International Journal of Earth Sciences

, Volume 100, Issue 1, pp 45–62 | Cite as

Magmatic and metamorphic evolution of the Shotur Kuh metamorphic complex (Central Iran)

  • Mahmoud Rahmati-Ilkhchi
  • Shah Wali Faryad
  • František V. Holub
  • Jan Košler
  • Wolfgang Frank
Original Paper

Abstract

Metamorphic basement rocks, that are exposed beneath the very low-grade to unmetamorphosed Upper Jurassic-Eocene formations north of the Torud fault zone within the Great Kavir Block, were investigated to elucidate the origin of their protoliths and the pressure and temperature conditions of metamorphism. The basement, previously assumed as a pre-Cambrian metamorphic complex, is mostly formed by amphibolite-facies orthogneisses (tonalite, granodiorite, and granite) with amphibolites and small amounts of metasediment-micaschists. Major- and trace-element geochemistry in combination with U–Pb age dating of zircon showed that the protoliths formed during Late Neoproterozoic continental arc magmatism that has also been identified in other tectonic blocks of Central Iran. In addition to quartz, feldspar(s), micas in orthogneisses, and amphibole + plagioclase in amphibolite, all rocks may contain garnet that shows prograde zoning. Kyanite was found only in some Al-rich amphibolite together with gedrite. The PT conditions of the rocks, based on conventional geothermobarometry and the pseudosection method, show a medium-pressure amphibolite-facies metamorphism. Ar–Ar age dating of muscovite reveals that this metamorphism occurred in the Middle Jurassic (166 Ma) and related to the closure of the Neotethyan basin.

Keywords

Late Neoproterozoic arc magmatism Jurassic Barrovian type metamorphism Central Iran 

Introduction

The Central Iranian Terrane is a composite of fold-and-thrust belts with basement blocks covered by thick sedimentary sequences. Comparative stratigraphy and tectonics, combined with geochemistry of igneous rocks in Central Iran, led to the identification of several crustal blocks (e.g., Stöcklin 1968, 1977; Takin 1972; Crawford 1977; Berberian and Berberian 1981). Ramezani and Tucker (2003) summarised tectonic and geochronological data from Central Iran and focused on igneous rocks in the Saghand area that previously had been thought to represent the Precambrian basement. The igneous/metaigneous rocks occur along a north–south trending belt, the Kashmar-Kerman Tectonic Zone (KKTZ), which separates the Yazd block to the west and the Tabas block to the east in Central Iran (Fig. 1a). Based on U–Pb ages and geochemical data, Ramezani and Tucker (2003) postulated that these rocks are remnants of a Late Neoproterozoic arc that developed along the Proto-Tethyan margin of the Gondwanaland supercontinent. Metamorphism of these rocks that reached amphibolite-facies condition occurred in Eocene times. However, several basement rocks exposed beneath the Paleozoic-to-Mesozoic sedimentary sequences west of the KKTZ are still assumed to be Pre-Cambrian metamorphic basement, even though detailed information about their composition and metamorphic evolution is lacking. Bagheri and Stampfli (2008) investigated amphibolite-facies metaigneous and metasedimentary rocks exposed along the boundary between the Yazd and Great Kavir blocks and postulated a Paleozoic accretionary complex with a 320 Ma of greenschist-to-amphibolite facies metamorphism and a Permian–Triassic (280–230 Ma) high-pressure metamorphism. It seems that boundaries between different crustal blocks were significant in the formation of igneous and metamorphic rocks during the geological evolution of Central Iran.
Fig. 1

a Simplified tectonic map of eastern Iran showing constituent crustal blocks (compiled from Ramezani and Tucker 2003; Alireza Nadimi 2006). KKTZ Kashmar-Kerman Tectonic Zone, DBF Dehshir Fault, TFZ. b Schematic map of the Torud area with important faults (after Hushmandzadeh et al. 1978). c Simplified geological map of the Rezveh area (Rahmati-Ilkhchi 2002)

In this work, we focus on the Shotur Kuh metamorphic complex in the Great Kavir Block (GKB) (Fig. 1a). It forms a tectonic window beneath the Mesozoic and Cenozoic sediments near the Torud fault zone. Similar to other crystalline windows exposed in Central Iran, it has been assumed to represent a Pre-Cambrian metamorphic basement. In addition to geochemistry and metamorphic evolution, we present new U–Pb zircon and Ar/Ar mica ages. The implications of our results for the metamorphic and tectonic evolution of the GKB along the Torud tectonic zone, as well as relationships to other igneous and metamorphic complexes in Central Iran, are discussed.

Geological setting

From east to west, the Central Iranian Terrane consists of four major crustal domains (Fig. 1a): the Lut Block, the Tabas Block, the Yazd Block, and the Great Kavir Block (see for e.g., Berberian and Berberian 1981; Bagheri and Stampfli 2008). These blocks are separated by a series of intersecting regional-scale faults. Although the stratified cover rocks are largely comparable among different blocks, locally significant facies and/or thickness variations occur across the domain boundaries. Each block features a particular overall deformation style and pattern of recent seismicity, distinguishable from those in the adjacent domains (Berberian and Berberian 1981). The Tabas and Yazd blocks are separated by a nearly 600-km long, arcuate, and structurally complex belt, the KKTZ, composed of variably deformed and fault-bound supracrustal rocks.

The Rezveh area with the Shotur Kuh metamorphic complex (SKMC) is a part of the Great Kavir Block (GKB) and occurs north of the Torud fault zone (Fig. 1b). This tectonic zone is followed by a series of Eocene volcanic rocks of dacitic and andesitic composition. The SKMC represents an E–W trending elliptical tectonic window (c. 20-km long and 11-km wide) that, together with its Pre-middle Triassic tectonic cover, is exposed beneath the Jurassic-Miocene sedimentary sequences (Fig. 1c). Most of the basement rocks have igneous precursors, and only small amounts of micaschists and phyllites, which represent the lower part of the Pre-middle Triassic (Permian-Lower Triassic?) sequence with metamorphosed limestone and dolomite are present. These rocks occur along the southern and northern boundaries of the SKMC. The main rocks are orthogneisses of tonalitic, granodioritic, and granitic compositions with various amounts of amphibolites (probably former dykes) that have lengths up to several tens of metres and widths from centimetres to 20 m. The amount and thickness of the amphibolite bodies decreases from meta-granodiorite to granite, and they are most common in the central and western part of the complex. In some cases, amphibolite and orthogneiss may form parallel stripes and bands that are locally deformed into isoclinal folds. The metagranite occurs along the northern border and in the central part of this complex.

Metapelitic rocks (black shale-phyllite-garnet micaschist) rimming the southern border of the complex form the lower part of the Pre-middle Triassic (Permian-Lower Triassic?) formation. They intercalate with calcitic and dolomitic marbles that become a massif carbonatic sequence upward with few or no pelitic rocks. Marbles and garnet micaschists are isoclinally folded together with metaigneous rocks and are tectonically juxtaposed with phyllites, which show a lower degree of metamorphism. The Pre-middle Triassic pelitic–carbonatic rocks are covered by the Upper Triassic-Lower Jurassic Shemshak formation, which is represented by the intercalation of psammopelitic rocks with conglomerates. They show metamorphic fabrics similar to those in the Pre-middle Triassic rocks, but in lower greenschist facies conditions that are characterised by the presence of chlorite, fine-grained white mica and quartz. The uppermost (Middle Jurassic) part of the Mesozoic sequence, beneath the Neogene sediments, is characterised by very low-grade sandstones, shales and conglomerates, which contain pebbles of the basement rocks (Rahmati-Ilkhchi 2002).

Analytical methods

We studied the rocks using bulk-rock geochemistry, mineral composition, and age dating of zircon and micas. Powdered whole-rock samples were analysed in the Activation Laboratories (Actlabs), Ltd. (Ancaster, Ontario) using the lithium metaborate/tetraborate fusion technique followed by a combination of (1) inductively coupled plasma emission spectroscopy (ICP-OES) for major elements plus Sc, Be, and V; and (2) inductively coupled plasma mass spectrometry (ICP-MS) for all other trace elements. The analytical results (see Table 1) are illustrated in geochemical diagrams using the GCDkit software (Janoušek et al. 2006).
Table 1

Representative whole-rock chemical analyses of the Shotur Kuh orthogneisses and amphibolites

Rock

Orthogneiss

Amphibolite

I-10

P8

P160

P109

P122

P203

P147

P183

P108

P202

P182

I-5

P179

P29

P68

P74

SiO2

63.72

64.29

71.74

72.21

73.66

73.72

74.54

75.16

75.91

76.89

46.21

49.58

51.12

51.57

49.91

51.6

TiO2

0.57

0.73

0.12

0.44

0.24

0.38

0.19

0.18

0.12

0.06

0.69

0.72

1.65

2.12

2.26

1.40

Al2O3

13.42

16.30

12.11

13.23

13.36

12.71

13.06

13.02

12.27

12.22

18.36

16.35

15.35

14.91

13.73

17.46

Fe2O3 tot.

10.42

5.88

1.67

3.77

2.57

2.80

2.02

1.84

1.54

0.90

8.88

8.63

11.98

13.25

13.41

13.45

MnO

0.18

0.05

0.01

0.04

0.04

0.03

0.01

0.02

0.04

0.01

0.12

0.12

0.18

0.20

0.21

0.22

MgO

1.74

3.06

0.40

0.81

0.58

0.56

0.39

0.36

0.27

0.13

12.19

8.80

6.95

5.75

5.32

2.49

CaO

5.68

2.22

2.81

3.13

1.63

1.43

0.91

1.56

1.15

0.44

8.65

11.47

9.59

9.29

9.19

8.29

Na2O

3.21

3.88

3.19

3.39

3.36

2.89

3.43

3.48

2.78

3.21

2.80

2.68

2.27

1.63

2.65

2.44

K2O

0.47

2.47

3.35

1.33

3.23

4.61

4.30

3.86

4.08

4.59

0.20

0.53

0.31

0.43

1.69

1.33

P2O5

0.06

0.22

0.59

0.12

0.08

0.11

0.23

0.05

0.05

0.06

0.09

0.07

0.23

0.30

0.34

0.45

LOI

0.30

1.09

2.86

0.89

1.09

0.56

1.20

0.28

0.79

0.68

1.62

0.90

1.20

1.47

1.11

0.61

Total

99.20

100.20

98.86

99.36

99.83

99.81

100.30

99.81

98.98

99.19

99.63

99.85

100.80

100.90

99.82

99.74

mg

24.9

50.8

32.2

29.9

30.9

28.4

27.7

27.9

25.8

22.3

73.1

66.9

53.5

46.2

44.0

26.8

A/CNK

0.83

1.24

0.87

1.04

1.11

1.03

1.09

1.02

1.11

1.11

0.89

0.63

0.71

0.74

0.60

0.85

MALI

−2.0

4.1

3.7

1.6

5.0

6.1

6.8

5.8

5.7

7.4

−5.6

−8.3

−7.0

−7.2

−4.9

−4.5

Rb

5

100

86

49

104

215

107

128

131

267

8

14

6

18

54

50

Sr

287

247

139

240

110

65

82

73

104

26

207

166

142

206

154

283

Ba

130

603

598

792

720

332

706

305

526

77

36

83.2

49

103

126

389

Y

23.1

35.4

33.0

14.2

23.6

59.6

32.8

46.2

21.7

52.0

19.0

17.5

42.3

41.0

46.1

41.3

Zr

609

220

132

215

161

226

127

139

101

64

66

48

127

198

223

209

Nb

11.4

16.4

10.4

7.7

9.3

11.3

6.5

8.8

5.8

6.7

3.0

1.7

7.1

19.5

16.4

14.4

Hf

15.2

5.3

3.7

4.7

4.1

6.4

4.0

4.8

3.1

3.0

1.5

1.5

3.3

4.4

5.9

4.8

Ta

0.6

1.3

0.9

0.4

0.8

1.2

0.9

0.9

1.0

1.8

0.2

0.1

0.4

1.1

1.2

0.8

Th

6.4

12.9

14.8

9.6

14.8

34.9

14.3

17.9

16.1

14.5

0.4

0.7

1.7

2.4

3.4

3.2

La

29.3

30.0

33.6

40.5

30.9

49.1

28.6

22.8

16.2

11.3

4.0

3.6

10.4

17.8

18.2

37.2

Ce

59.2

63.8

68.7

74.5

60.7

105

57.9

49.4

32.8

28.1

10.8

9.6

24.6

40.8

42.7

89.9

Sm

7.00

5.73

5.78

4.07

3.97

8.67

5.54

6.40

2.72

4.23

1.84

2.30

4.43

5.68

7.02

7.75

Eu

1.85

1.19

0.63

1.29

0.61

0.93

0.86

0.44

0.49

0.19

0.80

0.94

1.65

2.00

2.31

2.01

Gd

5.37

5.92

5.79

3.53

3.78

8.14

4.87

6.77

2.89

4.52

2.58

2.73

6.21

7.08

7.28

7.9

Yb

2.39

2.76

2.98

1.13

2.33

5.30

2.95

4.86

2.33

5.66

1.68

1.47

3.59

3.30

4.04

3.46

Major oxides are in weight percent, trace elements in parts per million (ppm)

Atomic ratio mg = 100 Mg/(Mg + Fetot.); molar ratio (alumina saturation index) A/CNK = Al2O3/(CaO + Na2O + K2O); Modified alkali-lime index (MALI) = Na2O + K2O – CaO (from wt%, Frost et al. 2001)

Mineral chemical analyses were carried out with a CAMECA SX 50 electron microprobe at the Institute of Mineralogy, Technische Universität Stuttgart, which is equipped with four wavelength-dispersive spectrometers. The following synthetic standards were used: pyrope (Si, Al, and Mg), andradite (Ca, Fe), jadeite (Na), spessartine (Mn), K-silicate glass (K), Ba-silicate glass (Ba), and NaCl (Cl), as well as natural rutile (Ti) and topaz (F). The operating voltage was 15 kV using beam currents between 10 and 15 nA. The beam was focused to 1–2 μm diameter, except for micas, for which an 8–10 μm beam was used. The peak counting time was 20 s. Some compositional profiles of garnet were obtained using the scanning electron microscope JEOL 6310 at the Institute for Petrology and Structural Geology, Charles University, Prague. Representative mineral analyses are given in Table 2.
Table 2

Selected microprobe analyses of minerals used for PT calculations

Sample

Garnet

Biotite

Chlorite

Amphibole

Plagioclase

Rock

Orthogneiss

Mica

Amphibolite

Orthogneiss

Mica

Amphibolite

Orthogneiss

Orthogneiss

Amphibolite

I-12

P95

P99

P34

P200

I-10a

I-10b

I-12

P95

P99

P200

I-10a

I-10b

P34

I-12

P95

P99

P34

I-10a

I-10b

c

r

SiO2

37.37

37.97

36.67

37.89

37.54

36.72

38.49

38.19

36.58

37.19

36.70

24.11

41.43

42.39

39.50

63.71

62.87

60.65

62.3

61.81

63.42

TiO2

0.07

0.04

0.03

0.00

0.11

0.03

0.12

0.00

2.03

2.71

2.43

0.06

0.72

0.73

1.15

      

Al2O3

21.10

21.38

20.48

20.85

20.97

20.92

20.75

21.38

17.63

19.10

17.79

21.69

14.68

13.50

15.22

22.82

23.05

24.85

23.68

24.59

22.98

Fe2O3

0.75

0.00

0.94

0.73

0.00

0.00

2.16

0.03

    

2.38

0.00

 

0.01

0.14

0.00

0.04

0.12

0.00

FeO

32.84

32.35

32.82

31.89

28.15

32.15

21.48

26.73

17.87

16.37

18.12

25.79

14.38

21.03

17.62

0.00

0.00

0.00

0.00

0.00

0.00

MnO

2.69

2.25

1.69

1.54

1.54

0.76

3.84

0.47

0.00

0.04

0.00

0.16

0.16

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

MgO

3.84

3.89

3.81

3.86

2.63

2.48

2.10

1.31

11.61

12.03

11.16

15.11

8.65

5.51

8.24

0.00

0.00

0.00

0.00

0.00

0.00

CaO

2.21

2.61

2.58

3.27

8.61

6.36

11.41

12.22

0.02

0.34

0.00

0.06

11.34

11.09

11.30

3.94

4.93

6.44

5.04

6.37

4.22

Na2O

0.07

0.00

0.01

0.07

0.11

0.01

0.43

0.00

0.11

0.00

0.12

0.03

1.79

1.78

2.26

9.35

8.74

7.86

8.82

7.66

9.31

K2O

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

9.75

9.10

9.73

0.03

1.32

1.13

0.70

0.11

0.16

0.18

0.00

0.14

0.00

Total

100.90

100.50

99.03

100.10

99.75

99.43

100.80

100.30

95.59

96.88

96.05

87.05

96.85

97.17

96.00

99.94

99.89

99.98

99.88

100.69

99.93

Si

2.942

3.022

2.948

3.032

2.993

2.967

3.023

3.029

2.755

2.726

2.754

2.209

6.254

6.480

6.081

2.811

2.785

2.700

2.763

2.723

2.805

Ti

0.004

0.003

0.002

0.000

0.007

0.002

0.007

0.000

0.115

0.149

0.137

0.004

0.082

0.084

0.133

0.00

0.00

0.000

0.000

0.00

0.000

Al

1.958

2.006

1.940

1.966

1.959

1.993

1.921

1.998

1.565

1.650

1.573

2.342

2.612

2.433

2.762

1.186

1.204

1.304

1.238

1.277

1.198

Fe3+

0.044

0.000

0.057

0.044

0.074

 

0.128

0.002

    

0.270

0.059

0.000

0.000

0.005

0.000

0.000

0.004

0.000

Fe2+

2.162

2.009

2.206

2.134

1.792

2.172

1.411

1.774

1.126

1.004

1.137

1.976

1.815

2.688

2.269

0.00

0.00

0.000

0.001

0.00

0.000

Mn

0.179

2.154

0.115

0.105

0.103

0.052

0.255

0.032

0.000

0.002

0.000

0.012

0.020

0.000

0.000

0.00

0.00

0.000

0.000

0.00

0.000

Mg

0.451

0.222

0.457

0.460

0.311

0.551

0.246

0.156

1.303

1.315

1.248

2.064

1.947

1.256

1.892

0.00

0.00

0.000

0.000

0.00

0.000

Ca

0.186

0.461

0.222

0.280

0.731

0.299

0.960

1.039

0.002

0.027

0.000

0.006

1.834

1.816

1.864

0.186

0.234

0.307

0.239

0.301

0.200

Na

0.011

0.152

0.001

0.010

0.017

 

0.065

0.000

0.016

0.000

0.019

0.000

0.524

0.527

0.674

0.800

0.751

0.679

0.758

0.654

0.798

K

0.000

0.000

0.000

0.000

0.000

0.000

0.000

0.000

0.937

0.851

0.932

0.003

0.254

0.220

0.137

0.006

0.009

0.010

0.000

0.008

0.000

alm

0.726

0.414

0.735

0.714

0.610

0.707

0.491

0.591

      

An

0.188

0.236

0.308

0.240

0.312

0.200

py

0.151

0.046

0.152

0.154

0.105

0.179

0.086

0.052

      

Ab

0.806

0.755

0.681

0.760

0.679

0.800

grs

0.063

0.095

0.074

0.094

0.248

0.097

0.334

0.346

      

Or

0.006

0.009

0.011

0.000

0.008

0.000

sps

0.060

0.444

0.038

0.035

0.035

0.017

0.089

0.011

             

XMg

0.17

0.10

0.172

0.177

0.15

0.20

0.15

0.08

0.54

0.57

0.52

0.51

0.52

0.32

       

c core, r rim

aMicaschist

U–Pb dating of zircons was performed by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) using a New Wave UP-213 (213 nm, Nd:YAG) laser system coupled to a Thermo Finnigan Element 2 ICP-MS instrument at the Department of Earth Science, University of Bergen. The analytical technique and data reduction were modified from Košler et al. (2002). Ar–Ar age dating was obtained by measuring the 40Ar*/39Ar isotopic ratio using the mass spectrometry VG 5400 at the Central European Ar-Laboratory in Bratislava.

Petrography

The main rock groups of the SKMC are orthogneisses that are derived from rocks of granitic, granodioritic, and tonalitic compositions. The metatonalite are characterised by the presence of amphibolite bodies, and they are exposed in the central and eastern parts of the complex. Phyllites to micaschists studied here come from the southern border of the SKMC.

Orthogneisses

The rocks are medium- to coarse-grained and (partly) display augen textures with various degrees of foliation. At least three varieties of orthogneisses can be distinguished based on the present mineral assemblages: amphibole + plagioclase + biotite + quartz ± garnet (metatonalite); plagioclase + biotite + quartz ± garnet; and plagioclase + biotite + muscovite ± garnet (metagranodiorite to metagranite). Grain size and augen structure in the orthogneisses increase from tonalites to granites. In addition to feldspars, quartz and biotite, some tonalitic and granitic varieties may additionally contain amphibole and muscovite, respectively. Zircon, apatite, rutile, titanite, allanite, and monazite are accessory phases. Metagranites are strongly foliated and they are rich in muscovite (up to 10 vol%); together with biotite and thin quartz bands, they define the foliation (Fig. 2a). Such rocks occur along the northeastern border and also in the central part of this complex. Garnet is not found often but occurs in all compositional varieties of the orthogneisses. The relict K-feldspar is slightly perthitic, but microcline and myrmekite can be also observed. An interesting feature in metatonalites and metagranodiorites is the presence of dark-brown allanite, which may form up to 1-mm long crystals oriented parallel to the foliation. It is usually rimmed by epidote.
Fig. 2

Microphotos of an muscovite-bearing orthogneiss metagranite (a) and a micaschist–phyllite (d) from the Shotur Kuh metamorphic complex. b and c are backscattered images of gedrite- and kyanite-bearing amphibolite

The rocks show various degrees of retrogression and recrystallisation, which is characterised by the presence of several mineral phases. Large plagioclase crystals contain small grains of zoisite–clinozoisite, and also, though rarely, of calcite in the cores. Both feldspars can be partly replaced by fine-grained white mica. Garnet is usually cracked, and along cracks and rims it is partially replaced by fine-grained biotite and chlorite. The fine-grained variety of biotite is also present in some mylonitized orthogneisses. Compared with the coarse-grained, red-brown biotite, which shows textural equilibrium with garnet, the fine-grained variety has light-brown and green colours.

Amphibolites

Amphibolites are mostly garnet-free, and the major phases are plagioclase and amphibole, which define the foliation of the rock. Porphyroblastic garnet (up to 5 mm in size) and quartz can each be present up to 15 vol%. Accessory phases are zircon, apatite, titanite, and opaque minerals, secondary minerals involve epidote and chlorite which replace amphibole. Relict igneous clinopyroxene was observed in one sample (I-10a) and it forms large crystals (1–3 mm in size) that are mostly replaced by actinolite and epidote. Plagioclase crystals (1–2 mm in size) are partly replaced by sericite.

In one case (sample P182), kyanite and gedrite were found (Fig. 2b, c). This rock is coarse-grained and well foliated, and consists of calcic amphibole, plagioclase, gedrite, and accessory rutile, kyanite, biotite, and phengite. Gedrite, forming up to 7 mm long crystals, show a weak pleochroism from pale to yellow grey, forms intergrowths with calcic amphibole and contains inclusions of plagioclase. Phengite was found in kyanite, which together with partial replacement of amphibole by chlorite along the cleavage indicates very weak retrogression.

Micaschists

The fine-grained rocks grading from phyllites to micaschists come from the southern border of the SKMC. The micaschists are strongly foliated and consist of quartz, white mica, garnet, biotite, and small amounts of plagioclase. In contrast to muscovite-bearing orthogneisses they contain a very fine graphitic substance that is indicative of a sedimentary protolith. Garnet crystals (up to 0.5 mm in size) form idioblasts, with pressure shadows consisting of quartz and large flakes of biotite and muscovite (Fig. 2d). The rocks are retrogressed to various degrees, with garnet grains being replaced by calcite and chlorite, biotite being replaced by chlorite, and plagioclase containing numerous flakes of fine-grained white mica.

Geochemistry

To characterise the tectono-magmatic evolution of the rocks, 16 samples (10 orthogneisses and 6 amphibolites) were selected for whole-rock analyses. Representative data are listed in Table 1. Since we did not recognise any particular evidence for major-element transport during post-magmatic evolution and regional metamorphism, we assumed that the present compositional variations of the orthogneisses and amphibolites may largely reflect the original characteristics of the igneous rocks. No chemical element, however, may be considered as fully “immobile”, and therefore all of the geochemical conclusions should be regarded with caution.

Orthogneisses

Orthogneisses represent a heterogeneous group of subalkaline rocks with SiO2 ranging from 63.7 to 76.9 wt% and variable amounts of alkalis (Na2O 2.8 to 3.9 wt%, K2O 0.5 to 4.6 wt%). Based on normative composition (CIPW), they correspond to tonalite, granodiorite, and monzogranite, with small amounts of syenogranite and alkali-feldspar granite.

Except from two samples of metaluminous tonalite (sample I-10) and granite (P160) and one strongly peraluminous granodiorite (P8) with A/CNK = 1.245, all other orthogneisses have weak-to-moderate peraluminous composition with A/CNK = 1.01–1.12. Almost all samples correspond to calc-alkaline rocks. The only exception is the tonalite gneiss (I-10) with unusually high Fetot (10.42 wt% Fe2O3) and low K2O (0.5 wt%).

Some geochemical characteristics of the orthogneisses are shown on a spiderdiagram (Fig. 3) and on two element–element variation diagrams proposed by Pearce et al. (1984) for discrimination of the tectonomagmatic position of granitic rocks (Fig. 4). Most orthogneisses have high large-ion lithophile elements (LILE) and Th (Fig. 3) and display high LILE/HFSE elemental ratios. Only the tonalitic gneiss (I-10) has low LILE (K, Rb, and Ba) but significantly higher Zr and Hf contents, which are close to those in “ocean-ridge granites” (ORGs) (Fig. 3). Geochemical features of this sample are generally similar to some dacitic volcanic rocks from the Cambrian Volcano-Sedimentary Unit (Lower Cambrian dacite porphyry JR95-G47from Zarigan Mountain; Ramezani and Tucker 2003), however, it is more evolved and thus has higher SiO2 and lower CaO.
Fig. 3

Spidergram of selected elemental abundances in orthogneisses normalised to the “ocean-ridge granite” (ORG). Normalising values are from Pearce et al. (1984)

Fig. 4

Position of the studied orthogneisses in selected discrimination diagrams for the tectonic setting of granitoid rocks according to Pearce et al. (1984)

Amphibolites

Mafic rocks associated with orthogneisses are generally low in SiO2 (46.2 to 51.6 wt%), high in total Fe as Fe2O3 (8.6–13.5 wt%) and have variable contents of MgO (12.2–2.5 wt%), Na2O (1.6–2.8 wt%), and K2O (0.2–1.7 wt%). They correspond to subalkaline basalts or gabbros of tholeiitic affinities. The high MgO and mg-value of sample P182 suggests a primitive composition. However, this rock contains gedrite and kyanite and is rich in Al2O3 (18.36 wt%). Samples P182 and I-5 have low TiO2, P2O5 and incompatible elements including rare earth elements (Fig. 5) that could result from accumulation of early magmatic phases (calcic plagioclase, olivine, and orthopyroxene). In contrast, sample P74 shows an evolved basaltic composition with high total Fe but with a very low mg-number and increased contents of incompatible elements. Geochemical features of amphibolite samples are illustrated in a spiderdiagram (Fig. 5) and selected widely used tectonic discrimination plots (Fig. 6). They show high Th/Ta ratios and plot both in different fields of within plate, MORB, and volcanic arc basalts. Their geotectonic interpretation is discussed later.
Fig. 5

Spidergram of MORB-normalised abundances of selected trace elements after Pearce (1983) for amphibolites associated with orthogneisses from the Shotur Kuh metamorphic complex

Fig. 6

Selected discrimination plots for the tectonic setting of analysed amphibolites. a Zr/4–2 × Nb–Y (Meschede 1986); b Th–Hf/3–Ta (Wood 1980). AI within-plate alkali basalts, AII within-plate alkali basalts and within-plate tholeiites; B E-type MORB; C within-plate tholeiites and volcanic-arc basalts; D N-type MORB and volcanic-arc basalts

Mineral chemistry

Orthogneisses

Garnet

Garnet from muscovite-free varieties is mostly rich in Fe and Ca (Alm59–66, Grs16–18, Py15–18, and Sps1–7). The amphibole- or epidote-bearing varieties have garnets with relatively low Mg and higher Ca contents (Fig. 7). The most Ca-rich garnet comes from the orthogneiss containing allanite rimmed by epidote (sample I-8b). Most garnets show weak but clear prograde zoning characterised by the decrease of Mn and XFe = Fe2+/(Mg + Fe2+) from the core towards the rim. In some cases, garnet may show a retrograde zoning (an increase of Mn and XFe) at the outermost rim part of the grains (Fig. 8). The variation of Ca differs from sample to sample but mostly shows an increase from core to rim. Garnet in the muscovite-bearing variety, including that with microcrystalline texture, has relatively high Fe (Alm69–73) and Mg (Py14–20) and low Ca (Grs5–9) abundances. Similar to garnet from Ms-free rocks, slight zoning from core to rim can be observed in some grains. The spessartine content ranges between 3 and 8 mol%.
Fig. 7

Composition of garnet from orthogneisses, amphibolites, and micaschists in the Shotur Kuh complex. m and am are muscovite- and amphibole-bearing varieties, respectively

Fig. 8

Compositional profiles of garnet from muscovite-bearing orthogneiss (a sample P99) and from micaschist (b sample P200)

Micas

Biotite from muscovite-free samples has lower XMg = 0.26–0.36 compared with that in the muscovite-bearing variety (XMg = 0.51–0.55), but the two varieties have similar Al (1.6–1.7 atoms per formula unit (a.p.f.u)) and Ti (0.1–0.2 a.p.f.u.) contents. Biotite associated with chlorite, muscovite, and albite in mylonitized orthogneiss has XMg = 0.32–0.34, which is similar to that in the muscovite-free orthogneisses, but with lower Al (1.4 a.p.f.u.) content. Muscovite has low Si (3.0–3.1 a.p.f.u.) with relatively high paragonite content (10–20 mol%), and only the microcrystalline variety of orthogneiss has low (3 mol%) paragonite content.

Plagioclase

Plagioclase in the muscovite-free varieties has a high anorthite content (An22–31 mol%), and some higher An contents come from the cores of large plagioclase grains. The coarse-grained, muscovite-free variety has An18–24, while that from the microcrystalline variety has relatively high Ca content (An23–30).

Other minerals

Epidote was analysed in muscovite-free orthogneisses, and it has XAl (Al/(Al + Fe3+)) = 0.75–0.86.

Amphibolite

Garnet

Garnet from amphibolite is mostly almandine-grossular solid solution. Its composition varies from sample to sample between Alm49–64, Grs31–34, Py3–9, and Sp1–9. Garnet from biotite-amphibole gneiss (sample P34) has higher Mg and lower Ca (Alm56–63, Grs21–25, Py10–13, and Sps2–9) and it is mostly homogeneous, although slight zoning characterised by an increase in Mg and Fe and a decrease in Mn and Ca towards the rim can be seen.

Amphiboles

Amphibole in kyanite-free amphibolites is mostly pargasite and ferropargasite, with Si = 6.2–6.4 a.p.f.u. and XMg = 0.53–0.30. Ferropargasite is present in sample I-10a, along with relatively high-Fe garnet (Fig. 7). The A-site of amphibole is occupied by 0.50–0.65 atoms, and NaM4 position is in the range of 0.25–0.45 a.p.f.u. Amphibole from biotite-bearing amphibolite has XMg = 0.38–0.45 with NaM4 about 0.2, and Mg-rich tschermakite with XMg = 0.9 (NaM4 = 0.40–0.44 a.p.f.u., Si = 6.3 a.p.f.u.) is associated with gedrite in kyanite-bearing amphibolite (sample P182). The coexisting gedrite, which has XMg = 0.7, is rich in Al (Al2O3 = 17.3 wt%) and Na (Na2O = 2.3 wt%), and the A-site occupancy is about 0.4 a.p.f.u. Considering the coupled substitution of tschermakite (Ts): AlVI + AlIV = MgVI + SiIV and edenite (Ed): NaA + AlIV = A + SiIV (Robinson et al. 1971), the analysed orthoamphibole shows Ed/Ts = 0.33 and has a composition close to ideal gedrite.

The analysed gedrite in the kyanite-bearing amphibolite has Al = 2.9 a.p.f.u. and Na = 0.55 a.p.f.u., which are slightly lower compared with typical literature data (Spear 1980). Robinson et al. (1971) emphasised the importance of Na in the A-site of gedrite and postulated an ideal end-member gedrite composition of Na0.5(Mg,Fe2+)2(Mg,Fe2+)3.5Al1.5Si6Al2O22(OH)2. This formula represents a combination of the edenite and tschermak’s substitutions in a ratio of one to three. The analysed gedrite has lower A-site occupancies and AlVI compared with ideal gedrite, with A = 0.5 and AlVI = 2.0 (Spear 1980).

Plagioclase

Plagioclase composition in amphibolite is An30–31 in sample I-10a and An18–20 in sample I-10b. Oligoclase–andesine with 28–30 mol% of An occurs in the biotite-bearing amphibolite, and more Ca-rich andesine with An33–35 is present in the kyanite-bearing amphibolite (sample P182).

Micas

Biotite from the kyanite-bearing amphibolite (P182) has XMg = 0.7. The phengite analysed in this sample has a high Si content (Si = 3.36–3.40 a.p.f.u).

Micaschist

Garnet (sample P200) is rich in almandine and shows strong zoning (Fig. 8) with a Mn-rich core (Alm49,Grs16,Py3,Sps31), a progressive increase in Mg and a decrease in Mn and XFe (Alm70,Grs18,Py10,Sps1) towards the rim. Biotite has XMg = 0.52–0.55, the silica and paragonite contents in muscovite correspond to 3.0 a.p.f.u. and 11–16 mol%, respectively. Plagioclase is pure albite.

PT conditions of metamorphism

Most of the studied rocks have simple mineral assemblages containing feldspars, biotite ± garnet in orthogneisses, and amphibole ± garnet in amphibolite. Plagioclase is not always totally equilibrated and may have preserved its igneous composition in the cores of large porphyroclasts. In one case, relics of clinopyroxene were also observed in amphibolite. To constrain the PT conditions of metamorphism, well-foliated and recrystallised samples of orthogneisses and amphibolites were selected. Temperature was calculated using the exchange thermometers of garnet–biotite (Ganguly et al. 1996; Holdaway 2000), garnet–muscovite (Hynes and Forest 1988), and amphibole–garnet (Graham and Powell 1984; Perchuk and Lavrentjeva 1983; Dale et al. 2000). Because of the lack of Al–silicate in garnet-bearing gneisses, pressure was estimated using the garnet–biotite–plagioclase–quartz barometer of Wu and Cheng (2006) and the plagioclase–garnet–amphibole–quartz of Kohn and Spear (1990) in combination with the occurrence of kyanite in amphibolite. Possible plausibility of the calculated PT conditions was tested by the pseudosection method, using the Perplex 07 software (Connolly 2005: version 07) with the internally consistent thermodynamic dataset of Holland and Powell (1998, 2003) for the muscovite-bearing orthogneiss (P99). The solution models of minerals that were used are plagioclase (Newton et al. 1980) biotite, chlorite, and garnet (Holland and Powell 2003).

Three orthogneiss samples (I-12, P95, and P99) and one micaschist sample (P200) were selected for garnet–biotite thermometry, and the temperatures obtained using three calibrations for the orthogneisses are in the range of 577–645°C (Table 3) but show a slight increase from the calibration of Ganguly et al. (1996) through Holdaway (2000) to Wu and Cheng (2006). Relatively higher temperatures (615–695°C) were obtained for the garnet-bearing amphibolite and amphibole–garnet-bearing orthogneiss using garnet–amphibole thermometry with the calibration of Graham and Powell (1984) and Dale et al. (2000). The calibration of Perchuk and Lavrentjeva (1983) gave about 70°C lower temperatures for these samples.
Table 3

Results of exchange thermobarometry for gneisses and amphibolites

Gneiss sample

Grt–Bt: T (°C)

Grt–Bt–Pl–Qtz: P (kbar)

Ganguly

Holdaway

Wu

Wu

I-12

577 ± 6

608 ± 7

631 ± 5

7.45 ± 0.14

P95

596 ± 42

610 ± 58

629 ± 20

6.82 ± 1.53

P99

602 ± 46

645 ± 75

633 ± 16

7.62 ± 0.71

P200

541 ± 42

544 ± 46

578 ± 37

 

Amphibolite sample

Grt–Hbl: T (oC)

Grt–Hbl–Pl–Qtz: P (kbar)

GP

PL

DHP

KS

DHP

I-10a

660 ± 35

530 ± 54

615 ± 21

10.1 ± 0.5

9.8 ± 0.1

I-10b

684 ± 26

558 ± 16

680 ± 14

11. ± 0.4

11.8 ± 0.5

P34

668 ± 48

601 ± 49

695 ± 24

9.9 ± 0.3

10.2 ± 0.4

Ganguly Ganguly et al. (1996), Holdaway Holdaway (2000), Wu Wu and Cheng (2006), GP Graham and Powell (1984), PL Perchuk and Lavrentjeva (1983), DHP Dale et al. (2000), KS Kohn and Spear (1990). Grt garnet, Bt biotite, Pl plagioclase, Hb Hornblende, Qtz quartz

Pressures obtained by the exchange barometry of Wu and Cheng (2006) for the orthogneisses were in the range of 6.8–7.6 kbar, but the garnet–plagioclase–amphibole–quartz barometry in amphibolite gave higher pressures of 10–12 kbar using the calibration of Kohn and Spear (1990) and Dale et al. (2000). The higher pressure for sample I-12b is due to its more sodic plagioclase composition (An18) compared with the other samples that have plagioclase with An29–31. Similar pressures (10 ± 0.3 kbar at 570 ± 46°C) for garnet–amphibole–plagioclase-equilibria were obtained using average PT (Thermocalc, version 2.7) and the thermodynamic dataset of Holland and Powell (1998).

The results of pseudosection with isopleths of XCa, XFe, and XMn in garnet for sample P99 indicate 7.2 kbar/610°C for core compositions and 8.4 kbar/660°C for rim compositions of garnet (Fig. 9). These PT conditions are confirmed by all three isopleths of XCa, XFe, and XMn that intersect one another in the four variant fields with garnet, muscovite, biotite, and plagioclase. This temperature is about 20°C higher than the average temperatures obtained by all exchange thermometers, but it is within the error range resulting from analytical measurements and calculations. Pressure conditions obtained from the pseudosection fit well with those derived from the exchange barometry in orthogneisses. Such pressures are confirmed also by the presence of kyanite in the Al-rich metabasite (sample P182), which, according to data from the literature (e.g., Cooper 1980; Ward 1984; Spear 1982), indicates a pressure greater than 6 kbar. For amphibolite-facies temperatures, Selverstone et al. (1984) suggested that kyanite reflects pressures higher than those appropriate for the stability of the common amphibolite assemblage. Experimental studies corroborate this idea, and it was predicted (Schreyer and Seifert 1969) that Ged + Ky should be stabilised at pressures of about 10 kbar and that they are a precursor of Tlc + Ky “whiteschists” (Schreyer 1973).
Fig. 9

Pseudosection (MnNCKFMASH system) constrained for garnet–muscovite-bearing gneiss (sample P99) from the Shotur Kuh complex. All fields contain excess quartz. The open and filled circles indicate intersections of isopleths (XFe, XCa in squares and XMn in ellipses) for core and rim compositions of garnet

The phyllitic-micaschist (sample P200) gave temperatures of 541–578°C (Table 3), which are 50–60°C lower than those in the orthogneisses. Pressure conditions were not estimated, but they seem to be close to that in the orthogneisses.

Geochronology

The age of igneous crystallisation of the Shotur Kuh orthogneiss–amphibolite complex was obtained by laser ablation ICP-MS U–Pb dating of zircon extracted from an orthogneiss sample of granodiorite composition (P160). The data constrain the crystallisation age of the granodiorite to be at 547 ± 7 Ma (2 sigma concordia age; Fig. 10, Table 4). A similar age of 566 ± 31 Ma (n = 8, MSWD = 1.3) has been recently obtained for the Shotur Kuh Complex orthogneisses using ion microprobe and thermal-ionisation zircon dating (Hassanzadeh et al. 2008). Both ages are consistent with the U–Pb zircon ages of granitic and volcanic rocks reported from several localities in Central Iran, including Saghand area (Ramezani and Tucker 2003) and Sanandaj-Sirjan region (Hassanzadeh et al. 2008).
Fig. 10

U–Pb concordia diagram for zircon from orthogneiss (sample 160) in the Shotur Kuh complex

Table 4

U–Pb and Pb–Pb laser ablation ICP-MS data for zircons from sample P160

Analysis no.

Isotopic ratios

Ages (Ma)

207Pb/235U

±1 sigma

206Pb/238U

±1 sigma

207Pb/206Pb

±1 sigma

207Pb/235U

±1 sigma

206Pb/238U

±1 sigma

207Pb/206Pb

±1 sigma

1

0.6685

0.0294

0.0839

0.0026

0.0578

0.0007

519.8

22.8

519.1

16.0

522.6

25.8

2

0.7143

0.0315

0.0889

0.0032

0.0583

0.0007

547.3

24.1

549.0

19.5

540.2

27.8

3

0.6824

0.0284

0.0920

0.0024

0.0538

0.0012

528.2

22.0

567.6

14.8

361.6

49.4

4

0.6750

0.0243

0.0838

0.0025

0.0584

0.0006

523.8

18.8

518.7

15.4

545.9

21.7

5

0.7023

0.0181

0.0904

0.0015

0.0564

0.0005

540.1

13.9

557.6

9.5

467.1

20.4

6

0.6947

0.0343

0.0848

0.0027

0.0594

0.0010

535.6

26.5

524.9

16.5

581.5

36.3

7

0.7496

0.0460

0.0936

0.0046

0.0581

0.0009

568.0

34.9

576.9

28.2

532.5

32.4

8

0.6739

0.0348

0.0883

0.0025

0.0554

0.0015

523.1

27.0

545.3

15.3

427.4

61.9

9

0.7252

0.0474

0.0935

0.0030

0.0562

0.0009

553.7

36.2

576.4

18.7

461.5

37.3

10

0.7203

0.0445

0.0897

0.0035

0.0582

0.0012

550.9

34.0

553.9

21.7

538.2

43.3

11

0.7500

0.0297

0.0936

0.0026

0.0581

0.0009

568.2

22.5

576.5

16.2

535.3

34.8

12

0.7280

0.0285

0.0887

0.0022

0.0596

0.0009

555.4

21.8

547.6

13.7

587.5

33.4

13

0.7397

0.0283

0.0892

0.0031

0.0601

0.0006

562.2

21.5

550.9

18.9

608.2

22.3

Ar–Ar age dating was performed on metamorphic muscovite from a metagranite sample showing well-preserved, equilibrated fabric in amphibolite-facies conditions. However, a distinct low-temperature alteration is clearly visible and caused severe alteration of biotite and also minor alteration of feldspars. The age pattern of the muscovite from this sample corresponds well with the crystallisation sequence observed in thin section. It shows a distinct saddle-shaped pattern which points to a moderate age rejuvenation due to diffusional loss of radiogenic Ar during low-temperature reheating (Fig. 11). A single sample may not be sufficient to make far-reaching conclusions, but from our experience we conclude that the mica flakes contain age domains that are older than 166 Ma (probably about 170–180 Ma), which reflects the original cooling after metamorphism. A Lower Jurassic age of 171.8 ± 2.7 was also obtained by K–Ar dating of muscovite (sample I-13), although biotite from the same sample gave 208 Ma (Table 5). A younger age of 141 Ma was obtained for biotite in sample I-12.
Fig. 11

Ar/Ar ages of 166 Ma for muscovite in micaschist (sample P98 from the Shotur Kuh complex)

Table 5

Results of K–Ar dating for muscovite and biotite in the Shotur Kuh metamorphic complex

Samples

Rock type

Mineral

K2O (wt%)

% Rad 40Ar/tot 40Ar

Rad 40Ar (10−11 mol/g)

Age (±2σ) in Ma

I-12

Orthogneiss

Biotite

8.446

93.3

178.6

141.2 ± 2.2

I-13

Orthogneiss

Biotite

8.383

90.3

266.9

208.7 ± 3.2

I-13

Orthogneiss

Muscovite

8.454

96.4

219.4

171.8 ± 2.7

Discussion

Evidence for a Pre-Cambrian to Cambrian magmatic arc

The new geochronological data for crystallisation ages of granitoid protoliths of orthogneises from the SKMC are quite comparable with the granitoid ages reported by Ramezani and Tucker (2003) from the Saghand area in East-Central Iran who interpreted the igneous rocks as a magmatic arc of Late Neoproterozoic to Early Cambrian age (Fig. 12). This interpretation was recently confirmed by geochronological data from granites and orthogneisses from central Iran (Hassanzadeh et al. 2008). These ages, which range from Late Neoproterozoic to Early Cambrian, match the mostly juvenile Arabian–Nubian shield and Peri-Gondwanan terranes that have formed after the main phase of Pan-African orogeny. Consistent with paleogeographic reconstructions for the Late Neoproterozoic (Powell et al. 1980; Lawver and Scotese 1987; De Wit et al. 1988; Unrug 1997) the arc was formed by the closure of the Chapedony Ocean (Proto-Tethys, Fig. 13) (Ramezani and Tucker 2003). Following the Pan-African orogeny and the consolidation of the basement, the Precambrian craton of Iran, Pakistan, central Afghanistan, southeastern Turkey, and Arabia became a relatively stable continental platform with epicontinental shelf deposits (mainly clastics) and was characterised by a lack of major magmatism or folding (Nadimi 2007).
Fig. 12

The “modified alkali-lime index” (MALI) versus SiO2 diagram after Frost and Frost (2008) for Shotur Kuh Metamorphic Complex orthogneisses (field labelled SKMC, circles represent individual analyses). Other fields represent: BSG Boneh Shurow granitic gneisses from Central Iran (Ramezani and Tucker 2003), CGS Cambrian Granitic Complex from Central Iran (Ramezani and Tucker 2003), BM granitoids of the Bitlis Massif in SE Turkey (Ustaömer et al. 2008). For comparison, variation fields for typical magmatic arc intrusive complexes of the Peninsular Ranges and Tuolomne Batholith are displayed (from Frost and Frost 2008). Terminology for individual igneous suites: a alkalic, a-c alkali-calcic, c-a calc-alkalic, and c calcic

Fig. 13

Schematic Gondwanaland reconstruction in Early Cambrian based on Molleweide projection of tectonic plates at 540 Ma by the PLATES Project (1999), with modifications by Ramezani and Tucker (2003). Precambrian cratons are: Af African, Am South America, An Antarctic, Au Australia, and In India, Ch South China, L Lhasa, and MT Menderes-Taurus Blocks. Stars refer to Peri-Gonwana arc plutons

Despite large compositional variability, the overall geochemistry of the orthogneiss varieties from the Shotur Kuh complex is compatible with an origin of their granitoid protoliths within a magmatic arc. Based on the trace-element discrimination diagrams (Pearce et al. 1984), the majority of orthogneisses plot within the “volcanic arc granite” field (Fig. 4). We suggest that the arc was situated at a continental margin, where melting and recycling of abundant older continental material could play a significant role. Some orthogneiss analyses are very comparable with rocks described by Ramezani and Tucker (2003) as being members of the Late Neoproterozoic to Early Cambrian arc igneous assemblage from the Saghand area (see Fig. 1 and Table 6)—namely, with the “Cambrian Granitoid Suite” (CGS) and with the acidic orthogneiss “Boneh-Shurow Granitic Gneiss” (BSG) dated as Late Neoproterozoic. When comparing these orthogneisses with Late Neoproterozoic to Early Cambrian igneous and metaigneous rocks in the nearby regions they show close similarity to continental arc-related granitoids and metagranitoids in SE Turkey (Ustaömer et al. 2009). Figure 12 displays geochemical signatures of the Shotur Kuh orthogneisses in the Frost et al. (2001) diagram and compare them with some plutonic rocks from other parts of the peri-Gondwanan and also much younger arc-related plutonic complexes from North America. Despite of their rather broad scatter, the Shotur Kuh orthogneisses fit well with arc-related calcic and calc-alkalic suites and together with other granitoid or metagranitoid complexes from the peri-Gondwanan arc they correspond to silica-rich, evolved rocks. Although some geochemical features, namely the high total Fe and Zr contents, of the tonalitic orthogneiss (sample I-10) appear to be far from typical arc assemblages they are broadly similar to some subvolcanic members of the Cambrian Volcano-Sedimentary Unit of Saghand, which has been interpreted as arc-related (Ramezani and Tucker 2003).
Table 6

Comparison of selected geochemical parameters for the studied orthogneisses with other Late Neoproterozoic to Cambrian granitoids in Central Iran and SE Turkey

 

Orthogneisses, Shotur Kuh metamorphic complex (this study)

Cambrian Granitoid Suite (Ramezani and Tucker 2003)

Boneh-Shurow granitic gneisses (Ramezani and Tucker 2003)

Bitlis Massif, SE Turkey (Ustaömer et al. 2009)

SiO2

64.3–75.9 (63.72)

65.9–76.3

71.7–77.7

68.0–78.3

K2O

1.33–4.61 (0.47)

2.86–4.92

2.7–4.8

0.2–4.6

K2O/Na2O

0.39–1.6 (0.146)

0.8–1.39

0.6–1.3

0.02–1.2

MALI

1.6–6.8 (−2.0)

2.9–8.7

5.8–7.2

1.6–7.8

A/CNK

0.866–1.24 (0.833)

0.923–1.004

0.94–0.99

0.82–1.09

Th/Ta

9.9–29.3

17.5–44.4

11–22

Eu/Eu*

0.20–1.02

0.38–0.59

0.30–0.68

Values in parentheses represent the orthogneiss I-10 that is significantly different from other rock analyses representing the Shotur Kuh metamorphic complex. MALI is “Modified Alkali-Lime Index” Na2O + K2O − CaO (Frost et al. 2001); A/CNK is molar ratio Al2O3/(CaO + Na2O + K2O)

The tectonic positions of mafic rocks (amphibolites) spatially associated with orthogneisses, are difficult to interpret, mainly because of metamorphism and deformation, which could have modified their original geological relationships. Some small amphibolite bodies could represent original mafic enclaves. However, the majority of mafic bodies may represent metamorphosed primary dykes.

The geochemical indications for the original tectonic setting of mafic rocks are rather ambiguous. In various discrimination diagrams, amphibolites plot in the fields of “within-plate basalts” and MORBs, and also of volcanic arc basalts. Because of the relatively high Th contents, most samples are shifted towards the “Calc-alkaline basalt” field or plot within this field (Fig. 6). The increased Th/Ta ratios (Fig. 5) may indicate that the original basaltic magmas were derived from a mantle source affected by subduction-related processes.

Some geochemical features of amphibolites are consistent with the origin of their protolith magmas at a destructive plate margin—i.e., in a magmatic arc (increased LILE/HFSE and Th/Ta element ratios, Fig. 5). However, the “subduction fingerprint” is rather weak and could also be explained in terms of partial melting of the sub-continental lithospheric mantle (with the long-term “subduction memory”) and/or contamination of mafic magmas by continental crust (including protoliths of orthogneisses). Except for the two most mafic samples that are probably affected by the accumulation of early mineral phases, the rocks display relatively high TiO2 content, which is in contrast to the composition of typical volcanic arc magmas. Mafic to intermediate rocks with strong, typical fingerprints of subduction-related magmas are absent, and therefore the arc igneous assemblage is incomplete and the outcropping rock complexes involving metamorphosed mafic dykes may represent an extensional setting of the continental back-arc rather than a true arc near the volcanic front. However, even such a setting can be supposed to be a part of a much greater Late Neoproterozoic–Early Paleozoic orogenic system that was active along the Proto-Tethyan margin of the Gondwanaland supercontinent, extending at least from its Arabian margin to the Himalayan margin of the Indian subcontinent (Ramezani and Tucker 2003).

Metamorphic evolution

Mineral assemblages observed in the studied rocks indicate a medium-pressure, Barrovian-type metamorphism for which a maximum pressure of ca. 8 kbar and a temperature of about 650°C were obtained. This metamorphism produced garnet, biotite, plagioclase, quartz, and muscovite in orthogneisses, and garnet, amphibole, plagioclase, and quartz in amphibolite. The overlaying pre-middle Triassic sedimentary sequences were also affected by this metamorphism, but they show a lower temperature (550°C). Garnet, both in metaigneous and metapelitic rocks, indicates prograde zoning from cores to rims of grains. This zoning is strong in metapelites and weak in metaigneous rocks, where it shows additional modification at rims (retrograde zoning) that probably occurred during cooling. The weak prograde zoning in metaigneous rocks could be the result of either higher temperatures or longer relaxation times at high temperature. Lithological similarity but different PT conditions between Pre-middle Triassic pelitic rocks and garnet micaschists suggest either tectonic juxtaposition of single unit with differing metamorphic grade or the micaschists belong to an older basement unit, which reached PT conditions close or similar to the orthogneisses.

The limitation of the Middle- to Lower-Jurassic ages of 165 Ma for the amphibolite-facies metamorphism, obtained by Ar–Ar data on muscovite, is supported by the presence of the Upper Triassic–Lower Jurassic Shemshak formation and by the Middle-Jurassic psammopelitic rocks with conglomerates containing pebbles of basement orthogneisses and amphibolites. The stratigraphically well-defined Shemshak formation in Central Iran (e.g., Assereto 1966; Stampfli 1978) shows lower greenschist facies metamorphism with fabrics comparable to those in the underlying limestone–dolomite formation (Rahmati-Ilkhchi et al. 2008). As the Middle-Jurassic sediments with conglomerates show only very low-grade metamorphic conditions, the greenschist facies overprint (fine-grained white mica, chlorite with albite, epidote, and, locally, also brown-green biotite replacing garnet or plagioclase) in the basement rocks could be related to retrogression during exhumation.

The results of our petrological and geochronological study agree well with the geodynamic scenario for closure of the Neotethyan basin in the southwestern part of Iran (e.g., Stöcklin 1968; Berberian and King 1981; Sengör 1991; Glennie 2000; Sheikholeslami et al. 2008; Ghasemi and Talbot 2006). This basin was formed by continental rifting, separating the Iranian and the Arabian plates (Fig. 14). Formation of a magmatic arc with an accretionary prism during the Early Cimmerian orogeny (Sheikholeslami et al. 2008; Paul et al. 2003) is coeval with the Middle Jurassic cooling age of white mica from the Shotur Kuh complex, which defines the collisional processes and metamorphism of the basement rocks of the Central Iran Block.
Fig. 14

Geodynamic reconstruction of the southwest margin of central Iran and the north Gondwanian margin from Late Paleozoic to Cenozoic (modified from Berberian and King 1981 and Sheikholeslami et al. 2008). a Development of rift-type basin from Permian. b Subduction of the Neotethys oceanic basin in Late Triassic (Early Cimmerian phase) and formation of an accretionary prism and a magmatic arc (granite/granodiorite of Chah Dozdan). c Collision and generation of an orogenic prism with metamorphism and exhumation of the basement units (Saghand-Shotur Kuh metamorphic complexes) and development of the volcanic arc of Urumieh-Dokhtar zone. A Arabian plate, Z Zagros, Ss Sanandaj-Shirin, CI Central Iran, and Al Albros

The significance of the Shotur Kuh metamorphic complex for the geological history of the Great Kavir Block

The investigated metaigneous rocks from the Shotur Kuh complex show arc-related signatures similar to those described by Ramezani and Tucker (2003) and confirm a uniform geotectonic position for the Central Iran micro-continent during the Late Neoproterozoic, including the GKB. This interpretation is also supported by comparative stratigraphy of Paleozoic to Mesozoic sedimentary sequences among different crustal blocks (Stöcklin 1968; Takin 1972; Crawford 1977; Berberian and Berberian 1981). Most basement rocks in Central Iran are exposed along tectonic zones separating individual blocks. In some cases, these zones are followed by Paleozoic to Cenozoic ophiolites with or without blueschists (Fotoohi Rad et al. 2005; Bagheri and Stampfli 2008), which suggests large-scale tectonic activity with rifting and subduction. The Sabzevar zone with ophiolites and blueschists is the best example of a back-arc basin that was opened during Cretaceous north from the Neothethys ocean (Bagheri and Stampfli 2008). The SKMC is exceptional, since it occurs in the central part of the GKB. However, the positions of the SKMC and some mafic and ultramafic rocks near Kuh Siah Poshteh (100 km from the Shotur Kuh complex) along the E–W trending Torud fault zone may suggest that the Torud fault zone was also active prior to Tertiary volcanism. Considering the N–S compression of the Shotur Kuh complex with south-vergent isoclinal fold axes (Rahmati-Ilkhchi et al. 2008), the Torud tectonic zone may represent a large-scale thrust zone that was responsible for formation (?) and exhumation of the amphibolite-facies metamorphic rocks during the Jurassic. The rapid increase in the degree of metamorphism from top to bottom in the Pre-middle Triassic rocks in contact with orthogneisses is either due to tectonic reduction of relatively thick sequences or the result of heating of partially exhumed orthogneiss–amphibolite rocks to the upper-crustal level. To confirm or reject the hypothesis of large thrust and underplate tectonics along this zone, more detailed structural analyses and dating of mafic and ultramafic rocks is needed.

Conclusions

Geochemical analyses and U–Pb zircon age data from orthogneisses in the Shotur Kuh metamorphic complex indicate that these rocks are derived from granitoid protoliths (granite, granodiorite, and tonalite) formed during Late Neoproterozoic arc-related magmatism. This magmatism, which is known also from other zones and blocks in Central Iran, confirms that the GKB, along with the rest of the blocks of the Central Iran micro-continent, were part of the Neoproterozoic–Early Paleozoic orogenic system that was active along the Proto-Tethyan margin of the Gondwanaland supercontinent. In the Shotur Kuh complex, these rocks were affected by Barrovian-type metamorphism that reached amphibolite-facies conditions in the deeper part of the exposed basement with orthogneisses and amphibolites, and lower amphibolite facies in the sedimentary cover. Both igneous and sedimentary rocks show evidence of prograde metamorphism with subsequent cooling. The results of petrological studies with structural data about N–S compression and north-dipping structures suggest that the Lower- to Middle-Jurassic metamorphism and exhumation of the Shotur Kuh complex occurred under conditions of south-vergent thrust tectonics along the E–W zone. The latter was probably related to the closure of the Neotethyan basin and subsequent collision during the Early Cimmerian orogeny.

Notes

Acknowledgments

This work is a part of the PhD thesis of M. Rahmati-Ilkhchi at the Institute of Petrology and Structural Geology (Charles University in Prague) and was supported by the Geological Survey of Iran as well as by the Ministry of Education, Youth and Sports of the Czech Republic (Research Project MSM21620855). K. Schulmann (Strasbourg) is thanked for facility to obtain K–Ar age data of biotite and muscovite. Christoph Hauzenberger and Marlina A. Elburg are thanked for careful and thorough reviews of the manuscript. We also thank Ingo Braun for helpful comments and editorial responsibility.

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Copyright information

© Springer-Verlag 2009

Authors and Affiliations

  • Mahmoud Rahmati-Ilkhchi
    • 1
    • 2
  • Shah Wali Faryad
    • 1
  • František V. Holub
    • 1
  • Jan Košler
    • 3
  • Wolfgang Frank
    • 4
  1. 1.Institute of Petrology and Structural GeologyCharles University in PraguePragueCzech Republic
  2. 2.Geological Survey of IranTehranIran
  3. 3.Department of Earth Science and Centre for GeobiologyUniversity of BergenBergenNorway
  4. 4.Central European Argon LaboratorySlovak Academy of SciencesValasška, BratislavaSlovakia

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