Granulites of the South Muya block (Baikal-Muya Foldbelt): Age of metamorphism and nature of protolith

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ScienceDirect Russian Geology and Geophysics 57 (2016) 451–463 www.elsevier.com/locate/rgg

Granulites of the South Muya block (Baikal–Muya Foldbelt): Age of metamorphism and nature of protolith S.Yu. Skuzovatov a,*, E.V. Sklyarov b,c, V.S. Shatsky a,d,e, K.-L. Wang f, K.V. Kulikova g, O.V. Zarubina a a

A.P. Vinogradov Institute of Geochemistry, Siberian Branch of the Russian Academy of Sciences, ul. Favorskogo 1a, Irkutsk, 664033, Russia b Institute of Earth’s Crust, Siberian Branch of the Russian Academy of Sciences, ul. Lermontova 128, Irkutsk, 664033, Russia c Far Eastern Federal University, ul. Sukhanova 8, Vladivostok, 690950, Russia d V.S. Sobolev Institute of Geology and Mineralogy, Siberian Branch of the Russian Academy of Sciences, pr. Akademika Koptyuga 3, Novosibirsk, 630090, Russia e Novosibirsk State University, ul. Pirogova 2, Novosibirsk, 630090, Russia f Institute of Earth Sciences, Academia Sinica, 128, Academia Road, Sec. 2, Nangang, Taipei, 11529, Taiwan g Institute of Geology, Komi Science Center, Ural Branch of the Russian Academy of Sciences, ul. Pervomaiskaya 54, Syktyvkar, 167982, Russia Received 19 November 2014; accepted 4 March 2015

Abstract High-pressure mafic granulites and garnet pyroxenites occur within the South Muya block as boudins or lenses among metamorphic rocks of the Kindikan Group. Their primary minerals crystallized at 670–750 °C and 9.5–12.0 kbar. Granulite metamorphism peaked at 630 Ma, according to LA-ICP-MS U–Pb zircon ages. Judging by their major- and trace-element compositions and Hf isotope ratios in zircons, the South Muya granulites were derived from differentiated within-plate basalts, which, in turn, resulted from melting of juvenile mantle source and Meso- or Paleoproterozoic crust. The events of granulite and eclogite metamorphism in the South and North Muya blocks, respectively, were coeval and the two blocks were spatially close to each other at the onset of Late Baikalian subduction and collision events. © 2016, V.S. Sobolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. Keywords: granulites; garnet pyroxenites; zircon; U–Pb age; continental subduction; Central Asian Orogenic Belt; South Muya block

Introduction The Muya segment of the Baikal–Muya Foldbelt comprises the Early Precambrian Muya block interpreted as a cratonic terrane (Bulgatov and Gordienko, 1999). It looks like a sheet bounded by thrusts in seismic images (Bulgatov, 1988), but its tectonic position remains controversial. Pressures and temperatures inferred from mineral parageneses of eclogitic boudins and lenses among rocks of the Dzhaltuk and Osinovka groups of the North Muya block previously timed as Archean or Early Proterozoic correspond to high-pressure metamorphism during continental subduction in the Baikal–Muya belt (Avchenko et al., 1989; Shatskii et al., 2014; Shatsky et al., 2012, 2015). On the other hand, high-pressure parageneses within the Paleoproterozoic Kindikan metamorphic complex

* Corresponding author. E-mail address: [email protected] (S.Yu. Skuzovatov)

of the South Muya block are poorly studied and may owe their origin to active stress metamorphism of basaltic rocks in zones of large faults (Doronina and Sklyarov, 1995; Grudinin and Menshagin, 1988). The highest grades in the South Muya block were inferred for diaphtoritic and granitic eclogite-like garnet-clinopyroxene-plagioclase-quartz-feldspar rocks which underwent peak metamorphism at 650 °C and 12 kbar (Bozhko et al., 1999). Thus, it remains unknown when and how the two blocks in the Muya terrane became juxtaposed and at which PT conditions they were subducted and exhumed. Recent geochronological studies of the Baikal–Muya rocks led to considerable revision to the previous timing of the Muya terrane as Early Precambrian. Rytsk et al. (2007, 2011) obtained Neoproterozoic ages for the metamorphic basement of the Baikal–Muya belt (0.6–1.0 Ga), while dating of the North Muya eclogite-gneiss complex (Shatsky et al., 2012) placed high-pressure metamorphism at 630 Ma. Of special interest in this respect are isotope ages of high-pressure rocks in the South Muya block that occur among the Kindikan

1068-7971/$ - see front matter D 201 6, V.S. So bolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. http://dx.doi.org/10.1016/j.rgg.201 + 6.03.007

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Group gneisses, schists, and calciphyres (Doronina and Sklyarov, 1995).

Samples and methods Metamorphic rocks were sampled in the area of Serebryakovsky and Dlinnyi (Long) Creeks within the South Muya block (Fig. 1). Rock-forming and accessory phases were identified using a JEOL JXA8200 electron-microprobe analyzer, at an accelerating voltage of 20 kV (analyst L.A. Pavlova). Major oxides were determined on a Bruker S4 Pioneer wavelength dispersive X-ray fluorescence spectrometer. Both

EPMA and XRF studies were performed at the share-use Center of Isotope-Geochemical studies (Institute of Geochemistry SB RAS, Irkutsk, Russia). Trace elements were measured by ICP-MS employing fusion with lithium metaborate, on a high-resolution Agilent 7700s quadrupole mass spectrometer at the Baikal Center for Nanotechnologies of the Irkutsk National Research Technological University. The ages and Hf isotope compositions of zircons were determined at the Institute of Earth Sciences, Academia Sinica (Taipei, Taiwan). In situ U–Pb dating was performed on a high-resolution Element XR ICP mass spectrometer with the Photon-Machines ANALYTE 193 nm ArF excimer laser ablation system (50 µm spot diameter). For estimating the Hf

Fig. 1. Simplified geology of the central Muya zone in the Baikal–Muya belt (Rytsk et al., 2011). 1, Quaternary sediments; 2, layered intrusions; 3, diorites and granitoids (Lesnoy complex), 4, Ust’-Kelyan basalt-rhyolite complex; 5, Param volcanic-sedimentary complex; 6, Dzhaltuk and East Gorbylok metamorphic complexes; 7, synmetamorphic gneiss-granite Ilier complex; 8, Nadporozhnaya volcanic-sedimentary and Shaman carbonate-schist complexes; 9, amphibolites and amphibole schists of Param–Shaman zone; 10, Param ultramafic complex; 11, Paleozoic granitoids of Barguzin complex; 12, Kindikan metamorphic complex; 13, Muya gabbro-diorite-plagiogranite complex; 14, Cambrian–Late Vendian clastic-carbonate complex; 15, Padrin volcanoplutonic complex; 16, Early Vendian Tallain gabbro-diorite-plagiogranite and Zaoblachnyi gabbro-gabbronorite complexes; 17, Karalon–Mamakan basalt-rhyolite complex; 18, Zhanok volcanoplutonic complex; 19, Siberian craton; 20, Osinovka amphibolite complex; 21, tectonic sutures. Box frames study area.

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Fig. 2. Textures and structures of analyzed samples (BSE images). a, b, Mafic granulites; c, garnet pyroxenite; d, poikilitic ingrowths in clinopyroxene from garnet pyroxenites.

isotope composition, a similar laser ablation system and a NuPlasma multicollector ICP mass spectrometer were applied to the same points, at the MC-ICP-MS parameters as reported by Griffin et al. (2000b). For more details of the analytical procedures see (Lan et al., 2009).

Results Petrography and mineralogy. The metamorphic complex within the Kindikan basement inlier consists of granulite and amphibolite facies gneisses; garnet-pyroxene-quartz-plagioclase granulites (including carbonate-bearing varieties); calciphyres; plagioclase-biotite-amphibole mafic rocks; and amphibolites. The complex comprises massive blocks consisting of amphibolites, garnet amphibolites, and garnet pyroxenites, mainly in cores of large folds, and also contains strongly altered rocks with residual primary minerals typical of garnet peridotite. Most granulites have granoblastic textures and massive structures, with matrix mineral grains from 0.1 to 0.4 mm (Fig. 2a), while some granulites are more or less strongly banded (samples Mu-10-24 and Mu-10-48). Mineral assemblages in granulites correspond to granulite, amphibolite, and

greenschist facies. Early parageneses include garnet (25%), oligoclase An18-22 (20%), K-feldspar (15%), clinopyroxene (15%), quartz (10%), sporadic orthopyroxene grains, and accessory ilmenite (up to 5% in a few cases), rutile, and pyrrhotite. Garnet is an almandine-grossular-pyrope variety, with Fe# = 0.84–0.89 (Fig. 3). Garnets have up to 0.7 mm grain sizes, are subhedral or anhedral, optically homogeneous, and most of them lack core-rim major-element zoning. In some cases, there is reverse zoning with pyrope (XPrp) decreasing and grossular (XGrs) increasing rimward, which holds in euhedral garnet grains in the granulite groundmass. Garnets enclose quartz, ilmenite, K-feldspar (Fig. 2b), as well as zircon and fluorapatite, which is present also in the rock matrix. Anhedral clinopyroxene occurs as salite (Fig. 4) and contains up to 6 mole% of jadeite; K-feldspar of the matrix is often perthitic. K-feldspar inclusions in garnet occur either as isolated phases or enter quartz-feldspar intergrowths. This paragenesis is preserved the best in massive granulites (Mu10-24, Mu-10-82 and Mu-10-124). Sample Mu-10-46 bears quite abundant euhedral biotite. Later parageneses are ferrous amphiboles (ferrous hornblende, ferropargasite, ferroedenite) and actinolite after clinopyroxene, as well as biotite, chlorite substituting for garnet mainly, and titanite after rutile. Clinopyroxene is remnant in Mu-10-46 and Mu-10-48 which contain

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Fig. 4. Compositions of pyroxenes in analyzed samples. Legend same as in Fig. 3.

Fig. 3. Compositions of garnets in analyzed granulites, pyroxenites, and peridotites. 1, granulite; 2, garnet pyroxenite, Serebryakovsky Creek; 3, garnet ultramafics, Dlinnyi Creek.

later amphibole and biotite as predominant mafic minerals while feldspars are strongly sericitized. Remnant garnet pyroxenites found in amphibolite boudins have granonematoblastic textures and massive structures. They consist of almandine-pyrope-grossular garnet (Fig. 3), diopside, pargasite (with Fe# lower than in amphibole from granulites), less often edenite, pyrrhotite, late epidote, chlorite, and actinolite (Fig. 2c). Subhedral and anhedral garnet grains, 0.3–0.7 mm, are slightly zoned, with 0.8–1.8% CaO, lower in the rim than in the core, and enclose clinopyroxene (mostly in grain cores), quartz, rutile, and pyrrhotite. Unzoned or weakly zoned clinopyroxene (diopside) contains 1.5–6.7% Al2O3 and 0.3–1.1% Na2O. Diopside in garnet pyroxenites is remarkable by poikilitic ingrowths withplagioclase partly substituted by an aggregate of albite, K-feldspar, epidote, and chlorite (Fig. 2d); pargasite-edenite amphibole substitutes for diopside (see Table 1 for representative analyses). Garnet-pyroxene rocks west of the main study area (in the vicinity of the Dlinnyi Creek) have notably coarser grain sizes and are more strongly altered. Their remnant primary minerals (up to 40 vol.%) are coarse (up to several mm) anhedral diopsides and anhedral pyrope-almandine-grossular segregations. Garnet is the most highly magnesian among the analyzed samples (Fe# 0.44–0.50) while the grossular component varies from 8 to 27 mole%. Diopside contains 1.2–2.9 wt.% Na2O and 1.8–4.2% Al2O3 (occasionally up to 8.9%); the jadeite component reaches 12%, while the tschermakite (Ca-Tschermak molecule) component is from 1.6 to 5.3% and reaches 12% in diopside having the highest Al2O3. Minerals occupying the greatest part of the rock volume experienced retrograde metamorphism, including low-temperature alteration, and are mostly chlorite, ferrous antigorite, as well as pargasite (after diopside) and later actinolite. Anhedral garnet segregations, like the whole rock, are cut by a dense web of thin chlorite veins. Garnets are locally replaced by epidote, with small (to

20 µm) hercynite particles. Much finer hercynites (within 1 µm) occur in most strongly chloritized zones of the groundmass. Abundant chlorite and serpentine may substitute for primary orthopyroxene and olivine, which allows interpreting the rocks as garnet peridotite, though no residual magnesian minerals have been identified in them. Note that altered peridotites in the Dlinnyi Creek area coexist with garnet pyroxenites similar to those of the Serebryakovsky Creek locality. Thermobarometry. Equilibrium temperatures for the Mu10-124 sample of garnet-clinopyroxene-plagioclase-K-feldspar-quartz granulite obtained by the classical geothermobarometry applied to garnet rim and matrix clinopyroxene are 720–740 °C (Ellis and Green, 1979). The Krogh Ravna garnet-clinopyroxene Fe2+-Mg geothermometer gives lower values and a larger range: 660–690 °C (Ravna, 2000). The pressure estimated with the geobarometer of Kohn and Spear (1990) is 10.5 kbar. PT calculations using TWQ software (Berman, 2007) show that the primary mineral assemblage of Mu-10-124 formed at T = 750 °C and P = 9.5 kbar. The equilibrium pressures and temperatures are 690–710 °C (Ellis and Green, 1979) and 620–640 °C (Ravna, 2000) at 11.5– 12 kbar for the Mu-10-82 granulite sample and 670–700 °C and 610–650 °C at 10.5 kbar, respectively, for Mu-10-24. These PT values fit the ranges 500–800 °C and 5–15 kbar suggested earlier for the Kindikan Group granulites (Doronina and Sklyarov, 1995). The equilibrium temperatures calculated for the compositions of remnant garnet and diopside (spatially close but not in contact), from the garnet ultramafic rock subjected to retrograde metamorphism (Mu-14-16, Dlinnyi Creek) are 700 to 830 °C (Ellis and Green, 1979) or 580– 710 °C according to Ravna (2000). The equilibrium pressures obtained from jadeite percentages in pyroxene from Mu-14-16 (Holland, 1980) reach 10 kbar, but they may be interpreted as minimum values in the absence of quartz. The temperatures calculated for pairs of garnet and diopside rims from the garnet pyroxenite Mu-10-80 are, respectively, 720–750 °C (Ellis and Green, 1979) and 640–670 °C (Ravna, 2000) but are higher for diopside inclusions and adjacent zones of garnet cores: 770–780 °C and 680–690 °C, obtained with the geothermometers of Ellis and Green (1979) and Ravna (2000), respectively.

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Table 1. Representative analyses of minerals from South Muya granulites and garnet pyroxenites Component

Mu-10-82 Grt 1

Mu-14-16 Cpx 1 Ap

Amp Grt 2

core rim

Cpx 2 Bt

Ilm

Kfs

Grt 3 Incl

core rim

SiO2, wt.% 38.1 38.6 51.4

Grt 1 Cpx 1 Amp Ep

Sp

Cpx 2 Grt 2

Kfs

0.18 42.7 37.9 38.2 51.2

35.7 64.5 –

64.3 37.2 64.0 40.1 53.5

42.6 39.3

2.31 54.4

40.3

3.53 –

–

–

TiO2

–

0.12

–

1.89 –

Al2O3

20.3 19.7 1.20

–

10.5 20.1 19.8 1.31

–

–

–

Ab

–

0.11

1.95 –

–

13.9 21.1 –

52.0

17.8 19.7 17.3 20.9 3.55

–

–

–

0.28

14.6 18.3

58.1 2.00

0.24

20.4

–

–

–

–

Cr2O3

–

–

–

FeO

32.6 33.1 17.0

0.17 22.6 31.7 32.6 16.9

27.3 0.00 49.1

–

33.2 0.80 19.6 4.54

7.80 13.1

34.4 4.57

20.2

MnO

0.86 0.91 0.14

–

–

–

0.42

–

0.89 –

0.54 –

–

0.65 –

0.64

MgO

3.01 2.41 9.87

–

7.17 3.10 2.79 9.82

–

–

2.44 –

9.51 14.6

15.2 0.18

5.09 16.0

12.2

7.24 –

10.1 21.7

–

–

–

–

–

0.79 0.82 0.11

–

–

–

6.80 –

–

–

–

–

2.67

–

CaO

6.83 6.88 21.2

54.0 11.1 7.72 7.43 21.8

–

3.60 –

–

11.6 26.2

0.26 22.2

6.37

Na2O

–

–

0.75

–

2.02 –

–

0.66

–

11.1 –

0.97 –

0.44 –

2.09

3.92 –

–

1.34

–

K2O

–

–

–

–

1.93 –

–

–

11.4 0.45 –

15.4 –

15.9 –

–

0.80 –

–

–

–

P2O5

–

–

–

41.6 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

F

–

–

–

3.53 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

Total

101.7 101.6 101.7

99.5 99.9 101.4 101.7 101.9 98.6 100.7 101.4

Si, ppm

3.01 3.06 1.96

0.01 6.48 3.00 3.02 1.95

Ti

–

0.00

–

0.22 –

Al

1.89 1.84 0.05

–

1.89 1.88 1.85 0.06

1.27 1.10 –

0.98 1.87 0.96 1.86 0.15

2.47 1.71

1.90 0.09

1.81

Cr

–

–

–

–

–

–

–

–

Fe

2.15 2.19 0.54

0.01 2.86 2.10 2.16 0.54

1.77 0.00 1.03

0.00 2.23 0.03 1.24 0.14

0.94 0.87

0.80 0.14

1.27

Mn

0.06 0.06 0.00

–

–

0.00 –

0.01

–

0.06 –

0.03 0.00

–

0.18

0.02 0.00

0.04

Mg

0.35 0.28 0.56

–

1.62 0.37 0.33 0.56

0.79 –

–

–

0.29 –

1.07 0.79

3.25 0.02

0.21 0.86

1.37

Ca

0.58 0.58 0.87

4.95 1.81 0.66 0.63 0.89

0.00 0.17 –

0.00 0.62 0.00 0.82 0.85

1.78 2.22

0.01 0.86

0.51

Na

–

–

0.06

–

0.59 –

–

0.05

0.00 0.95 –

0.09 –

0.04 –

0.15

1.09 –

–

0.09

–

K

–

–

0.00

–

0.37 –

–

0.00

1.13 0.03 –

0.92 –

0.95 –

0.00

0.15 –

–

0.00

–

P

–

–

–

3.01 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

F

–

–

–

0.95 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

Total

8.05 8.02 4.04

Component

–

–

0.00

–

–

0.00

–

0.00

0.05 0.05 0.00

8.95 15.8 8.06 8.05 4.04

98.5 100.7 98.5 100.8 100.2 98.4 99.8

100.8 100.7 100.1

2.78 2.85 –

3.01 3.00 3.01 3.03 1.95

6.14 3.11

0.06 1.97

3.04

0.21 –

–

0.21 –

–

–

–

–

7.95 5.09 2.02

Mu-10-124 Grt 1

0.98

–

–

–

–

–

0.01

–

0.00

5.00 8.07 5.00 8.04 4.04

–

16.0 8.10

0.01

0.00

2.99 4.02

8.05

Mu-10-24 Incl

Incl

Incl

core

rim

Kfs

Ilm

Ap

SiO2

38.8

38.3

63.9

–

0.94

TiO2

–

–

–

50.20 –

Amp Ab

Kfs

Bt

Cpx

Grt 1 core

Cpx

Amp Ap

Ilm

Ab

Kfs

rim

Grt 2 core

41.0

64.2

63.8

36.4

51.9

38.7

38.5

52.2

42.0

–

–

1.87

–

–

2.00

0.16

–

–

0.16

2.08

–

50.78 –

63.6

Incl rim

Bt

Kfs

63.8

39.2

38.9

63.6

38.2

–

–

–

–

–

Al2O3 19.52 19.26 17.94 –

–

10.60 21.33 17.82 12.74 1.60

19.52 19.47 1.59

10.38 –

–

21.61 17.81 19.69 19.54 18.50 14.45

Cr2O3 –

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

47.66 0.10

–

29.84 30.34 0.57

30.60

–

1.09

–

–

1.52

1.55

–

0.12

–

–

–

–

2.86

2.79

–

4.36

FeO

29.78 30.42 0.57

48.12 1.26

22.82 0.13

–

32.55 16.60 30.69 30.32 16.48 22.82 0.18

MnO

1.47

1.59

–

1.35

–

0.11

–

–

–

0.28

1.54

1.56

0.30

0.13

MgO

2.69

2.44

–

–

–

6.30

–

–

4.32

9.11

3.07

2.45

9.06

6.62

–

–

–

CaO

6.44

6.72

–

0.13

53.2

11.0

4.11

–

–

20.5

5.14

6.50

20.0

10.7

54.7

–

4.05

–

6.26

6.43

–

–

Na2O

–

–

1.54

–

–

1.91

8.54

0.88

–

0.81

–

–

0.89

1.90

–

–

8.97

0.72

–

0.08

1.01

–

K2O

–

–

15.9

–

–

1.67

0.22

16.3

9.82

–

–

–

–

1.64

–

–

0.42

16.0

–

–

15.8

9.60

P2O5

–

–

–

–

40.7

–

–

–

–

–

–

–

–

–

40.9

–

–

–

–

–

–

–

F

–

–

–

–

3.81

–

–

–

–

–

–

–

–

–

3.79

–

–

–

–

–

–

–

Total

98.7

98.7

99.8

99.8

100.0 97.3

98.5

98.8

97.8

100.9 98.6

98.8

100.7 98.2

99.6

99.5

98.8

98.3

99.3

99.6

99.5

97.3

(continued on next page)

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S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

Table 1 (continued) Component

Mu-10-124 Grt 1

Mu-10-24 Incl

Incl

Incl

Amp Ab

Kfs

Bt

Cpx

core

rim

Kfs

Ilm

Ap

Si

3.12

3.10

2.98

0.97

0.08

6.42

2.87

3.00

2.90

Ti

–

–

–

–

–

0.22

–

–

0.12

Al

1.85

1.84

0.99

–

–

1.96

1.12

0.99

Cr

–

–

–

1.03

–

–

–

–

Fe

2.00

2.06

0.02

0.03

0.09

2.99

0.00

Grt 1

Cpx

Amp Ap

Ilm

Ab

Kfs

core

rim

1.98

3.12

3.11

1.99

6.48

–

0.98

2.85

0.00

–

–

0.00

0.24

–

–

–

1.19

0.07

1.85

1.85

0.07

1.89

–

–

–

0.00

–

–

0.00

–

–

1.02

–

2.17

0.53

2.07

2.04

0.53

2.95

0.01

Grt 2

Incl

Bt

core

rim

Kfs

3.00

3.13

3.11

2.97

3.00

–

–

–

–

0.00

1.14

0.99

1.85

1.84

1.02

1.34

–

–

–

–

–

–

0.02

0.00

0.00

1.99

2.03

0.02

2.01

Mn

0.10

0.11

–

–

–

0.01

–

–

0.00

0.01

0.10

0.11

0.01

0.02

–

–

–

–

0.10

0.10

–

0.01

Mg

0.32

0.29

–

0.00

–

1.47

–

–

0.51

0.52

0.37

0.29

0.52

1.53

–

–

–

–

0.34

0.33

–

0.51

Ca

0.55

0.58

0.00

–

4.87

1.84

0.20

–

0.00

0.84

0.44

0.56

0.82

1.78

5.03

–

0.19

0.00

0.54

0.55

0.00

0.00

Na

–

–

0.14

–

–

0.58

0.74

0.08

0.00

0.06

–

–

0.07

0.57

–

–

0.78

0.07

–

0.01

0.09

0.00

K

–

–

0.95

–

–

0.33

0.01

0.98

1.00

0.00

–

–

0.00

0.32

–

–

0.02

0.96

–

–

0.94

0.96

P

–

–

–

–

2.94

–

–

–

–

–

–

–

–

–

2.98

–

–

–

–

–

–

–

F

–

–

–

–

1.03

–

–

–

–

–

–

–

–

–

1.03

–

–

–

–

–

–

–

Total

7.95

7.98

5.07

2.03

9.02

15.8

4.95

5.04

7.89

4.01

7.96

7.97

4.00

15.8

9.05

2.02

4.98

5.02

7.95

7.98

5.04

7.82

Compo- Mu-10-80 nent Grt 1 core

rim

Cpx 1

Amp

Ep

Cpx 2

Incl

Grt 2

Cpx 3

Ab 1

Ep

Kfs

core

rim

Grt 3

Incl

Incl

core

rim

Cpx

Cpx

SiO2

39.3

39.5

52.9

45.7

38.8

54.1

66.1

39.6

64.6

40.1

40.3

53.2

40.1

39.2

51.9

52.0

TiO2

0.12

–

0.45

0.36

–

0.16

–

0.11

–

0.14

–

0.21

0.10

–

0.51

0.61

Al2O3

19.7

19.7

3.15

11.8

25.9

1.59

19.6

29.5

17.0

19.4

20.3

1.86

20.6

20.5

4.48

4.76

Cr2O3

–

–

–

–

–

–

–

0.16

–

–

–

–

–

–

–

–

FeO

22.6

24.2

7.37

13.1

4.88

7.08

–

2.15

0.15

22.1

24.0

7.11

22.2

23.5

5.84

6.45

MnO

0.47

0.59

–

0.11

–

–

–

–

–

0.52

0.58

–

0.49

0.60

–

–

MgO

7.11

6.52

13.2

12.3

–

14.1

–

0.11

–

6.64

6.79

14.6

6.93

6.46

14.3

13.6

CaO

10.1

9.35

22.3

11.7

24.1

24.1

2.43

24.3

0.00

11.5

9.78

23.5

10.9

10.2

23.2

23.3

Na2O

–

–

0.85

2.09

–

0.49

11.54

–

0.35

–

–

0.65

–

–

0.93

0.90

K2O

–

–

–

0.40

–

–

–

–

16.4

–

–

–

–

–

–

–

P2O5

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

F

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

Total

99.4

99.8

100.2

97.6

93.7

101.5

99.7

96.0

98.5

100.4

101.7

101.1

101.4

100.5

101.1

101.6

Si

3.05

3.07

1.95

6.69

3.12

1.97

2.93

3.09

3.04

3.09

3.07

1.95

3.05

3.03

1.89

1.89

Ti

0.01

–

0.01

0.04

0.00

0.00

–

0.01

–

0.01

–

0.01

0.01

–

0.01

0.02

Al

1.81

1.80

0.14

2.04

2.46

0.07

1.02

2.71

0.94

1.76

1.82

0.08

1.84

1.87

0.19

0.20

Cr

–

–

0.00

–

–

0.00

–

–

–

–

–

0.00

–

–

0.00

0.00

Fe

1.47

1.57

0.23

1.61

0.33

0.22

0.00

0.14

0.01

1.43

1.52

0.22

1.41

1.52

0.18

0.20

Mn

0.03

0.04

0.00

0.01

–

0.00

–

–

–

0.03

0.04

0.00

0.03

0.04

0.00

0.00

Mg

0.82

0.76

0.73

2.70

–

0.76

–

0.01

–

0.76

0.77

0.80

0.78

0.74

0.78

0.74 0.91

Ca

0.84

0.78

0.88

1.84

2.08

0.94

0.12

2.03

0.00

0.95

0.80

0.92

0.89

0.84

0.91

Na

–

–

0.06

0.59

–

0.03

0.99

–

0.03

–

–

0.05

–

–

0.07

0.06

K

–

–

0.00

0.08

–

0.00

0.00

–

0.98

–

–

0.00

–

–

0.00

0.00

P

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

F

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

–

Total

8.03

8.03

4.00

15.6

7.99

4.00

5.06

7.98

5.00

8.02

8.02

4.03

8.02

8.04

4.03

4.02

Note. Grt, garnet; Cpx, clinopyroxene; Ap, apatite; Amp, amphibole; Bt, biotite; Ab, albite; Ilm, ilmenite; Kfs, K-feldspar; Ep, epidote; Sp, spinel; Incl, inclusion in mineral; Mu-10-80 and Mu-14-16 are garnet pyroxenites; Mu-10-24, Mu-10-82, and Mu-10-124 are granulites.

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463 Table 2. Major- and trace-element compositions of South Muya granulites Component

Mu-10-24 Mu-10-46 Mu-10-48 Mu-10-82 Mu-10-124

SiO2, wt.%

48.3

45.1

46.7

50.8

49.8

TiO2

0.56

2.67

2.60

1.85

2.45

Al2O3

7.45

13.8

13.4

16.3

13.2

FeO

11.5

20.6

21.1

14.8

17.4

MnO

0.18

0.27

0.33

0.20

0.32

MgO

10.9

3.71

3.05

2.47

2.73

CaO

18.2

8.01

7.54

7.00

6.73

Na2O

1.24

2.90

2.74

3.28

3.27

K2O

0.39

1.04

0.76

2.44

2.38

P2O5

0.03

1.62

1.31

0.63

1.10

LOI

1.46

0.89

1.03

0.17

0.45

Total

100.2

100.8

100.6

100.2

100.1

Cs, ppm

0.16

0.13

0.60

0.04

0.81

Rb

6.99

22.0

18.9

30.9

39.3

Ba

146

400

315

1 449

1 755

Th

0.34

2.28

1.06

2.06

1.65

U

0.10

0.93

0.37

0.46

0.58

Nb

1.05

13.4

23.1

24.2

31.1

Ta

0.05

0.93

1.52

1.26

1.53

La

4.24

42.6

50.8

52.1

49.7

Ce

10.9

94.9

124

124

117

Pb

6.85

4.00

3.59

6.94

8.66

Pr

1.53

11.4

16.0

15.3

15.0

Sr

330

350

310

727

563

Nd

8.18

53.2

70.8

68.4

73.3

Zr

37.5

146

54

389

825

Hf

1.46

3.62

1.76

8.88

16.8

Sm

2.31

11.4

15.4

14.6

16.6

Eu

0.60

2.90

2.61

2.97

5.16

Gd

2.93

12.0

16.0

15.1

18.4

Tb

0.49

1.60

2.27

2.16

2.66

Dy

3.33

9.33

13.8

13.1

16.2

Y

15.5

44.9

64.1

60.9

77.6

Ho

0.66

1.74

2.63

2.42

3.14

Er

2.07

4.98

7.49

6.84

9.12

Tm

0.28

0.61

0.99

0.89

1.22

Yb

1.72

4.01

6.68

5.65

8.28

Lu

0.26

0.58

0.94

0.77

1.22

V

463

77.9

52.6

61.0

47.3

Cr

91.0

0.62

3.04

2.65

3.33

Co

56.0

51.9

44.1

28.1

24.1

Ni

63.6

7.69

1.84

8.38

9.40

Cu

67.4

28.7

40.3

20.7

16.5

Zn

96.3

148

166

190

263

Ga

11.4

23.4

21.4

27.6

23.7

457

Major- and trace-element chemistry of granulites. The analyzed granulite samples have quite variable compositions (Table 2). REE show large ranges (Fig. 5a), between 251 and 337 ppm for most of samples, while REE in Mu-10-24 is much lower (39 ppm). REE-rich granulites have markedly higher contents of Al2O3, FeO, TiO2, MnO, alkalis (K2O, Na2O) and P2O5, as well as low Mg# (0.26–32), while the sample with low REE has high Ca#. REE spectra show negative slopes with (La/Yb)N from 1.67 to 7.21 and a negative Eu anomaly (Eu/Eu* = 0.90–0.51). The low-REE granulite has multielement spectra with Pb and Sr peaks and a P minimum (Fig. 5b), while the respective anomalies in the REE-rich Mu-10-82 granulite are negative or absent. All granulite samples are depleted in Th (Th and U in some cases) and in Ta-Nb. The REE spectra of Mu-10-48 and Mu-10-46 also have a deep Zr-Hf minimum. Iron group elements are largely variable in all samples. U–Pb dating of zircons and Hf isotope composition. Zircon grains from the Mu-10-124 granulite selected for dating are mostly colorless subisometric round or less often prismatic, with smooth contours (Fig. 6). They do not show distinct zoning or difference between core and rim in cathode luminescence (CL) images. Their Th/U ratios (0.4–0.9) are typical of igneous rocks, this being a feature of granulite zircons, along with a complex interior structure (Kaulina, 2010). However, these features do not have unambiguous implications for the genesis of zircons. The U–Pb ages of zircons (Table 3) are from 586 to 713 Ma, or Neoproterozoic (Fig. 6). Thirty one zircon grains show variable U–Pb ages with different degrees of discordance. The most concordant 206Pb/238U ages (discordance within 5%, 15 grains) vary between 606 and 713 Ma, with a distinct peak at 630 Ma and Th/U = 0.47–0.83. Less concordant 206Pb/238U ages (discordance above 5%, 17 zircons) are in the range 556–711 Ma at Th/U = 0.41–0.89. The data points of these zircons define a discordia with an upper intercept age of 715 63 Ma. The analyzed zircons have 176Hf/177Hf ratios from 0.282405 to 0.282561 (1σ = 0.000450–0.007600). Primary 176 Hf/177Hf(T) ratios estimated using the decay constant 176Lu 1.865 10–11 yr–1 (Bouvier et al., 2008) are 0.282392– 0.282542 (Table 4). The εHf(T) values calculated with the obtained U–Pb ages and chondritic 176Hf/177Hf (0.282785) and 176Lu/177Hf (0.0336) (Scherer et al., 2001) are positive, within 0.4–6.0 (Fig. 7, Table 4). The model age TDM estimated for a single-stage model, which corresponds to the time when the source magma separated from the mantle, is 1.0–1.2 Ga C (Table 4). A crust model age TDM of 1.2–1.5 Ga (Table 4) follows from a two-stage model for a depleted mantle source with the modern ratios 176Hf/177Hf = 0.283251 (Nowell et al., 1998) and 176Lu/177Hf = 0.0384 (Griffin et al., 2000a), assuming that the source magma separated from the depleted mantle reservoir and had an average crustal 176Lu/177Hf ratio of 0.015.

458

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

Fig. 5. Trace element patterns in granulites. a, REE spectra; b, multielement spider diagram.

Discussion The major-element contents of all granulites (except for low-REE Mu-10-24) correspond to those in continental rift basalts in the discrimination diagram of Velikoslavinsky and Glebovitsky (2005), and their compositions fall next to the field of flood basalts (Fig. 8). On the other hand, discrimination based on trace elements (especially, Th, Yb, and HFSE) fails to constrain unambiguously the protolith nature. Most of the granulite samples have their compositions (especially, Fe, Ti, P, Ba, Th, Sr, Nb, and Ta enrichments) typical of

differentiated within-plate basaltic melts. In the discrimination diagrams of Wood (1980) (Fig. 9a) and Pearce and Norry (1979) (Fig. 9b), REE-rich granulites fall in the fields of enriched within-plate (Mu-10-124 and Mu-10-82 with highest Zr), island-arc or E-MORB-type basalts with Ce/Pb (14–34) close to that in oceanic basalts (25), while the REE-depleted Mu-10-24 sample approaches mid-ocean ridge and island arc tholeiites. Y/Nb ratios in the granulites (2.5–3.3) are within the range for continental and oceanic tholeiite basalts (Pearce and Cann, 1973). The contents of Zr, Y and Ti correspond to a protolith composition most proximal to island arc basalts

Fig. 6. Concordia diagram for analyzed zircons from the South Muya granulites, according to LA-ICP-MS UPb dating and CL microphotographs of some grains.

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

459

Table 3. U-Th-Pb isotope ratios and ages of zircons from South Muya granulites Sam- Th/U Isotope ratios ple 207 1σ Pb No. 206 Pb

Age, Ma 207Pb 235

U

1σ

206

Pb

238

U

1σ

208

Pb

232

Th

1σ

Pb 1σ

Pb 1σ

Pb 1σ

Pb 1σ

207

207

206

208

206

235

238

232

Pb

U

U

D, %

Th

14-1

0.84 0.06493 0.00033 0.88355 0.01008 0.09871 0.00094 0.03135 0.00069 772

11

643

5

607

6

624

14

21

14-2

0.76 0.06104 0.00030 0.90368 0.01014 0.10740 0.00102 0.03212 0.00072 641

11

654

5

658

6

639

14

–2.7

14-3

0.89 0.06345 0.00030 0.94668 0.01009 0.10823 0.00101 0.03297 0.00075 723

11

676

5

662

6

656

15

8.4

14-4

0.59 0.06161 0.00035 0.88516 0.01108 0.10421 0.00100 0.03234 0.00079 661

13

644

6

639

6

643

15

3.3

14-5

0.53 0.06202 0.00046 0.81405 0.01220 0.09521 0.00096 0.03011 0.00081 675

17

605

7

586

6

600

16

13

14-6

0.65 0.06054 0.00034 0.83047 0.01030 0.09951 0.00095 0.03013 0.00078 623

13

614

6

612

6

600

15

1.8

14-7

0.60 0.06480 0.00039 0.95261 0.01238 0.10663 0.00103 0.03123 0.00085 768

14

679

6

653

6

622

17

15

14-8

0.80 0.06243 0.00030 0.87780 0.00968 0.10200 0.00094 0.03111 0.00085 689

11

640

5

626

5

619

17

9.1

14-9

0.75 0.06097 0.00034 0.80018 0.00973 0.09520 0.00090 0.02906 0.00083 638

13

597

5

586

5

579

16

8.2

14-10 0.48 0.06029 0.00037 0.74846 0.00978 0.09005 0.00086 0.02824 0.00085 614

14

567

6

556

5

563

17

9.4

14-12 0.64 0.06067 0.00037 0.90549 0.01195 0.10825 0.00106 0.03455 0.00098 628

14

655

6

663

6

687

19

–5.6

14-13 0.75 0.06053 0.00030 0.89147 0.01010 0.10682 0.00102 0.03338 0.00095 623

12

647

5

654

6

664

19

–5.0

14-14 0.83 0.06076 0.00034 0.83684 0.01047 0.09989 0.00097 0.03290 0.00098 631

13

617

6

614

6

654

19

2.7

14-15 0.66 0.06046 0.00038 0.86185 0.01151 0.10338 0.00102 0.03387 0.00105 620

15

631

6

634

6

673

21

–2.3

14-16 0.47 0.06166 0.00037 0.92576 0.01211 0.10888 0.00107 0.03491 0.00114 662

14

665

6

666

6

694

22

–0.6

14-18 0.74 0.06088 0.00035 0.88182 0.01111 0.10504 0.00102 0.03319 0.00114 635

13

642

6

644

6

660

22

–1.4

14-19 0.79 0.06283 0.00036 0.89090 0.01125 0.10284 0.00100 0.03241 0.00115 702

13

647

6

631

6

645

23

10

14-20 0.85 0.06248 0.00034 0.82808 0.01013 0.09611 0.00093 0.03061 0.00112 691

13

613

6

592

5

609

22

14

14-26 0.69 0.06286 0.00034 0.88845 0.01087 0.10251 0.00101 0.03101 0.00103 704

13

646

6

629

6

617

20

10

14-28 0.79 0.06308 0.00033 0.99383 0.01174 0.11427 0.00112 0.03449 0.00119 711

12

701

6

697

6

685

23

2.0

14-29 0.67 0.05947 0.00034 0.80751 0.01015 0.09849 0.00098 0.03014 0.00110 584

13

601

6

606

6

600

22

–3.8

14-31 0.68 0.06366 0.00035 0.91722 0.01142 0.10451 0.00104 0.03280 0.00125 730

13

661

6

641

6

652

24

12

14-32 0.78 0.06193 0.00032 0.89974 0.01058 0.10538 0.00104 0.03230 0.00128 672

12

652

6

646

6

643

25

3.9

14-36 0.71 0.06333 0.00034 0.87120 0.01057 0.09978 0.00099 0.03177 0.00132 719

12

636

6

613

6

632

26

14

14-41 0.41 0.06245 0.00039 0.87540 0.01193 0.10167 0.00104 0.03467 0.00153 690

15

638

6

624

6

689

30

9.6

14-44 0.78 0.06358 0.00036 0.88226 0.01122 0.10065 0.00102 0.03139 0.00143 728

13

642

6

618

6

625

28

15

14-45 0.90 0.06167 0.00034 0.99096 0.01232 0.11656 0.00117 0.03478 0.00165 663

13

699

6

711

7

691

32

–7.2

14-47 0.79 0.06237 0.00038 0.95757 0.01270 0.11137 0.00114 0.03479 0.00173 687

14

682

7

681

7

691

34

0.9

14-51 0.83 0.06414 0.00032 1.03394 0.01183 0.11690 0.00117 0.03552 0.00110 746

11

721

6

713

7

705

21

4.4

14-52 0.80 0.06296 0.00032 0.97178 0.01120 0.11195 0.00113 0.03355 0.00109 707

12

689

6

684

7

667

21

3.3

14-53 0.72 0.06108 0.00034 0.86747 0.01086 0.10301 0.00106 0.03097 0.00107 642

13

634

6

632

6

616

21

1.6

Note. D, Age discordance.

(Pearce and Cann, 1973), while all samples show typically crustal anomalies in Ti, Ta and Nb. Their La/Nb ratios (1.6–3.2) are close to that in suprasubduction volcanics (La/Nb ~ 3) but exceed those of within-plate varieties (La/Nb < 1); the lowest values are measured in the samples Mu-10-82 and Mu-10-124 with the highest enrichment in K, Ba, Rb, Zr, and Nb. It is difficult to judge about the origin of volcanics using discrimination models based on the least mobile elements (Zr, Ti, V, Y, Th, Hf, Nb, Ta, Sm, Sc), as Li et al. (2015) infer proceeding from synthesis of available trace-element data for basaltic rocks and from the estimated accuracy of such discrimination diagrams. According to Li et al. (2015), the

conditions of magma generation are indicated most reliably by mantle-normalized REE spectra. Basaltic rocks enriched in LILE and LREE but depleted in HFSE are interpreted as originated in island-arc (Winter, 2001) or within-plate (Li et al., 2003) settings. In the latter case, LILE and LREE enrichment and HFSE depletion record crustal contamination of magma or a magma source enriched by fluids released from a subducting slab. The granulite samples Mu-10-46 and Mu-10-48, with the lowest contents of SiO2 and alkalis and the highest TiO2, FeO, MgO and CaO, show depletion (even to zero) in Ti, Ta and Nb along with distinct Zr-Hf and Pb minimums. Compared to these, the Mu-10-82 and Mu-10-124 granulites enriched in SiO2, K2O and Na2O have notably

460

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

Table 4. Hf isotope compositions, εHf and Hf model ages of zircons from South Muya granulites Sample No.

176

Lu/177Hf

1σ

176

Hf/177Hf

1σ

Age, Ma

176

εHf(T)

TDM, Ma

TDM, Ma (176Lu/177Hf = 0.015)

14-01

0.001122

0.000059

0.282486

0.001800

607

0.282473

2.5

1089

1373

14-02

0.001141

0.000073

0.282492

0.002800

658

0.282478

3.8

1081

1330

14-03

0.001796

0.000160

0.282485

0.007600

662

0.282463

3.4

1110

1362

14-04

0.001292

0.000064

0.282470

0.001800

639

0.282455

2.6

1117

1395

14-05

0.000678

0.000030

0.282442

0.001400

586

0.282435

0.7

1138

1473

14-06

0.000937

0.000012

0.282475

0.000450

612

0.282464

2.3

1099

1390

14-07

0.000739

0.000071

0.282490

0.003300

653

0.282481

3.8

1073

1327

14-08

0.000997

0.000051

0.282477

0.002400

626

0.282465

2.7

1098

1379

14-09

0.000715

0.000012

0.282435

0.000480

586

0.282427

0.4

1149

1489

14-12

0.001102

0.000057

0.282488

0.002700

663

0.282474

3.8

1086

1335

14-13

0.000984

0.000076

0.282464

0.003500

654

0.282452

2.8

1116

1391

14-14

0.000670

0.000036

0.282454

0.001600

614

0.282446

1.7

1121

1429

14-15

0.000866

0.000020

0.282467

0.000860

634

0.282457

2.5

1108

1393

14-16

0.000951

0.000058

0.282474

0.002400

666

0.282462

3.4

1101

1361

14-18

0.000653

0.000052

0.282474

0.002500

644

0.282466

3.1

1093

1366

14-19

0.000918

0.000050

0.282461

0.002300

631

0.282450

2.2

1118

1410

14-20

0.001078

0.000023

0.282484

0.000810

592

0.282472

2.1

1091

1385

14-26

0.001313

0.000087

0.282507

0.002600

629

0.282492

3.6

1065

1318

14-27

0.001347

0.000085

0.282487

0.002400

629

0.282471

2.9

1094

1364

14-28

0.001007

0.000071

0.282405

0.003100

697

0.282392

1.6

1200

1498

14-29

0.000852

0.000032

0.282494

0.001500

606

0.282484

2.9

1070

1349

14-31

0.001161

0.000043

0.282487

0.002400

641

0.282473

3.3

1089

1352

14-32

0.001114

0.000084

0.282519

0.003700

646

0.282506

4.5

1042

1276

14-36

0.001639

0.000019

0.282483

0.001300

613

0.282464

2.3

1109

1390

14-44

0.001009

0.000030

0.282465

0.001200

618

0.282453

2.0

1115

1411

14-45

0.001095

0.000120

0.282457

0.005900

711

0.282442

3.8

1129

1376

14-47

0.001165

0.000064

0.282491

0.003400

681

0.282476

4.3

1083

1320

14-51

0.001760

0.000056

0.282528

0.002700

713

0.282504

6.0

1048

1236

14-52

0.001059

0.000052

0.282520

0.002800

684

0.282506

5.4

1040

1250

14-53

0.001638

0.000029

0.282561

0.001500

632

0.282542

5.5

997

1204

higher Ba (at a similar Rb concentration), a marked Nb-Ta minimum but no depletion in Pb and Zr-Hf. Therefore, most of the analyzed granulites are the most proximal to withinplate basalts. The composition variations may result from differentiation of melts within the crust and crustal contamination in the case of highest Ba and Zr enrichments. The idea of a differentiated protolith is supported by correlation of some major elements with the total of incompatible elements and with enrichment in the most strongly incompatible elements (including LREE and Zr), as well as by low Sr/Nd ratios (4.4–10.6). On the other hand, zircons from the most highly enriched sample Mu-10-124 show quite a narrow range of model ages and lack xenogenic grains. The massive granulite Mu-10-24, with CaO enrichment and gently sloping chondrite-normalized REE spectra, is rather compositionally proximal to pyroxenites and has very high Y/Nb and Zr/Nb ratios (14.7 and 35.5, respectively) typical

Hf/177Hf

C

of seafloor oceanic basalts. Therefore, its source composition may have been different from that of the other samples: a pyroxene-rich lower crustal cumulate (pyroxenite), indicated by major-element and trace-element patterns (especially, Sr/Nd as high as 40). The large range of trace elements in the granulites may likewise be due to their concentrating by some accessories, e.g., phosphates, which is evident in distinct correlation between REE and P2O5. Positive correlation of Ti with Ta and (to a lesser degree) with Nb shows rutile to be their concentrator. The granulites enriched in REE have also greater enrichments in K2O (a K peak), Na2O, P2O5 and most trace elements and lower Mg# relative to the North Muya eclogites (Shatsky et al., 2012 and unpublished evidence by Skuzovatov et al.). Furthermore, the granulites have higher (La/Yb)N ratios than eclogites and different trace-element patterns (Th-U-, Nb-Ta-

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

461

Fig. 7. Hf isotope composition of analyzed zircons from the South Muya granulites, according to in situ MC-LA-ICP-MS Lu-Hf data. Lines of continental crust evolution correspond to a crustal average 176Lu/177Hf ratio of 0.015 (Griffin et al., 2002).

and Pb minimum, K maximum, and much lower Cr and Ni contents). On the other hand, the trace-element patterns of depleted Mu-10-24 are similar to those in some eclogites. Therefore, similar material may have been present in the granulite protolith. The analyzed granulitic zircons lack oscillatory zoning but hold high Th/U ratios common to magmatic zircons. The distribution of data points along the discordia shows lead losses as a result of recrystallization during Neoproterozoic (Late Baikalian) subduction and collision marked by a distinct age peak at 630 Ma (Fig. 6). The same age was inferred for the North Muya eclogites (Shatsky et al., 2012), while the age of the oldest zircons (720 Ma) may mark the origin of the protolith (i.e. be the age of magmatism). Positive εHf(T) values, along with a Mesoproterozoic Hf model age of zircons (1.2–1.5 Ga), indicate participation of a Baikalian juvenile source in the granulite formation and melting of older (Mesoor Paleoproterozoic) lower crust. Different shares of the juvenile and old-crust components may account for the majorand trace-element variations. Note that the hypothetical age of within-plate magmatism (720 Ma), which produced the protolith of the granulites, marks the end of the post-collisional phase in the Early Baikalian orogenic cycle in the Baikal–Muya belt and the onset of the Late Baikalian orogeny. This is the age inferred for many mafic complexes in the southern periphery of the Siberian craton that result from plume activity and related breakup of Rodinia (Izokh et al., 1998; Mekhonoshin et al., 1993; Polyakov et al., 2013). The Zhanok volcanoplutonic complex, as well as layered mafic intrusions, originated about the same time within the Muya subzone (Rytsk et al., 2011), presumably, during accretion over a hotspot. This age is also close to the tentative age of the youngest uncontaminated protolith of the North Muya eclogites (Shatskii et al., 2014). However, the date cannot be unambiguously interpreted as the age of magmatism because zircons lack oscillatory zoning.

On the other hand, the spatial relations of the South and North Muya blocks in Early Baikalian time and the grades of later Late Baikalian metamorphism remain open to discussion. The peak of metamorphism at 630 Ma is close to the age of the North Muya eclogites (Shatsky et al., 2012). Most of dates we obtained approach the age ranges of metamorphic zircons from the North Muya eclogite-bearing schists and gneisses (Shatsky et al., 2015). The age proximity of metamorphism in the two blocks indicates that they both were involved into Late Baikalian subduction-collision processes, though at obviously different PT parameters. The maximum PT conditions of metamorphism for the South Muya eclogites were estimated at 900–950 °C and 11–19 kbar (Doronina and Sklyarov, 1995) but the high-pressure rocks of the block remained poorly studied so far, including in terms of geochronology. The calculated equilibrium parameters for the primary parageneses of the South Muya mafic rocks fall within the PT trend for the Kindikan Group rocks in the South Muya block, but are notably below

Fig. 8. Granulite compoisitons in discrimination diagram for island-arc and continental basalts (Velikoslavinsky and Glebovitsky, 2005).

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S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463

Fig. 9. Granulite compoisitons in discrimination diagrams of Wood (1980) (a) and Pearce and Norry (1979) (b). IAB, island-arc basalts; OIB, ocean island basalts; N-MORB, primitive mid-ocean ridge basalts; E-MORB, enriched mid-ocean ridge basalts; WPB, within-plate basalts.

their peak values (800 °C and 15 kbar) (Doronina and Sklyarov, 1995). The compositions of granulite minerals, reverse zoning in garnet, and the PT conditions in which the analyzed granulites and garnet pyroxenites formed, as well as the absence of decompression textures commonly shown by decompressed high-pressure and ultrahigh-pressure minerals, indicate that the granulites were exhumed from depths no more than 40 km and did not undergo a high-pressure stage in their history. On the other hand, the PT parameters of the last equilibrium in the mineral assemblage depend on kinetics of cooling and exhumation. According to simulations (Gerya et al., 2008), granulites may have been produced by exhumation of partially molten subducted crust in a hot subduction channel early during the collision as the temperature reached 900 °C and hotter. Inasmuch as the U–Pb isotopic system closed at 900 °C (Cherniak and Watson, 2001), the granulite protolith apparently experienced a peak of metamorphism at a hotter temperature. This model implies exhumation in the subduction channel, diffuse reequilibration of parageneses in the HT-HP region of granulite stability, and further cooling upon ascent to middle and upper crust depths. Recrystallization at high temperatures could account for the inherited high Th/U ratios as the U–Pb system was reset with the loss of typical zoning. Few temperature values obtained for garnet pyroxenite and ultramafics subjected to retrograde metamorphism (up to 780 °C and 830 °C, respectively) may correspond to greater depths of the protolith. In this case, the poikilitic texture of clinopyroxene in garnet pyroxenite may have resulted from breakdown of omphacite during eclogite-facies metamorphism of basaltic rocks. Genesis of the ultramafics and their relation with garnet pyroxenite likewise require further investigation.

Conclusions The granulites we studied were apparently derived from within-plate basaltic rocks, with their Late Baikalian age

(720 Ma) corresponding to the end of the Early Baikalian orogenic cycle in the Baikal–Muya belt. The composition of the precursor basalts resulted from mixing of a Late Baikalian juvenile mantle source with material of older (Meso- or Paleoproterozoic) continental crust. In Late Baikalian time, the South and North Muya blocks were jointly involved into subduction and collision events that peaked at 630 Ma. The subduction-collisional metamorphism with a peak at 630 Ma produced granulites after basaltic rocks, which bear no high-pressure signatures unlike the North Muya eclogites. Thus, it remains unclear how the two blocks became juxtaposed during the Baikalian orogeny and which were the PT trends in their metamorphic history. We greatly appreciate the aid of F.-L. Lin and Y.-S. Chien in the dating work. The study was supported by grant 13-05-00261 from the Russian Foundation for Basic research and was carried out as part of Integration Project 49, 2012-2014 of the Siberian Branch of the Russian Academy of Sciences.

References Avchenko, O.V., Gabov, N.F., Kozyrev, I.V., Konikov, A.Z., Travin, L.V., 1989. Eclogites of the North Muya block: composition and genesis. Izv. AN SSSR, Ser. Geol., No. 5, 68–82. Berman, R.G., 2007. winTWQ (version 2.3): a software package for performing internally-consistent thermobarometric calculations. Geological Survey of Canada, Open File 5462 (ed. 2.32). Bouvier, A., Vervoort, J.D., Patchett, P.J., 2008. The Lu–Hf and Sm–Nd isotopic composition of CHUR: constraints from unequilibrated chondrites and implications for the bulk composition of terrestrial planets. Earth Planet. Sci. Lett. 273, 48–57. Bozhko, N., Parfenova, O., Kirmasov, A., Clochko, A., 1999. Metamorphic history and paleotectonics of Precambrian complexes in the eastern Baikal–Muya belt. Bull. Moscow University, No. 2, 9–18. Bulgatov, A.N., 1988. Geological-geophysical model of the upper crust of Northern Transbaikal. Geologiya i Geofizika (Soviet Geology and Geophysics) 29 (9), 62–68 (54–58).

S.Yu. Skuzovatov et al. / Russian Geology and Geophysics 57 (2016) 451–463 Bulgatov, A.N., Gordienko, I.V., 1999. Terranes of the Baikal mountain province and their gold mineralization. Geologiya Rudnykh Mestorozhdenii 41 (3), 230–240. Cherniak, D.J., Watson, E.B., 2001. Pb diffusion in zircon. Chem. Geol. 172, 5–24. Doronina, N.A., Sklyarov, E.V., 1995. Eclogitic and granulitic metamorphism within the South Muya block. Doklady AN SSR 340 (6), 793–796. Ellis, D.G., Green, D.H., 1979. An experimental study of the effect of Ca upon garnet-clinopyroxene Fe-Mg exchange equilibria. Contrib. Mineral. Petrol. 71, 13–22. Gerya, T., Perchuk, L.L., Burg, J.-P., 2008. Transient hot channels: Perpetrating and regurgitating ultrahigh-pressure, high-temperature crust–mantle associations in collision belts. Lithos 103, 236–256. Griffin, W.L., Pearson, N.J., Belousova, E., Jackson, S.E., van Achterberg, E., O’Reilly, S.Y., Shee, S.R., 2000a. The Hf isotope composition of cratonic mantle: LAM-MC-ICPMS analysis of zircon megacrysts in kimberlites. Geochim. Cosmochim. Acta 64, 133–147. Griffin, W.L., Wang, X., Jackson, S.E., Pearson, N.J., O’Reilly, S.Y., Xu, X., Zhou, X., 2000b. Zircon chemistry and magma mixing, SE China: In-situ analysis of Hf isotopes, Tonglu and Pingtan igneous complexes. Lithos 61, 237–269. Grudinin, M.I., Menshagin, Yu.V., 1988. A find of garnet ultramafics and eclogites in the Archean South Muya block (Northern Baikal area). Dokl. Akad. Nauk SSR 299 (2), 434–437. Grudinin, M.I., Mitrofanov, V.G., 1998. Position and composition of the Kedrovsky anorthosite-gabbro intrusion, in: Sizykh, A.I. (Ed.), Precambrian Complexes in East Siberia [in Russian]. Irkutsk, pp. 68–77. Izokh, A.E., Gibsher, A., Zhuravlev, D.Z., Balykin, P.A., 1998. Sm-Nd data on the ages of mafic-ultramafic rocks in the eastern arm of the Baikal–Muya ophiolite belt. Dokl. Akad. Nauk SSR 360 (1), 88–92. Kaulina, V., 2010. Formation and Transformation of Zircons in Polymetamorphic Complexes [in Russian]. KNC RAN, Apatity. Kohn, M.J., Spear, F.S., 1990. Two new geobarometers for garnet amphibolites with applications to southeastern Vermont. Amer. Miner. 75, 89–96. Lan, C.-Y., Tadashi, U., Wang, K.-L., Yui, T.-F., Okamoto, K., Lee, C.-S., Tsutsumi, Y., Yokoyama, K., Hirata, T., Kon, Y., Orihashi, Y., Horie, K., Hidaka, H., Liou, J.G., 2009. Detrital zircon evidence for the antiquity of Taiwan. Geosci. J. 13, 233–243. Li, Z.X., Li, X.H., Kinny, P.D., Wang, J., Zhang, S., Zhou, H., 2003. Geochronology of Neoproterozoic syn-rift magmatism in the Yangtze Craton, South China and correlations with other continents: evidence for a mantle superplume that broke up Rodinia. Precambrian Res. 122, 85–109. Li, C., Arndt, N.T., Tang, Q., Ripley, E.M., 2015. Trace element discrimination diagrams. Lithos 232, 76–83. Mekhonoshin, A., Bognibov, V.I., Lomonosova, E.I., 1993. Rare-earth elements and petrogenesis of ultrabasie-basite massifs in Southern Siberia. Geologiya i Geofizika (Russian Geology and Geophysics) 34 (2), 43–49 (39–46). Nowell, G.M., Kempton, P.D., Noble, S.R., Fitton, J.G., Saunders, A.D., Mahoney, J.J., Taylor, R.N., 1998. High precision Hf isotope measure-

463

ments of MORB and OIB by thermal ionisation mass spectrometry: insights into the depleted mantle. Chem. Geol. 149, 211–233. Pearce, J.A., Cann, J.R., 1973, Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth Planet. Sci. Lett. 19, 290–300. Pearce, J.A., Norry, M.J., 1979. Petrogenetic implications of Ti, Zr, Y and Nb variations in volcanic rocks. Contrib. Mineral. Petrol. 69, 33–47. Polyakov, G.V., Tolstykh, N.D., Mekhonoshin, A., Izokh, A.E., Podlipsky, M.Yu., Orsoev, D.A., Kolotilina, B., 2013. Ultramafic–mafic igneous complexes of the Precambrian East Siberian metallogenic province (southern framing of the Siberian Craton): age, composition, origin, and ore potential Russian Geology and Geophysics (Geologiya i Geofizika) 54 (11), 1319–1331 (1689–1704). Ravna, E.G.K., 2000. The garnet-clinopyroxene geothermometer: an update calibration. J. Metamorph. Geol. 18, 211–219. Rytsk, E.Yu., Kovach, V.P., Kovalenko, V.I., Yarmolyuk, V.V., 2007. Structure and evolution of continental crust in the Baikal fold area. Geotektonika, No. 6, 23–51. Rytsk, E.Yu., Kovach, V.P., Yarmolyuk, V.V., Kovalenko, V.I., Bogomolov, E., Kotov, A.B., 2011. Isotopy and evolution of continental crust in the East Transbaikalian segment of the Central-Asian orogen. Geotektonika, No. 5, 17–51. Scherer, E., Munker, C., Mezger, K., 2001. Calibration of the lutetium-hafnium clock. Science 293, 683–687. Shatskii, V.S., Skuzovatov, S.Yu., Ragozin, A.L., Dril, S.I., 2014. Evidence of Neoproterozoic continental subduction in the Baikal–Muya fold belt. Dokl. Earth Sci. 459 (1), 1442–1445. Shatsky, V.S., Sitnikova, E.S., Tomilenko, A.A., Ragozin, A.L., Kozmenko, O.A., Iagoutz, E., 2012. Eclogite-gneiss complex of the Muya block (East Siberia): age, mineralogy, geochemistry, and petrology. Russian Geology and Geophysics (Geologiya i Geofizika) 53 (6), 501–521 (657–682). Shatsky, V.S., Malkovets, V.G., Belousova, E.A., Skuzovatov, S.Yu., 2015. Evolution history of the Neoproterozoic eclogite-bearing complex of the Muya dome (Central Asian Orogenic Belt): constraints from zircon U–Pb age, Hf and whole-rock Nd isotopes. Precambrian Res. 261, 1–11. Tsygankov, A.A., 2005. The Late Precambrian Magmatic History of the Baikal–Muya Volcanoplutonic Belt [in Russian]. Izd. SO RAN, Novosibirsk. Velikoslavinsky, S.D., Glebovitsky, V.A., 2005. A new discriminant diagram for classification of island-arc and continental basalts on the basis of petrochemical data. Dokl. Earth Sci. 401 (2), 308–310. Whitney, D.L., Evans, B.W., 2010. Abbreviations for names of rock-forming minerals. Am. Mineral. 95, 185–187. Winter, J.D., 2001. An Introduction to Igneous and Metamorphic Petology. Prentice Hall, New Jersey. Wood, D.A., 1980. The application of a Th-Hf-Ta diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary volcanic province. Earth Planet. Sci. Lett. 50, 11–30.

Editorial responsibility: V.V. Reverdatto