell, I.H., Allen, C.M., Guan, P., Pan, W., Chen, M., Yu, H.,Zhu, W., 2010. The Tarim picrite–basalt–rhyolite suite, a Permian floodbasalt from northwest China with contrasting rhyolites produced byfractional crystallization and anatexis. Contrib. Mineral. Petrol. 160, 407–425.

Tsikouras, B., Pe-Piper, G., Piper, D.J.W., Hatzipanagiotou, K., 2008. Triassicrift-related komatiite, picrite and basalt, Pelagonian continental margin,Greece. Lithos 104, 199–215.

Wang, Y.J., Zhao, G.C., Fan, W.M., Peng, T.P., Sun, L.H., Xia, X.P., 2007.LA-ICP-MS U-Pb zircon geochronology and geochemistry of Paleopro-terozoic mafic dykes from western Shandong Province: implication forbask-arc basin magmatism in the Eastern Block, North China Craton. Prec.Res. 154, 107–124.

Zorin, Yu.A., Sklyarov, E.V., Belichenko, V.G., Mazukabzov, A.M., 2009.Island arc–back-arc basin evolution: implications for Late Riphean–EarlyPaleozoic geodynamic history of the Sayan–Baikal folded area. RussianGeology and Geophysics (Geologiya i Geofizika) 50 (3), 149–161(209–227).

Editorial responsibility: A.E. Izokh


Documents

High-magnesium picrite–basalt associations of the Tunka terrane (Baikal–Hövsgöl region) as an indicator of the back-arc basin spreading