Geochemistry of Precambrian volcanosedimentary rocks of the Karsakpai Group in southern Ulutau (Central Kazakhstan)

Available online at www.sciencedirect.com

ScienceDirect Russian Geology and Geophysics 58 (2017) 935–948 www.elsevier.com/locate/rgg

Geochemistry of Precambrian volcanosedimentary rocks of the Karsakpai Group in southern Ulutau (Central Kazakhstan) N.V. Dmitrieva a,b,*, E.F. Letnikova a,b, I.A. Vishnevskaya a,b, P.A. Serov c a

V.S. Sobolev Institute of Geology and Mineralogy, Siberian Branch of the Russian Academy of Sciences, pr. Akademika Koptyuga 3, Novosibirsk, 630090, Russia b Novosibirsk State University, ul. Pirogova 2, Novosibirsk, 630090, Russia c Geological Institute, Kola Scientific Center of the Russian Academy of Sciences, ul. Fersmana 14, Apatity, Murmansk Region, 184209, Russia Received 5 October 2016; received in revised form 12 December 2016; accepted 22 December 2016

Abstract We have analyzed the isotope-geochemical features of volcanosedimentary rocks of the Karsakpai Group in southern Ulutau (Central Kazakhstan): mafic volcanics, siliceous and siliceous-ferruginous sediments, and quartz–sericite–chlorite schists. The close association of ferruginous quartzites with intraplate volcanics indicates that they formed in a tectonically active basin. The Nd isotope composition of ferruginous quartzites was governed by synchronous underwater volcanism, whereas the 143Nd/144Nd value of schists was additionally controlled by the Nd isotope composition of older sources. The Mesoproterozoic Nd model ages and positive εNd(t) values of the metaterrigenous rocks of the Karsakpai Group indicate the presence of Mesoproterozoic juvenile material in the provenance. The minimum Nd model ages suggest the lower boundary of sedimentation of 1.3 Ga. © 2017, V.S. Sobolev IGM, Siberian Branch of the RAS. Published by Elsevier B.V. All rights reserved. Keywords: Late Precambrian; volcanosedimentary rocks; sedimentary rocks of iron formation; rift volcanics; Sm–Nd isotope composition; Central Kazakhstan

Introduction The Ulutau block is one of the largest Precambrian sialic massifs in Central Kazakhstan. It includes several structural zones of N–S strike separated by large longitudinal ruptures (Fig. 1). A volcanosedimentary complex of the Karsakpai Group is of special interest among the metamorphic rocks of the Ulutau terrane. It is characterized by a predominance of basaltic volcanics, on the one hand, and the presence of persistent horizons of ferruginous quartzites, on the other. The drastic difference of this complex from others in the Precambrian section was reported by all researchers studying the Ulutau region (Filatova, 1962; Polovinkina, 1952; Satpaev, 1928; Zaitsev and Rozanov, 1971). The ferruginous-siliceous sediments are spatially and genetically related to mafic lavas. Strakhov (1947) related the formation of iron ores to the activity of fumaroles and hydrotherms. Polovinkina (1952) established a genetic relationship between iron formations and submarine eruptions of mafic lavas. Two large epochs of formation of ferruginous quartzites are recognized in the Pre-

* Corresponding author. E-mail address: [email protected] (N.V. Dmitrieva)

cambrian: Archean–Paleoproterozoic (3.20–1.89 Ga) (Klein and Beukes, 1993) and Neoproterozoic (~1000–635 Ma) (James et al., 1983; Klein, 2005). The previous researchers assigned the Karsakpai Group to Early Proterozoic deposits (Filatova and Rozanov, 1995; Rozanov, 1976). Comprehensive isotope-geochemical studies of the Karsakpai Group rocks have not been performed. In this work we consider the geochemical features of rocks and their isotope parameters in order to reconstruct the geodynamic settings of formation of the protoliths of metavolcanosedimentary rocks and to refine their age.

Geologic location and petrography of the objects of study The Karsakpai Group is localized within the Karsakpai structural zone composed mostly of effusive mafic rocks and subordinate felsic rocks and is overlain by the metamorphosed volcanoterrigenous Bozdak Group (Filatova, 1983) (Fig. 1). The latter is dated (by detrital zircons) at the Neoproterozoic (no older than 800 Ma) (Dmitrieva et al., 2016). The Karsakpai Group comprises mostly mafic volcanics and initially sedimentary rocks. Volcanic and sedimentary

1068-7971/$ - see front matter D 201 7, 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.2017.07.004 +

936

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Fig. 1. Map of the geologic structure of southern Ulutau, after Zaitsev and Filatova (1971), simplified. 1, Paleozoic and Mesozoic complexes; 2, Vendian sedimentary and volcanic strata; 3–8, Proterozoic complexes: 3, volcanosedimentary, of the Bozdak Group; 4, 5, volcanic and volcanosedimentary, of the Maityuba zone (II): 4, Koksui Group, 5, Maityuba and Zhiida Groups; 6, 7, of the Karsakpai zone (I): 6, Beleuta and Karsakpai Groups, 7, Aralbai Group; 8, amphibolites and schists of the Bekturgan Group; 9–11, intrusive rocks: 9, Late Ordovician granodiorites, 10, syenites of the Karsakpai complex, 11, Proterozoic granitoids (Zhaunkar and Aktas complexes); 12, discontinuities; 13, localities of sampling for Sm–Nd isotope studies. Rectangles mark the sampling areas.

rocks are grouped into thick members, thus forming coarse rhythmicity, which served as a basis for the stratigraphic division of the group into four formations (Rozanov, 1976) 700–800 to 1200 m in thickness. In the lower two formations, Burmash and Bolbraun, the lower members are metabasaltic, with impersistent interbeds of quartz–sericite phyllites and marbles, and the upper thinner (200–300 m) members are of quartzite–phyllite composition, with interbeds of carbonate, siliceous, and siliceous-ferruginous rocks. In the upper part of the group, in the Shagyrly and Biit Formations, metabasalts coexist with metarhyolites and their tuffs. Ferruginous quartzites are scarce here (Filatova and Rozanov, 1995). The presence of rocks with a relict parallel thin-laminated structure

in almost all parts of the section, the signs of cross-bedded structures in tuffites, and the persistence of volcanic and sedimentary rock sections over a large area evidence that most of the group rocks accumulated in a sea basin (Zaitsev and Rozanov, 1971). We studied metavolcanosedimentary rocks of two lower formations (Burmash and Bolbraun) in the region of Karsakpai Village and south of it (Fig. 1). In the region of Karsakpai Village, there are outcrops of the lower member of the Burmash Formation, composed mostly of greenschists (metamorphosed by basaltic tuffs and tuffites) with an impersistent interbed of quartz–sericite phyllites, schists, and quartzites. The Bolbraun Formation, developed south of Karsakpai Village, is made up of metamorphosed basaltic volcanics with interbeds of siliceous and siliceous-ferruginous rocks. The typical composition of metabasalts is as follows: actinolite, epidote, albite, chlorite, ore mineral, and, more seldom, biotite, apatite, and zoisite. Primary phenocrysts (plagioclase) are often albitized and saussuritized. The rocks are of porphyritic textures and massive (or, more seldom, schistose) structures. Chlorite–actinolite greenschists are intimately associated with metabasalts in the sections. In composition and structure they are close to the metabasalt groundmass. The mineral composition of the greenschists is as follows: amphibole, chlorite, epidote, zoisite, plagioclase, biotite, sericite, hematite, magnetite, and calcite. The rocks are of nematoblastic and lepidoblastic textures and schistose structures. They formed, most likely, through metamorphism of tuffaceous rocks. Metamorphosed sedimentary rocks are quartz–sericite– chlorite schists (phyllites), quartzite schists, quartzites, and ferruginous quartzites. The schists are of lepidogranoblastic and lepidoblastic textures and schistose structures. Some varieties preserved primary lamination. The rocks have the following mineral composition: quartz, sericite, chlorite, muscovite, hematite, magnetite, apatite, tourmaline, and sphene. The rocks are usually fine-grained, though sometimes they contain coarser quartz and tourmaline grains. Quartzites consist mostly of quartz (up to 80–85%), sericite, and muscovite and dispersed fine crystals of hematite and magnetite. They are of granoblastic, microgranoblastic, and, more seldom, lepidogranoblastic textures and massive, sometimes schistose, structures. Ferruginous quartzites contain ≥10–15% (sometimes, up to 50–60%) ore minerals. These are black, clearly banded rocks with the following mineral composition: hematite, magnetite, quartz, chlorite, sericite, and, more seldom, pyrite, apatite, albite, and carbonate. The rocks are of microgranoblastic and lepidogranoblastic textures.

Methods To study the lithologic composition and isotope parameters of metamorphic rocks of the Karsakpai Group, we sampled their most typical mineralogic/petrographic varieties. Analyses of rock-forming elements were performed by a wet chemical technique at the Analytical Center of the Institute of the

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Earth’s Crust, Irkutsk (analysts G.V. Bondareva and E.G. Koltunova). Contents of trace and rare-earth elements were determined by ICP MS on a high-resolution ELEMENT mass spectrometer (detection limits of 0.005 to 0.1 ppm) at the Analytical Center of the Institute of Geology and Mineralogy, Novosibirsk (analyst I.V. Nikolaeva). The error of analyses was 2–5%. The specimens were prepared following the chemical technique borrowed from Nikolaeva et al. (2008). The Nd and Sm contents and Nd isotope composition for all samples, except for M-40-13, were determined at the Geological Institute, Apatity, following the technique described by Bayanova (2004). They were measured on a Finnigan MAT-262 (RPQ) seven-channel solid-phase mass spectrometer with a two-ribbon (Re–Ta) ion source. The average 143Nd/144Nd ratio in the La Jolla standard (Lugmair and Carlson, 1978) was 0.511835 ± 0.000018 (n = 15). The error of 147Sm/144Nd determination, an average over seven measured values in the BCR standard (Raczek et al., 2003), was 0.4% (2σ). The error of measurement of the Nd isotope composition in individual analysis was no more than 0.004%. The blank sample contained 0.3 ng Nd and 0.06 ng Sm. The accuracy of Sm and Nd analyses was ±0.5 %. The measured isotope ratios were normalized to 146Nd/144Nd = 0.7219 and then converted to 143Nd/144Nd = 0.511860 in the La Jolla standard. The Sm and Nd isotope compositions of the sample M-40-13 were determined at the Baikal Center of Common Use, Irkutsk, following a modified technique (Pin and Zalduegui, 1997). The 143Nd/144Nd ratios were measured on a Finnigan MAT-262 mass spectrometer. The 147Sm/144Nd ratios were calculated from the Sm and Nd contents measured by ICP MS. The determined isotope ratios were normalized to 146Nd/144Nd = 0.7219. Correction for fractionation was made by the Raleigh method. To control the validity of obtained data, a JNd-1 Nd standard (Tanaka et al., 2000) was used (143Nd/144Nd = 0.512102 ± 0.000009, n = 20). The εNd(T) value and model age TNd(DM) were calculated from the modern CHUR (chondrite uniform reservoir) values, 143 Nd/144Nd = 0.512638 and 147Sm/144Nd = 0.1967 (Jacobsen and Wasserburg, 1984), and DM (depleted mantle) values, 143 Nd/144Nd = 0.513151 and 147Sm/144Nd = 0.2136 (Goldstein and Jacobsen, 1988).

Results Contents of major, trace, and rare-earth elements in the representative samples of the Karsakpai Group are listed in Tables 1 and 3. Metasedimentary rocks Major elements. Analysis of ferruginous quartzites showed that Fe2O3tot (29.67–71.13%) and SiO2 (25.74–67.38%) are major components and Al2O3 (0.57–4.45%) is next in content. The contents of TiO2, MgO, MnO, CaO, Na2O, K2O, and P2O5 do not exceed 0.5% (Table 1). The calculated Pearson

937

correlation coefficients (Table 2) demonstrate significant positive correlations of TiO2 with Al2O3, K2O, and P2O5, of Al2O3 with K2O and P2O5, of Fe2O3tot with MgO and CaO, and between CaO and MgO and a strong negative correlation between Fe2O3tot and SiO2. In barren quartzites, as in the ferruginous ones, major components are SiO2 (90.73–92.32%) and Fe2O3 (4.91–6.98%). The highest contents of Al2O3 (8.68–16.69), TiO2 (0.38–2.47%), and K2O (0.45–3.71%) (components typical of clastic rocks) are observed in quartz– sericite–chlorite schists. The schists of the Bolbraun Formation are characterized by domination of K2O over Na2O and reduced Fe2O3tot, TiO2, and Fe2O3tot + MgO contents as compared with the schists of the Burmash Formation (Table 1). According to the classification by Pettijohn et al. (1972), the schists of the Bolbraun Formation correspond in composition to schists and wackes, and those of the Burmash Formation, to ferruginous sandstones. Trace and rare-earth elements. The contents of trace elements in the metasedimentary rocks of the Karsakpai Group vary significantly. The ferruginous and barren quartzites are depleted in almost all trace elements relative to the PostArchean Australian Shale (PAAS) (Taylor and McLennan, 1985) (Fig. 2). The schists and quartzite schists of the Karsakpai Group are enriched in all trace elements relative to the quartzites and are close in their contents to PAAS. Compared with the schists of the Burmash Formation, the schists of the Bolbraun Formation are enriched in Y, Zr, Rb, Hf, and Th (elements geochemically related to felsic rocks) and are depleted in Sc, V, Cr, and Co (elements typical of mafic igneous rocks). The contents of REE and Y are listed in Table 1. The studied metasedimentary rocks are characterized by varying total REE contents. The schists show the maximum values (ΣREE = 60–241 ppm), especially the Bolbraun Formation ones; the barren quartzites demonstrate the minimum values (ΣREE = 8–17 ppm); and the ferruginous quartzites show intermediate values (ΣREE = 28–75 ppm), except for one sample with ΣREE = 241 ppm. The PAAS—normalized REE patterns of the Karsakpai Group metasediments are presented in Fig. 3. All rocks are depleted in LREE. The (La/Yb)PAAS values vary from 0.12 to 0.68. The schists are characterized by the maximum (La/Yb)PAAS values close to those of the Karsakpai Group basalts (0.5). All rocks show no Eu anomaly ((Eu/Eu*)PAAS = 0.9–1.1), except for the Burmash Formation schists ((Eu/Eu*)PAAS = 1.4), no (or a weak negative) Ce anomaly ((Ce/Ce*)PAAS = 0.83–1.03), and low Y/Ho ratio (24–34). Thus, the highest contents of Al2O3, TiO2, and K2O (components typical of clastic rocks) are observed in quartz– sericite–chlorite schists. These rocks also have the highest ΣREE and elevated (relative to the quartzites) contents of other trace elements. In (La/Yb)PAAS values they are similar to the Karsakpai Group basalts. These schists are, most likely, metamorphosed argillaceous rocks containing a clastic material, mostly mafic. The Bolbraun Formation schists are characterized by domination of K2O over Na2O, reduced (relative to the Burmash Formation rocks) contents of Fe2O3tot, TiO2,

938

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Table 1. Contents of major (wt.%) and trace (ppm) elements in representative samples of the Karsakpai Group metasediments Component

Burmash Formation

Bolbraun Formation

D-47-09

D-53-09

D-49-09

D-54-09

D-34-09

M-39-13 D-33-09

M-43-13 M-38-13 D-32-09

M-46-13 M-40-13 M-41-13

1

2

3

4

5

6

7

8

9

10

11

12

13

SiO2

57.06

68.82

79.22

90.73

25.74

31.75

32.64

44.39

61.21

67.38

64.96

80.21

92.32

TiO2

2.47

1.6

0.38

0.04

0.2

0.23

0.17

0.21

0.41

0.07

0.89

0.63

0.04

Al2O3

11.9

9.1

11.2

0.25

1.35

1.76

0.8

1.76

4.45

0.57

16.69

8.68

0.91

Fe2O3

8.58

4.97

0.71

6.98

69.38

63.41

65.21

51.35

29.12

30.59

2.14

5.69

4.91

FeO

3.57

3.83

1.31

1.48

1.59

0.82

0.67

0.76

0.50

0.85

2.70

0.25

0.48

MnO

0.13

0.09

0.02

0.03

0.02

0.05

0.01

0.03

0.01

0.01

0.08

0.01

0.01

MgO

3.93

2.8

0.62

0.11

0.4

0.16

0.05

0.05

0.11

0.05

3.65

0.53

0.13

CaO

5.81

3.52

0.53

0.53

0.38

0.12

0.05

0.05

0.05

0.05

0.32

0.05

0.05

Na2O

3.44

4.14

4.1

0.03

0.06

0.02

0.04

0.04

0.05

0.06

0.22

0.13

0.06

K2O

0.45

0.59

0.95

0.01

0.13

0.01

0.01

0.11

0.01

0.03

3.71

2.31

0.04

P2O5

0.26

0.14

0.11

0.25

0.25

0.21

0.1

0.09

0.50

0.25

0.14

0.04

0.11

LOI

1.75

0.66

0.97

0.05

0.8

1.13

0.76

0.95

3.27

0.53

3.98

1.82

0.46

Total

99.63

100.39

100.49

100.61

100.36

99.81

100.55

99.88

99.90

100.46

99.56

100.44

99.58

Sc

31

19

7

1.61

18

8.0

19

6.0

25

3.0

18.2

15.9

1.61

V

259

180

53

14.2

110

136

160

68

37

38

49

66

11.1

Cr

175

64

45

47

47

33

26

43

37

34

24

44

25

Co

37

15

14

4.4

6

5

6

7.5

5.3

2.2

6.0

2.5

1.71

Ni

75

34

28

13.0

9

3

<5

15.2

18.8

6.6

78

28

5.2

Cu

87

58

190

29

24

45

24

18.3

45

18.0

44

53

18.4

Zn

102

52

44

9.2

50

24

48

18.3

29

8.0

109

42

11.4

Rb

3.9

6.9

31

0.39

6

1.15

1.16

4.2

0.26

0.95

109

119

1.12

Sr

158

50

84

9.2

84

107

38

22

1 293

86

29

20

203

Y

26

17.2

22

8.6

11

27

37

12.7

64

19.3

50

38

3.8

Zr

161

95

116

15.1

43

75

68

58

75

12.1

271

247

19.6

Nb

15.0

9.6

5.6

1.00

7

7.9

4.5

3.1

4.3

1.19

14.3

19.3

0.50

Ba

100

592

138

6.8

61

94

28

92

682

2 363

1 189

1 512

49

La

12.1

9.0

17.4

2.4

<5

11.2

7.6

8.1

36

3.3

43

29

1.22 2.4

Ce

31

19.9

35

5.5

12

27

18.3

16.5

68

7.9

97

56

Pr

4.4

2.8

4.3

0.72

–

3.3

2.5

2.1

10.0

1.00

12.1

7.5

0.30

Nd

19.4

12.9

16.3

3.0

<9

13.7

9.9

8.0

50

4.4

45

28

1.24

Sm

5.2

2.9

3.6

0.81

–

3.5

2.2

1.43

16.9

1.11

9.2

5.8

0.33

Eu

1.73

1.01

0.90

0.21

–

0.80

0.55

0.29

4.9

0.31

2.1

1.08

0.11

Gd

5.8

3.7

4.1

1.09

–

3.8

3.1

1.35

22

1.63

9.6

5.9

0.56

Tb

0.93

0.53

0.63

0.20

–

0.62

0.67

0.23

3.2

0.34

1.50

0.97

0.14

Dy

5.3

3.2

3.3

1.28

–

3.8

4.4

1.65

16.3

2.7

8.5

5.8

0.88

Ho

1.05

0.58

0.68

0.28

–

0.91

1.12

0.37

2.7

0.63

1.70

1.34

0.14

Er

2.5

1.53

1.90

0.81

–

2.8

3.5

1.21

5.8

1.80

5.0

3.9

0.40

Tm

0.35

0.22

0.30

0.13

–

0.43

0.60

0.23

0.70

0.30

0.78

0.64

0.060

Yb

2.1

1.20

1.90

0.81

–

3.0

4.1

1.69

3.9

2.0

5.2

4.3

0.41

Lu

0.29

0.18

0.29

0.12

–

0.46

0.64

0.27

0.54

0.30

0.76

0.64

0.060

Hf

4.1

2.6

3.3

0.17

–

1.31

0.99

0.99

1.99

0.20

8.4

7.5

0.34

Ta

0.98

0.72

0.44

<0.05

–

1.10

3.9

0.72

0.57

<0.05

0.90

1.61

0.060

Th

1.2

0.7

4.7

0.1

1.5

1.9

1.0

1.5

6.4

0.3

8.3

5.4

1.7

U

0.3

0.4

3.0

0.5

0.7

0.8

0.9

0.7

3.2

1.1

2.1

0.9

0.8

(continued on next page)

939

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948 Table 1 (continued) Component

Burmash Formation

Bolbraun Formation

D-47-09

D-53-09

D-49-09

D-54-09

D-34-09

M-39-13 D-33-09

M-43-13 M-38-13 D-32-09

M-46-13 M-40-13 M-41-13

1

2

3

4

5

6

7

8

9

10

11

Fe2O3tot

12.51

9.18

2.15

8.61

71.13

64.31

65.95

52.19

29.67

31.53

5.12

5.97

5.44

Fe2O3tot + MgO

16.44

11.98

2.77

8.72

71.53

64.47

66.00

52.24

29.78

31.58

8.77

6.50

5.57

12

13

ΣREE

92

60

91

17

–

75

59

43

241

28

241

150

8

(Eu/Eu*)PAAS

1.44

1.42

1.08

1.01

–

1.03

0.95

0.96

1.15

1.04

1.03

0.86

1.15

(Ce/Ce*)PAAS

0.95

0.91

0.93

0.97

–

1.03

0.96

0.92

0.83

1.01

0.98

0.88

0.94

(La/Yb)PAAS

0.42

0.55

0.67

0.21

–

0.27

0.14

0.35

0.68

0.12

0.61

0.49

0.22

Y/Ho

24.9

29.3

31.8

30.5

–

29.7

33.0

34.3

23.8

30.9

29.7

28.6

27.0

Note. Burmash Formation: 1–2, quartz–sericite–chlorite schists, 3, quartzite schist, 4, quartzite; Bolbraun Formation: 5–10, ferruginous quartzites, 11, quartz–sericite–chlorite schist, 12, quartzite schist, 13, quartzite. ΣREE does not include Y; (Eu/Eu*)PAAS = 2⋅EuPAAS/(SmPAAS + GdPAAS), (Ce/Ce*)PAAS = 2⋅CePAAS/(LaPAAS + PrPAAS), (La/Yb)PAAS = LaPAAS/YbPAAS, PAAS-normalized (Taylor and McLennan, 1985). Table 2. Pearson correlation coefficients for major components of the Karsakpai Group ferruginous quartzites SiO2

TiO2

Al2O3

Fe2O3tot

MnO

MgO

CaO

Na2O

K2O

SiO2

1

TiO2

0.06

1

Al2O3

0.31

0.96

1

Fe2O3tot

–0.99

–0.21

–0.45

1

MnO

–0.55

–0.58

–0.60

0.56

1

MgO

–0.70

–0.35

–0.42

0.66

–0.15

1

CaO

–0.76

–0.48

–0.54

0.73

–0.11

0.99

1

Na2O

0.40

–0.18

–0.05

–0.37

–0.89

0.52

0.52

1

K2O

–0.41

0.65

0.85

0.36

–1.00

1.00

1.00

0.20

1

P2O5

0.50

0.70

0.80

–0.61

–0.63

0.05

–0.10

0.33

0.17

P2O5

1

Note. Bold values are discussed in the text. Correlation is significant at the 0.05% level (critical value of the correlation coefficient of 0.55).

(Fe2O3tot+MgO), Sc, V, and Cr, and elevated contents of Y, Zr, Rb, Hf, and Th, which might indicate the presence of a terrigenous material formed during the erosion of felsic rocks. Metavolcanic rocks Analysis of the Karsakpai Group metabasaltoids yielded the following results (wt.%): SiO2—46.83–51.03, Al2O3—13.55– 15.30, MgO—5.21–6.8, FeOtot—10.00–15.56, Na2O—2.97– 4.06, K2O—0.12–1.00, total alkalies—3.17–4.87, and TiO2— 1.24–3.46 (Table 3). In SiO2 and Na2O + K2O contents the rocks correspond to basalts (Fig. 4a). On the Al2O3–(TiO2 + FeOtot)–MgO diagram (Jensen, 1976), the rocks fall in the field of high-Fe tholeiites (Fig. 4c). The studied rocks have medium MgO contents (5.21–6.80 wt.%). According to alkaline-element patterns, they are classified as low- and mediumK (Fig. 4b), essentially sodic (Na2O/K2O > 4%). The metabasites are moderately differentiated varieties: Their Mg# value varies from 41 to 57. Higher values are observed in the metabasalts of the Bolbraun Formation. All studied rocks, except for one sample from the Bolbraun Formation, are of high-Ti composition (TiO2 > 1.5 wt.%).

The metabasaltoids are enriched in LREE and are characterized by fractionated chondrite-normalized (Boynton, 1984) REE patterns ((La/Yb)n = 3.8–5.3). The maximum (La/Yb)n value is determined for the low-Ti sample (M-42-13, Table 3) from the Bolbraun Formation. The (La/Sm)n value varies from 1.5 to 2.9, being higher in the rocks of the Bolbraun Formation. All samples show a negligible Eu anomaly (Eu/Eu* = 0.9–1.1) (Table 3, Fig. 5a). The basaltoids are significantly enriched in LILE and LREE relative to N-MORB. The primitive-mantle-normalized (Sun and McDonough, 1989) patterns of elements are shown in Fig. 5b. They lie between the patterns of E-MORB and OIB. A specific feature of the metabasaltoids of the Burmash Formation is a positive Nb anomaly (Nb/Nb* = 1.2–1.3), which is typical of basalts resulted from the impact of mantle plumes. The metabasalts of the Bolbraun Formation show a weak negative Nb anomaly (0.76–0.91), which drastically increases in the low-Ti sample (0.31), with the synchronous appearance of a negative Ti anomaly. This suggests the influence of crustal material on the composition of the volcanics. The values of (Nb/Th)PM (0.26–1.05 to 1.31–1.58) and (Nb/La)PM (0.37–0.78 to 1.00–1.08) increase from the

940

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

compared with the standard depleted mantle (εNd(1000) = +7.9 (Goldstein and Jacobsen, 1988)). The studied ferruginous quartzites show similar εNd(t) values, +3.2 to +5.0. The model Nd age of the schists is within 1.8–1.3 Ga, and εNd(t) = –0.3 to +4.7. Discussion

Fig. 2. PAAS-normalized (Taylor and McLennan, 1988) trace-element patterns of the Karsakpai Group metasediments: ferruginous (solid lines) and barren (dotted lines) quartzites (a) and schists and quartzite schists of the Bolbraun (solid lines) and Burmash (dotted lines) Formations (b). The rock compositions are given in Table 1.

metabasalts of the Bolbraun Formation to those of the Burmash Formation. The (Th/La)PM values are <1 (0.72–0.93) in all samples, except for the low-Ti one (Table 3). Thus, the metabasaltoids of the Karsakpai Group correspond to essentially sodic low- and medium-K high-Fe high-Ti (except for one sample from the Bolbraun Formation) tholeiitic basalts and are moderately differentiated varieties. In traceelement composition they are similar to intraplate volcanics. The basalts of the Bolbraun Formation might have been impacted by crustal material. The Nd isotope classification of the Karsakpai Group rocks The metabasltoids of the Karsakpai Group are characterized by both positive and negative εNd(t) values (+3.4 to +8.0 and –6.8) calculated for their minimum possible age of ~1000 Ma (Table 4). According to these data, the metabasaltoids of the Burmash Formation are no older than 1000 Ma (εNd(1000) = +8.0), otherwise they should have formed from a strongly depleted mantle source with much higher εNd values as

Metasedimentary rocks. Data on the distribution of major elements in ferruginous rocks give an insight into the rock genesis (hydrothermal or hydrogenous (Bostrom, 1970; Cox et al., 2013; Gurvich, 1998)). The strong positive correlation of K with Al and Ti suggests the same mechanism of their supply into the sediments, most likely, from a detrital source. The positive correlations between Mg and Ca obviously indicate their coexistence in carbonate minerals. Moreover, the Karsakpai Group metasediments show a positive correlation of Zr with Al2O3 and TiO2 (Fig. 6), which suggests the presence of a clastic impurity in the sediments (Ewers and Morris, 1981). The contents of Al2O3, TiO2, Zr, and Nb are maximum in the Karsakpai Group schists, which agrees with the detrital source of material. The high MgO + Fe2O3tot contents (12–16%) and reduced (relative to PAAS) K2O/Na2O ratios (0.1–0.2) point to a significant portion of mafic rocks in the source of the Burmash Formation schists. The Bolbraun Formation schists show domination of K2O over Na2O and reduced Fe2O3tot, TiO2, and (Fe2O3tot + MgO) contents, which suggests the presence of terrigenous material formed during the erosion of igneous and metamorphic felsic rocks. Hydrothermal fluids are characterized by high Fe/Ti, Fe/Al, and Si/Al ratios (Bostrom, 1970; Bostrom et al., 1969; Gurvich, 1998). The ferruginous quartzites of the Karsakpai Group show high Si/Al ratios (>4) typical of sediments formed mostly at the expense of hydrothermal component (Gurvich, 1998). As seen from the Fe/Ti–Al/(Al + Fe + Mn) diagram (Bostrom, 1973), often used for the determination of clastic and hydrothermal components in ferruginous rocks, the Karsakpai Group quartzites might have resulted from the mixing of hydrothermal and detrital components. The clastic material was, most likely, basalts or volcanic sediments rather than typical PAAS (Fig. 7a). The schists apparently lack a significant impurity of exhalative component. The difference in the chemical composition of quartzites and schists is seen from the Al–Fe–Mn ternary diagram (Bostrom, 1973): Quartzites fall in the field of hydrothermal sediments, and schists fall in the field of nonhydrothermal deposits (Fig. 7b). Study of REE patterns is one of crucial geochemical tools for understanding the genesis of iron formations (Bau and Dulski, 1996; Bekker et al., 2010; Fryer, 1977; Klein and Beukes, 1989). Such patterns and the presence/absence of element anomalies provide information about the participation of hydrothermal solutions in the above genesis and help to reconstruct the redox conditions in water basins and the degree of interaction of the solutions with seawater. Figure 3 presents the REE patterns of the studied rocks in comparison with those of hydrothermal solutions (Douville et al., 1999; Michard et

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

941

Fig. 3. PAAS-normalized (Taylor and McLennan, 1988) trace-element patterns of the metasedimentary rocks of the Karsakpai Group of the Burmash (a) and Bolbraun (b) Formations in comparison with sea water (×106) (Zhang and Nozaki, 1996) (1), high-temperature hydrothermal fluids (×106) (average for high-temperature hydrothermal fluids of the East Pacific uplift, 13º N, 17º–19º E (Douville et al., 1999)) (2), low-temperature fluids (Michard et al., 1993) (3), and typical metabasites of the Karsakpai Group (sample D-50-09). The numerals follow the sample numbers in Tables 1 and 3.

al., 1993) and seawater (Zhang and Nozaki, 1996). The hydrothermal solutions of the East Pacific Rise are slightly enriched in LREE and show a strong positive Eu anomaly (Bau and Dulski, 1999; Douville et al., 1999; Michard and Albarède, 1986; Michard et al., 1983); the low-temperature solutions show no Eu anomaly (Michard et al., 1993). The seawater is enriched in HREE and has negative Ce and positive Y anomalies (Elderfield and Greaves, 1982; Piepgras and Jacobsen, 1992; Zhang and Nozaki, 1996). The PAASnormalized REE patterns of the ferruginous-siliceous rocks of the Karsakpai Group show their LREE depletion ((La/Yb)PAAS = 0.12–0.27 to 0.68). This might be due to the interaction of hydrothermal solutions with seawater, the inheritance of the REE characteristics of detrital material, or both factors. As seen from Fig. 3, the degree of LREE depletion varies, which seems to reflect the different amounts

of detrital material. The schists and quartzite schists show elevated (La/Yb)PAAS ratios (0.42–0.67) close to those in the Karsakpai Group basalts (0.46–0.57) or somewhat higher, which suggests the participation of felsic La-enriched rocks in the formation of detrital material. Yttrium is chemically very similar to REE (Henderson, 1984), particularly, Ho (Bau and Dulski, 1999). Clastic sedimentary rocks (e.g., PAAS (Taylor and McLennan, 1985)), upper continental crust (Rudnick and Gao, 2003), and oceanic crust (Sun and McDonough, 1989) are characterized by Y/Ho close to those in chondrite (~26), because Y and Ho have close ionic radii, charge densities, and degrees of oxidation. Hydrothermal solutions do not have superchondrite Y/Ho ratios (Douville et al., 1999) in contrast to seawater, in which Ho is more reactive (Bau et al., 1996; Bau and Dulski, 1999). The Karsakpai Group quartzites have a somewhat lower

942

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Table 3. Contents of major (wt.%) and trace (ppm) elements in representative samples of Karsakpai Group metabasaltoids Component

Burmash Formation

Bolbraun Formation

D-50-09

D-51-09

D-52-09

D-46-09

M-45-13

M-44-13

M-42-13

1

2

3

4

5

6

7

SiO2

46.83

47.24

48.34

49.44

47.82

49.54

51.03

TiO2

3.46

3.08

2.91

2.50

2.44

2.27

1.24

Al2O3

14.65

13.75

13.55

14.6

14.62

14.01

15.30

Fe2O3

8.63

5.45

8.41

3.89

5.24

4.43

3.55

FeO

7.79

9.65

6.13

8.44

8.33

8.75

6.80

MnO

0.17

0.19

0.17

0.19

0.20

0.20

0.16

MgO

5.21

5.61

5.33

5.54

6.80

6.02

6.37

CaO

4.58

5.22

6.25

8.00

7.60

7.34

8.54

Na2O

3.82

4.1

3.69

4.06

2.97

4.05

3.71

K2O

0.66

0.77

1.00

0.28

0.20

0.12

0.42

P2O5

0.38

0.35

0.27

0.19

0.28

0.25

0.23

LOI

2.94

3.32

3.40

2.94

3.31

3.33

2.99

Total

99.88

99.68

99.89

100.21

99.94

100.40

100.40

Sc

32

37

38

31

41

43

33

V

390

333

341

295

370

363

177

Cr

59

71

110

182

80

94

166

Co

38

37

45

32

42

47

32

Ni

67

47

56

80

43

47

34

Cu

140

104

129

106

87

155

26

Zn

140

148

147

143

152

139

96

Rb

12.1

8.6

18.7

8.9

8.4

2.0

4.1

Sr

131

105

218

169

439

178

262

Y

37

34

31

29

33

34

32

Zr

192

210

192

190

167

194

145

Nb

17.6

19.5

16.2

16.5

13.7

13.8

9.5

Cs

1.50

1.38

3.4

2.4

0.33

0.67

0.39

Ba

578

207

609

387

291

76

225

La

16.9

18.3

14.5

15.4

16.9

18.2

25

Ce

39

42

35

36

37

40

48

Pr

5.8

5.7

5.0

5.0

5.1

5.5

6.6

Nd

26

25

22

22

21

23

25

Sm

6.6

6.8

6.1

5.8

5.8

6.0

5.3

Eu

2.3

2.1

2.0

1.89

2.1

1.98

1.61

Gd

7.5

7.6

7.0

6.4

6.3

6.8

5.6

Tb

1.16

1.19

1.13

1.05

1.07

1.14

0.95

Dy

6.7

6.7

6.4

6.0

6.4

7.0

5.5

Ho

1.30

1.25

1.16

1.11

1.27

1.34

1.10

Er

3.2

3.2

3.0

2.8

3.5

3.6

3.1

Tm

0.48

0.43

0.40

0.39

0.51

0.54

0.48

Yb

2.7

2.6

2.4

2.4

3.0

3.4

3.1 0.46

Lu

0.35

0.38

0.35

0.35

0.45

0.49

Hf

5.1

5.3

4.9

4.8

4.3

4.9

4.4

Ta

1.20

1.26

1.10

1.19

0.92

0.99

0.60

Th

1.6

1.7

1.3

1.3

1.6

2.1

4.3

U

0.7

0.6

0.5

0.9

0.7

0.6

0.6

Na2O + K2O

4.48

4.87

4.69

4.34

3.17

4.17

4.13 (continued on next page)

943

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948 Table 3 (continued) Component

Burmash Formation

Bolbraun Formation

D-50-09

D-51-09

D-52-09

D-46-09

M-45-13

M-44-13

M-42-13

1

2

3

4

5

6

7

Na2O/K2O

0.83

0.79

0.59

0.47

0.39

0.55

0.43

FeO*

15.56

14.56

13.70

11.94

13.05

12.74

10.00

Mg#

41

45

45

49

52

50

57

(La/Yb)n

4.20

4.73

4.06

4.33

3.77

3.65

5.28

(La/Sm)n

1.62

1.69

1.49

1.67

1.82

1.90

2.93

Eu/Eu*

1.02

0.89

0.93

0.94

1.05

0.94

0.90

Nb/Nb*

1.15

1.20

1.28

1.28

0.91

0.76

0.31

(Nb/Th)PM

1.31

1.41

1.51

1.58

1.05

0.78

0.26

(Nb/La)PM

1.00

1.03

1.08

1.04

0.78

0.73

0.37

(Th/La)PM

0.77

0.73

0.72

0.66

0.74

0.93

1.42

Note. Burmash Formation: 1–3, chlorite–actinolite schists, 4, metabasalt; Bolbraun Formation: 5, chlorite–actinolite schist, 6–7, metabasalts. FeO*, Total iron; Mg# = 100⋅Mg/(Mg + Fe2+), where Fe2+/Fetot = 0.85, Eu/Eu* = 2⋅Eun/(Smn + Gdn), (La/Yb)n = Lan/Ybn, (La/Sm)n = Lan/Smn, chondrite-normalized (Boynton, 1984); Nb/Nb* = NbPM/((ThPM × LaPM)0.5, (Nb/Th)PM = NbPM/ThPM, (Nb/La)PM = NbPM/LaPM, (Th/La)PM = ThPM/LaPM, primitive-mantle(PM)-normalized (Sun and McDonough, 1989).

Fig. 4. SiO2–(Na2O + K2O) (a), SiO2–K2O (b) (Le Maitre et al., 1989), and MgO–(FeO* + TiO2)–Al2O3 (Jensen, 1976) (c) classification diagrams for the metabasalts of the Burmash (1) and Bolbraun (2) Formations of the Karsakpai Group. BK, Basaltic komatiites; CA, calc-alkalic andesites; CB, calc-alkalic basalts; CD, calc-alkalic dacites; CR, calc-alkalic rhyolites; PC, picrites; HFT, high-Fe tholeiites; HMT, high-Mg tholeiites; TA, tholeiitic andesites; TD, tholeiitic dacites; TR, tholeiitic rhyolites.

944

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Table 4. Sm-Nd isotope data on the Karsakpai Group rocks Sm, ppm

Nd, ppm

147

143

TNd(DM), Ma TNd(DM-2st), Ma εNd(T)

2.60

11.80

0.1341

0.512412 ± 9

1414

M-43-13 (ferruginous quartzite)

1.30

7.70

0.1053

0.512292 ± 9

M-40-13 (quartzite schist)

5.80

28.0

0.1247

0.512408 ± 10

M-46-13 (schist)

9.50

44.2

0.1300

0.512184 ± 4

M-42-13 (metabasite)

5.30

24.5

0.1310

M-45-13 (metabasalt)

8.80

32.5

0.1630

D-49-09 (quartzite schist)

3.90

13.50

D-50-09 (metabasite)

7.40

29.0

Sample

Sm/144Nd

Nd/144Nd ± 2σmeas

Bolbraun Formation M-39-13 (ferruginous quartzite)

1351

+3.2

1208

1240

+5.0

1273

1258

+4.7

1758

1677

–0.3

0.511859 ± 14

2372

2211

–6.8

0.512592 ± 15

–

1367

+3.4

0.1761

0.512590 ± 9

–

1510

+1.7

0.1550

0.512776 ± 8

–

983

+8.0

Burmash Formation

Note. The εNd(T) and TNd(DM-2st) values were calculated for an age of 1000 Ma.

average Y/Ho ratio (~31) than seawater (>44), which might be due to an impurity of detrital material. The schists show Y/Ho ratios close to those in chondrite. The studied Karsakpai Group rocks show no or a weak negative Ce anomaly ((Ce/Ce*)PAAS = 0.88–1.03) and no or

a positive Eu anomaly ((Eu/Eu*)PAAS = 0.9–1.4). The absence of a Eu anomaly might be inherited from low-temperature hydrothermal solutions or high-temperature hydrotherms that lost their positive Eu anomaly during the interaction with seawater, basalts, and terrigenous sediments. The spread in Eu

Fig. 5. Chondrite-normalized (Boynton, 1984) REE (a) and PM-normalized (Sun and McDonough, 1989) trace-element (b) patterns of the metabasites of the Burmash (solid lines) and Bolbraun (dotted lines) Formations of the Karsakpai Group in comparison with E-MORB, N-MORB, and OIB (Sun and McDonough, 1989). The numerals follow the sample numbers in Table 3.

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

945

Fig. 6. Zr–Al2O3 and Zr–TiO2 diagrams demonstrating positive correlations between the components of the rocks of the Burmash (light symbols) and Bolbraun (black symbols) Formations of the Karsakpai Group. Triangles mark quartzites, and squares mark quartzite schists and schists.

Fig. 7. Fe/Ti–Al/(Fe + Mn + Al) (a) and Fe–Al–Mn (b) diagrams (Bostrom, 1973) for the Karsakpai Group metasediments. The mixing lines were calculated for hydrothermal sediments (metal-bearing sediments of the Red Sea (Gurvich, 1998)), PAAS (Taylor and McLennan, 1985), N-MORB, and OIB (Sun and McDonough, 1989). 1, 2, Burmash Formation: 1, schists, 2, quartzite; 3–5, Bolbraun Formation: 3, ferruginous quartzites, 4, schists, 5, quartzites.

anomaly values is, most likely, due to the varying contribution of a detrital component, which agrees with the varying degrees of LREE depletion and Y/Ho ratios. To sum it up, the formation of the ferruginous-siliceous sediments of the Karsakpai Group is intimately related to submarine eruptions of mafic lavas and the activity of hydrothermal solutions (probably, with an impurity of a detrital component). We have established that the Karsakpai Group schists lack an exhalative material. These are rocks with a significant portion of erosion products of mafic igneous deposits. The Bolbraun Formation schists might additionally contain erosion products of more mature La- and Th-enriched felsic rocks. The 143Nd/144Nd ratios (εNd(t)) can be used to establish the source of the material of sedimentary rocks (Cox et al.,

2016; Goldstein et al., 1984). The studied ferruginous quartzites are characterized by εNd(t) = +3.2 to +5.0. The Nd model ages of the schists are within 1.8–1.3 Ga, and εNd(t) = –0.3 to +4.7. Thus, the formation of the ferruginous-siliceous rocks of the Karsakpai Group is apparently related to hydrothermal activity and submarine volcanism, which resulted in sediments with 143Nd/144Nd close to those in the mantle rocks. The isotope-geochemical data of the studied metaterrigenous rocks of the Karsakpai Group (the Meso-Proterozoic model Nd ages and positive εNd(t) values) testify to the existence of Mesoproterozoic juvenile material in the provenances. The more ancient Nd model age TNd(DM) = 1.8 Ga and εNd(t) = –0.3 for the chlorite schist reflect the supply of sedimentary material from more ancient sources.

946

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

(Pearce, 2008) (Fig. 8), the composition points of the Bolbraun Formation metabasalts are shifted from the field of intraplate basaltic melts to the field of continental crust. Thus, the Karsakpai Group metabasalts are similar in trace-element composition to intraplate volcanics. According to isotope-geochemical data, there was no significant influence of contamination processes on the composition of the Burmash Formation basalts. The Bolbraun Formation basalts have geochemical features indicating their contamination with crustal material.

Conclusions

Fig. 8. Th/Yb–Nb/Yb composition diagram (Pearce, 2008) for the Karsakpai Group metabasaltoids. Designations follow Fig. 4. The MORB–OIB mantle sequence and compositions of the Mariana Arch rocks (Ma) are after Pearce (2008), and the average composition of continental crust (CC) is after Rudnick and Gao (2003).

Metavolcanics. Since Na, K, and other LILE (Rb, Sr, Ba) are the most mobile during postmagmatic transformations (Humphris and Thompson, 1978; Rudnick et al., 1985), we used only immobile HFSE and REE for the elucidation of the geodynamic nature of the rocks and for petrogenetic reconstructions. There is no correlation between the LREE contents and LOI (rLa-LOI = –0.13) in the Karsakpai Group basaltoids, whereas La and Th are in good correlation (Table 3). This indicates that the contents of LREE and HFSE, including Th, in these rocks are their primary characteristics and reflect the composition of their sources but are not related to secondary transformations (Turkina and Nozhkin, 2008). As seen from the primitive-mantle-normalized trace-element patterns (Fig. 5, Table 3), the metabasalts of the Burmash Formation are characterized by (Nb/La)PM and (Nb/Th)PM ≥ 1, which is usually considered (Hofmann, 1997; Sun and McDonough, 1989) to be an identification feature of basalts resulted from the mantle plume activity. On the Th/Yb–Nb/Yb (Pearce, 2008) diagram (Fig. 8), the composition points of these rocks fall in the field of intraplate basaltic melts. The sample with indicator ratios closest to the primitive-mantle ones ((Nb/La)PM = 1.0; (Nb/Th)PM = 1.3; (Th/La)PM = 0.8)) has εNd(1000) = +8.0 (Table 4), which is close to the composition of the depleted mantle of the same age (εNd(1000) = +7.9). Most likely, contamination processes did not seriously affect the composition of the Burmash Formation basalts. The Bolbraun Formation metabasalts are characterized by indicator ratios (Nb/Th)PM and (Nb/La)PM < 1 and (La/Sm)n > 1. The presence of a negative Nb anomaly and the Th and La enrichment of the rocks suggest the influence of crustal material on their composition. Crustal contamination is argued for by a clear negative correlation of (Nb/La)PM and Nb/Nb* with (La/Sm)n and correlation of εNd(1000) = +3.4 to –6.8 with (La/Sm)n and (Nb/Th)PM. On the Th/Yb–Nb/Yb diagram

The analysis of the isotope-geochemical parameters of the Karsakpai Group volcanosedimentary rocks (mafic volcanics, siliceous and ferruginous-siliceous sediments, and quartz– sericite–chlorite schists) has shown that the quartz–sericite– chlorite schists are the richest in Al2O3, TiO2, and K2O, i.e., components specific to clastic rocks. These rocks are also characterized by the highest ΣREE contents and elevated (relative to the quartzites) contents of other trace elements. Their (La/Yb)PAAS values are close to those of the Karsakpai Group basalts. We have established that the Karsakpai Group schists lack an exhalative material. They are, most likely, metamorphosed argillaceous rocks containing clastic material, mostly mafic. The Bolbraun Formation schists seem to contain additionally erosion products of more mature La- and Th-enriched felsic rocks. The isotope-geochemical data on the metaterrigenous rocks (TNd(DM) = 1.8–1.3 Ga, εNd(t) = –0.3 to +4.7) testify to the presence of Mesoproterozoic juvenile material in the provenances. The more ancient Nd model age (TNd(DM) = 1.8 Ga) and εNd(t) = –0.3 of the chlorite schist indicate the supply of sedimentary material from more ancient sources. The ferruginous-siliceous sediments of the Karsakpai Group show a strong positive correlation of K with Al and Ti, which suggests the same mechanism of their supply into the sediment, most likely, from a detrital source. The ferruginous quartzites have high Si/Al values typical of sediments formed mostly at the expense of a hydrothermal component. The Fe/Ti and Al/(Al + Fe + Mn) ratios suggest that these quartzites resulted from the mixing of hydrothermal and detrital components. Moreover, the detrital material was, most likely, basalts or volcanic sediments. The REE patterns show that the difference in the degree of Eu anomaly and the variations in the degree of LREE depletion and Y/Ho values are apparently due to the varying contribution of the detrital component. The ferruginous quartzites are characterized by positive εNd(t) values (+3.2 to +5.0). All this leads us to the conclusion that the formation of these rocks was related to hydrothermal activity and submarine volcanism, which resulted in sediments similar in 143Nd/144Nd values to the mantle rocks. The Karsakpai Group metabasaltoids correspond to high-Fe tholeiitic basalts. They are high–Ti low-/medium-K sodic rocks. The basaltoids are moderately differentiated. In trace-

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

element composition they are similar to intraplate volcanics. According to the isotope-geochemical data, contamination processes did not seriously affect the composition of the Burmash Formation basalts. The Bolbraun Formation metabasalts show indicator ratios (Nb/Th)PM, (Nb/La)PM < 1, and (La/Sm)n > 1, a clear negative correlation of (Nb/La)PM and Nb/Nb* with (La/Sm)n, and correlation of εNd(t) = +3.4 to –6.8 with (La/Sm)n and (Nb/Th)PM. All this testifies to the influence of crustal material on the composition of the volcanics. Thus, the intimate relationship between the ferruginous quartzites and intraplate volcanics indicates that they formed in a tectonically active basin. The Nd isotope composition of these quartzites was determined by synchronous submarine volcanism, whereas the 143Nd/144Nd values of the schists were additionally controlled by the Nd isotope composition of more ancient sources. The Mesoproterozoic Nd model ages and positive εNd(t) values of the metaterrigenous rocks of the Karsakpai Group indicate the presence of Mesoproterozoic juvenile material in the provenances. The minimum Nd model ages suggest the lower boundary of sedimentation of 1.3 Ga. This work was supported by research projects 0330-20160014, 0330-2016-0015, and 0330-2016-0013 (“Geodynamic and structure-lithologic evolution of the Central Asian Fold Belt and Siberian Platform: relationships and regularities of tectonic-process and magmatism manifestation”) from the Institute of Geology and Mineralogy, Novosibirsk, and by grants 15-20516 mol_a_ved and 15-35-20501 mol_a_ved from the Russian Foundation for Basic Research.

References Bau, M., Dulski, P., 1996. Distribution of yttrium and rare-earth elements in the Penge and Kuruman iron-formations, Transvaal Supergroup, South Africa. Precambrian Res. 79, 37–55. Bau, M., Dulski, P., 1999. Comparing yttrium and rare earths in hydrothermal fluids from the Mid-Atlantic Ridge: implications for Y and REE behaviour during near-vent mixing and for the Y/Ho ratio of Proterozoic seawater. Chem. Geol. 155 (1/2), 77–90. Bau, M., Koschinsky, A., Dulski, P., Hein, J.R., 1996. Comparison of partitioning behaviours of yttrium, rare earth elements, and titanium between hydrogenetic marine ferromanganese crusts and seawater. Geochim. Cosmochim. Acta, 60, 1709–1725. Bayanova, T.B., 2004. The Age of Geologic Reference Complexes on the Kola Peninsula and Magmatism Duration [in Russian]. Nauka, St. Petersburg. Bekker, A., Slack, J.F., Planavsky, A., Krapez, B., Hofmann, A., Konhauser, K.O., Rouxel, O.J., 2010. Iron formation: the sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric processes. Econ. Geol. 105, 467–508. Bostrom, K., 1970. Submarine volcanism as a source for iron. Earth Planet. Sci. Lett. 9 (4), 348–354. Bostrom, K., 1973. The origin and fate of ferromanganoan active ridge sediments. Stockholm Contrib. Geol. 27, 149–243. Bostrom, K., Peterson, M.N.A., Joensuu, O., Fisher, D.E., 1969. Aluminumpoor ferromanganoan sediments on active oceanic ridges. J. Geophys. Res. 74 (12), 3261–3270. Boynton, W.V., 1984. Geochemistry of the rare earth elements: meteorite studies, in: Henderson, P. (Ed.), Rare Earth Element Geochemistry. Elsevier, Amsterdam, pp. 63–114.

947

Cox, G.M., Halverson, G.P., Minarik, W.G., Le Heron, D.P., Macdonald, F.A., Bellefroid, E.J., Strauss, J.V., 2013. Neoproterozoic iron formation: an evaluation of its temporal, environmental and tectonic significance. Chem. Geol. 362, 232–249. Cox, G.M., Halverson, G.P., Poirier, A., Le Heron, D., Strauss, J.V., Stevenson, R.S., 2016. A model for Cryogenian iron formation. Earth Planet. Sci. Lett. 433, 280–292. Dmitrieva, N.V., Letnikova, E.F., Shkol’nik, S.I., Vishnevskaya, I.A., Kanygina, N.A., Nikolaeva, M.S., Sharf, I.V., 2016. Neoproterozoic metavolcanosedimentary rocks of the Bozdak Group in southern Ulutau (Central Kazakhstan): isotope-geochemical and geochronological data. Russian Geology and Geophysics (Geologiya i Geofizika) 57 (11), 1551–1569 (1969–1991). Douville, E., Bienvenu, P., Charlou, J.L., Donval, J.P., Fouquet, Y., Appriou, P., Gamo, T., 1999. Yttrium and rare earth elements in fluids from various deep-sea hydrothermal systems. Geochim. Cosmochim. Acta 63 (5), 627–643. Elderfield, H., Greaves, M.J., 1982. The rare earth elements in seawater. Nature 296, 214–219. Fryer, B.J., 1977. Rare earth evidence in iron-formations for changing Precambrian oxidation states. Geochim. Cosmochim. Acta 41, 361–367. Ewers, W.E., Morris, R.C., 1981. Studies on the Dales Gorge member of Brockman iron formation. Econ. Geol. 76, 1929–1953. Filatova, L.I., 1962. Precambrian strata in Ulutau, in: Materials on Geology of Central Kazakhstan [in Russian]. Izd. MGU, Moscow, Vol. 5. Filatova, L.I., 1983. Stratigraphy and Historical and Geological (Formation) Analysis of Precambrian Metamorphic Strata in Central Kazakhstan [in Russian]. Nedra, Moscow. Filatova, L.I., Rozanov, S.B., 1995. Volcanic complexes of the Early Proterozoic Kazakhstan greenstone belt. Vestnik Moskovskogo Universiteta. Ser. 4. Geologiya, No. 6, 11–30. Goldstein, S.J., Jacobsen, S.B., 1988. Nd and Sr isotopic systematics of river water suspended material: implications for crustal evolution. Earth Planet. Sci. Lett. 87, 249–265. Goldstein, S.L., O’Nions, R.K., Hamilton, P.J., 1984. A Sm–Nd isotopic study of atmospheric dusts and particulates from major river systems. Earth Planet. Sci. Lett. 70, 221–236. Gurvich, E.G., 1998. Metalliferous Sediments of the World Ocean [in Russian]. Nauchnyi Mir, Moscow. Henderson, P., 1984. General chemical properties and abundances of the rare earth elements, in: Henderson, P. (Ed.), Rare Earth Element Geochemistry. Elsevier, Amsterdam, pp. 1–32. Hofmann, A.W., 1997. Mantle geochemistry: the message from oceanic volcanism. Nature 385, 219–229. Humphris, S.E., Thompson, G., 1978. Hydrothermal alteration of oceanic basalts by seawater. Geochim. Cosmochim. Acta 42, 107–125. Jacobsen, S.B., Wasserburg, G.M., 1984. Sm–Nd isotopic evolution of chondrites, II. Earth Planet. Sci. Lett. 67, 137–150. James, H.L., Trendall, A.F., Morris, R.C., 1983. Chapter 12. Distribution of banded iron formation in space and time. Dev. Precambrian Geol. 6, 471–490. Jensen, L.S., 1976. A New Cation Plot for Classifying Subalkalic Volcanic Rocks. Ontario Division of Mines, Miscellaneous Paper 66. Klein, C., 2005. Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin. Am. Mineral. 90 (10), 1473–1499. Klein, C., Beukes, N.J., 1989. Geochemistry and sedimentology of a facies transition from limestone to iron-formation deposition in the Early Proterozoic Transvaal Supergroup, South Africa. Econ. Geol. 84 (7), 1733–1774. Klein, C., Beukes, N.J., 1993. Sedimentology and geochemistry of the glaciogenic Late Proterozoic Rapitan iron-formation in Canada. Econ. Geol. 88 (3), 542–565. Le Maitre, R.W., Bateman, P., Dudek, A., Keller, J., Lameyre, J., Le Bas, M.J., Sabine, P.A., Schmid, R., Sorensen, H., Streckeisen, A., Woolley, A.R., Zanettin, B., 1989. A Classification of Igneous Rocks and Glossary of Terms: Recommendations of the International Union of Geological Sciences Subcommission on the Systematics of Igneous Rocks. Blackwell Scientific Publication, Oxford.

948

N.V. Dmitrieva et al. / Russian Geology and Geophysics 58 (2017) 935–948

Lugmair, G.W., Carlson, R.W., 1978. Sm–Nd systematics of KREEP, in: Lunar and Planetary Science IX. Lunar and Planetary Institute, Houston, Part II, pp. 669–671. Michard, A., Albarède, F., 1986. The REE contents of some hydrothermal fluids. Chem. Geol. 55, 51–60. Michard, A., Albarède, F., Michard, G., Minster, J.F., Charlou, J.L., 1983. Rare-earth elements and uranium in high temperature solutions from East Pacific Rise hydrothermal vent field (13ºN). Nature 303, 795–797. Michard, A., Michard, G., Stüben, D., Stoffers, P., Cheminée, J.-L., Binard, N., 1993. Submarine thermal springs associated with young volcanoes: the Teahitia vents, Society Islands, Pacific Ocean. Geochim. Cosmochim. Acta 57, 4977–4986. Nikolaeva, I.V., Palesskii, S.V., Koz’menko, O.A., Anoshin, G.N., 2008. Analysis of geologic reference materials for REE and HFSE by inductively coupled plasma-mass spectrometry (ICP-MS). Geochem. Int. 46 (10), 1016–1022. Pearce, J.A., 2008. Geochemical fingerprinting of oceanic basalts with applications to ophiolite classification and the search for Archean oceanic crust. Lithos 100, 14–48. Pettijohn, F., Potter, P., Siever, R., 1972. Sand and Sandstone. Springer-Verlag, New York. Piepgras, D.J., Jacobsen, S.B., 1992. The behavior of rare earth elements in seawater: precise determination of variations in the North Pacific water column. Geochim. Cosmochim. Acta 56 (5), 1851–1862. Pin, C., Santos Zalduegui, J.F., 1997. Sequential separation of light-rare-earth elements, thorium and uranium by miniaturized extraction chromatography: Application to isotopic analyses of silicate rocks. Anal. Chim. Acta 339, 79–89. Polovinkina, Yu.I., 1952. Mafic and Ultramafic Rocks of Karsakpai in the Context of Genesis of Ferruginous Quartzites [in Russian]. Gosgeoltekhizdat, Moscow. Raczek, I., Jochum, K.P., Hofmann, A.W., 2003. Neodymium and strontium isotope data for USGS reference materials BCR-1, BCR-2, BHVO-1, BHVO-2, AGV-1, AGV-2, GSP-1, GSP-2 and eight MPI-DING reference glasses. Geostand. Geoanal. Res. 27, 173–79. Rozanov, S.B., 1976. Early Proterozoic jaspilite-bearing spilite association in Kazakhstan, in: Precambrian Geology and Tectonics of Central Kazakhstan. Materials on Geology of Central Kazakhstan [in Russian]. Izd. MGU, Moscow, Vol. 11, pp. 11–178.

Rudnick, R.L., Gao, S., 2003. Composition of continental crust, in: Holland, H.D., Turekian, K.K. (Eds.), Treatise on Geochemistry, Vol. 3: The Crust. Elsevier-Pergamon, Oxford, pp. 1–64. Rudnick, R.L., McLennan, S.M., Taylor, S.R., 1985. Large ion lithophile elements in rocks from high-pressure granulite facies terrains. Geochim. Cosmochim. Acta 49, 1645–1655. Satpaev, K.I., 1928. The Karsakpai district and its ore potential. Narodnoe Khozyaistvo Kazakhstana, No. 1, 92–109. Strakhov, N.M., 1947. Iron ore facies and their analogs in the Earth’s history, in: Transactions of the Geological Institute of the USSR Academy of Sciences [in Russian]. Izd. AN SSSR, Moscow, Issue 73. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematic of oceanic basalts: implications for mantle composition and processes, in: Saunders, A.D., Norry, M.J. (Eds.), Magmatism in the Ocean Basins. Special Publication 42. Geological Society of London, pp. 313–345. Tanaka, T., Togashi, Sh., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, T., Yuhara, M., Orihashi, Y., Yoneda, Sh., Shimizu, H., Kunimaru, T., Takahashi, K., Yanagi, T., Nakano, T., Fujimaki, H., Shinjo, R., Asahara, Y., Tanimizu, M., Dragusanu, C., 2000. JNdi-1: a neodymium isotopic reference in consistency with LaJolla neodymium. Chem. Geol. 168, 279–281. Taylor, S.R., McLennan, S.M., 1985. The continental Crust: Its composition and Evolution. Blackwel, Oxford. Turkina, O.M., Nozhkin, A.D., 2008. Oceanic and riftogenic metavolcanic associations of greenstone belts in the northwestern part of the Sharyzhalgai uplift, Baikal region. Petrology 16 (5), 468–491. Zaitsev, Yu.A., Filatova, L.I., 1971. New data on the structure of the Precambrian strata of Ulutau, in: Issues of Geology of Central Kazakhstan. Materials on Geology of Central Kazakhstan [in Russian]. Izd. MGU, Moscow, Vol. 10, pp. 21–92. Zaitsev, Yu.A., Rozanov, S.B., 1971. Structure of the Proterozoic greenstone and iron ore series of the Karsakpai synclinorium in southern Ulutau, in: Issues of Geology of Central Kazakhstan. Materials on Geology of Central Kazakhstan [in Russian]. Izd. MGU, Moscow, Vol. 10, pp. 107–122. Zhang, J., Nozaki, Y., 1996. Rare earth elements and yttrium in seawater: ICP-MS determinations in the East Caroline, Coral Sea, and South Fiji basins of the western South Pacific Ocean. Geochim. Cosmochim. Acta 60 (23), 4631–4644.

Editorial responsibility: A.E. Izokh