Arabian Journal of Geosciences

, Volume 8, Issue 2, pp 849–866 | Cite as

Soltan Maidan Complex (SMC) in the eastern Alborz structural zone, northern Iran: magmatic evidence for Paleotethys development

Original Paper

Abstract

Soltan Maidan Complex is located in the eastern Alborz zone, Iran. Thickness of this complex ranges from 436 to 1,286 m and its age is Upper Ordovician–Silurian. It is the most important and voluminous Lower Paleozoic magmatic activity in Iran. This complex consists of mafic lava flows, agglomerates, and tuffs accompanied with some sedimentary intercalations. Field geology evidence indicate both subaerial and submarine eruptions during accumulation of this complex and also occurrence of a long-time exhumation event in the study area from Upper Ordovician to Middle Devonian time. The volcanic rocks of Soltan Maidan Complex indicate a range composition of basalt to basaltic-andesite and show transitional to alkaline nature of the primary magma along with evolution via fractional crystallization process. These rocks represent fairly high degrees of partial melting (between 14–20 %) from a common garnet peridotitic mantle source in an intra-continental rift setting. Relatively high Nb and Zr contents (i.e., 21.68 and 189.63 ppm, respectively) show the generation of magma from an enriched mantle source. Also, Th/La (with an average of 0.13) and Zr/Nb (with an average of 9.11) ratios of the samples illustrate mixing between an OIB-like magma with an EMI-type mantle source. The primary magma shows lack or minor role of the crustal contamination during its ascent. According to the many evidences, Soltan Maidan magmatism has occurred in the Upper Ordovician–Silurian times in relevant to a cycle of rift-related magmatic events and breakup of the northern margin of Gondwana in the process of Paleotethys development in the north of Iran.

Keywords

Soltan Maidan Basalt Paleotethys Rift Paleozoic Alborz 

Introduction

The Iranian plateau was divided into nine geological–structural zones by Stocklin (1968) including: Kopeh-Dagh, Alborz, Central Iran, Sanandaj-Sirjan, Imbricate Thrust Zagros, Simply Folded Belt Zagros, Khuzestan Plain, Makran and East Iran, and Lut Block zones (Fig. 1). Alborz zone, which is located in the north of Iran, is an active zone with a general E–W trend, formed by the Gondwana and Eurasia collision after closure of Paleotethys Ocean (Sengör and Burke 1978). This zone has evolved by the Cimmerian and Alpine orogenies during Triassic to the present day (Alavi 1996) and consists of Western, Central, and Eastern parts (Fig. 1).
Fig. 1

Geological–structural map of Iran (modified from Stocklin 1968) and location of the study area

Lower Paleozoic mafic rocks are generally seen in some parts of Iran (e.g., Alborz, Central Iran and Sanandaj-Sirjan zones). Soltan Maidan Complex (SMC) with Upper Ordovician–Silurian age is a part of eastern Alborz, representing the most important, widespread, thick, and voluminous lower Paleozoic magmatic activity in Iran. It has NE–SW trend and crops out in the eastern Alborz zone (Fig. 1). In different outcrops, SMC thickness varies from 436 to 1,286 m and dominantly composes of basalt to basaltic-andesite rocks associated with some shale, siltstone, sandstone, and conglomerate layers.

This paper describes field studies, petrography, geochemistry, petrogenesis, and tectonic setting of SMC in order to understand some aspects of the most voluminous lower Paleozoic magmatic activity in Iran and its relation to opening of the Paleotethys Ocean.

Geology

The exposures of the SMC investigated in this study are Kaludar (a) and Cheshmeh-Seyed (b) Valleys in northwest and north of Shahrud, respectively, and Khoshyeilaq village (c) in northeast of Shahrud (Figs. 1 and 2). SMC consists of a thick succession of basalt to basaltic-andesite lava flows (as the dominant volume), agglomerate, tuff and a number of sedimentary intercalations. Most of the volcanic rocks have been altered and minerals converted to chlorite, epidote, calcite, quartz, titanite (sphene), Fe-Ti oxides, and clay minerals.
Fig. 2

Simplified geological maps of the SMC, Ghelli and Padeha Formations around a Nekarman, b Qaleh-e-Nowkhareqan, and c Khoshyeilaq villages

Amygdaloidal texture and columnar joints are two main features of the volcanic rocks in the study area. The majority of lava flows consists of an upper and sometimes lower amygdaloidal-rich zones. The distribution pattern of amygdaloidal texture in SMC lava flows is a clue to recognize and separate lava flows from each other. The upper zone is more amygdaloidal, broader and contains larger amygdaloidals than the lower zone, probably as a result of coalescence of rising primary bubbles during solidification of lava flows. Amygdales are sometimes elongated which indicate the flow direction. These were often filled by secondary minerals such as chlorite, quartz, and calcite.

Delaloye et al. (1981) reported some K–Ar ages for SMC ranging from 633 to 181 Ma. Although those ages are widespread, the SMC rocks lay uncomfortably on the shale, siltstone, and sandstone of Ghelli Formation which are Late Ordovician in age (Ghavidel-Syooki and Winchester-Seeto 2002; Ghavidel-Syooki et al. 2011). SMC also overlay with an erosional surface by sandstones and conglomerates of Padeha Formation belonging to Early–Middle Devonian (Bozorgnia 1973; Ahmadzadeh Heravi 1975; Jenny 1977; Hamdi and Janvier 1981; Wendt et al. 2005; Aharipour et al. 2010). Moreover, paleontological studies of thin shale intercalations of the SMC by Ghavidel-Syooki et al. (2011) in the Khoshyeilagh area, which is another outcrop of this complex, show Late Ordovician–Middle Silurian ages. They pointed out that SMC is not younger than Gorstian (early Ludlow, early late Silurian). This complex is thus late Ordovician–Silurian in age.

Kaludar Valley

Kaludar Valley is the westernmost outcrop of SMC and is located in the north of the Nekarman village (with 36°32′N and 54°50′E), northwest Shahrud city (Figs. 1 and 2). SMC lies with a concordant contact on the olive-gray shale-siltstone and sandstone of Ghelli Formation with Upper Ordovician age (Fig. 3). The upper contact is erosional and covered by conglomerates and sandstones of the Padeha Formation which is Early–Middle Devonian in age. Aharipour et al. (2010) mentioned that Padeha Formation deposited in a syn-rift terrestrial environment during Early–Middle Devonian.
Fig. 3

Stratigraphic columns for the SMC, upper part of Ghelli Formation and lower part of Padeha Formation at Kaludar Valley, Cheshmeh-Seyed Valley, and Khoshyeilaq area sections

The Kaludar Valley contains a complete succession of the Paleozoic rocks from middle Ordovician to Devonian (Figs. 3 and 4a). General thickness of SMC in the Kaludar Valley is 436 m (Fig. 3). At least 20 lava flows were exposed in the Kaludar Valley with thicknesses between 3 and 58 m. They include an interlayer of sedimentary rocks as red to purple sandstone and micro-conglomerate of river channels with about 1 m in thickness.
Fig. 4

General view of the SMC in the Kaludar Valley, northern Nekarman village. a General view showing the complete succession of Paleozoic rocks from Upper Ordovician to Devonian in the study area. In this outcrop SMC is covered by sandstones and conglomerates of Padeha Formation and lies on the shale, siltstone and sandstone of Ghelli Formation. b General view showing well-developed primary columnar joints with the height of about 22 m at the base of SMC

Kaludar Valley is dominantly composed of tabular successions of green–black basalt to basaltic-andesite lavas. At this section, many lava flows (spatially at the basal parts) are characterized by sub-columnar to well-developed primary columnar joints (Figs. 3 and 4b).

The height of columnar joints at the base of the complex is up to 22 m (Fig. 4b). The presence of intercalated red to purple sandstone, micro-conglomerate and also primary columnar joints in some lava flows could be possibly evidence of subaerial rather than subaqueous eruptional environment.

There are some older and younger volcanic–subvolcanic rocks than SMC as sills, dykes, and basaltic lavas in the Kaludar Valley. The older ones were found to be form of basaltic lava flows and dykes in the Ghelli Formation (Fig. 3). Based on the field studies, these dykes are substantially younger than their host rocks; so they might fed Soltan Maidan volcanism during late Ordovician–Silurian times. The presence of these dykes indicates occurrence of extrusions along the fissures and explain some kind of relationship between the volcanism and the fault activities in this time. The volcanism was continuing from Silurian to Devonian with some basaltic lava flows, sills, and dykes (Fig. 3).

Cheshmeh-Seyed Valley

Cheshmeh-Seyed Valley is located in the northeast of Kaludar Valley and approximately 20 km away from it. It outcropped in the northwest of Qaleh-e-Nowkhareqan village (with 36°38′N and 55°04′E; Fig. 1). The Cheshmeh-Seyed Valley is similar to the Kaludar Valley and represents a complete rock succession of Upper Ordovician to Devonian. It is exposed between Ghelli and Padeha Formations with concordant contacts (Figs. 2 and 3).

SMC is 1,286 m in thickness and consists of nearly 50 typically sheet-like basaltic lava flows, some thin basaltic agglomerates (located in the middle parts of the complex), two tuff units, and two sedimentary intercalated units (Fig. 3) in the Cheshmeh-Seyed Valley. Compared to Kaludar Valley these are generally less altered. Total thickness and the number of lava flows in the Cheshmeh-Seyed Valley are more than Kaludar Valley but on the basis of the field studies, most properties of both lava flows (e.g., color, general composition, etc.) are the same. Also, some lava flows with primary columnar joints, exist in this area (Fig. 3).

Another difference between the basaltic rocks of the Kaludar Valley and the Cheshmeh-Seyed Valley is the presence of two thick units of sedimentary rocks in the latter, one with 14 m and the other with 85 m thickness (Figs. 3 and 5a). These units, which mainly consist of alternating layers of green shale-siltstone and sandstone, were situated in the middle and upper parts of complex, respectively (Fig. 3). Presence of some brachiopods and greenish color of the units are indicating of subaqueous condition during their accumulation. It is important to mention that a lens shape conglomerate with up to 5 m thickness lies on the top of the upper sedimentary layer containing enormous rounded pink and white granitic clasts (pebbles) with clast-size up to about 50 cm in diameter (discussed in the following sections) (Figs. 3 and 5b).
Fig. 5

The photographs show both an overview of SMC and a close up view of conglomerate within it in the Cheshmeh-Seyed Valley. a General overview showing middle to upper parts of SMC and lower part of the Padeha Formation. Two distinct intercalated sedimentary units are shown as dashed lines. b Close up view of a conglomerate unit contains enormous rounded pink granitic clasts

Khoshyeilaq area

The third outcrop of SMC is located approximately 40 km away from the Cheshmeh-Seyed Valley in the northeast direction, around the Khoshyeilaq village with 36°51′N and 55°21′E (Figs. 1 and 2). It is the easternmost outcrop of the SMC. Although its lower contact with Ghelli Formation is invisible due to folding and faulting (Fig. 3), it seems that its total primary thickness, compared to the other outcrops, is substantial.

In this area, volcanic rocks are extensively altered. Apart from the difference in the alteration degrees of igneous rocks and also particular position of lower contact of complex, general properties (e.g., general composition, presence of primary columnar joints, presence of sedimentary layers, etc.) are similar to Cheshmeh-Seyed Valley. There is an intercalated sedimentary layer of green shale-siltstone, sandstone, and conglomerate (up to 100 m in thickness) in the Khoshyeilaq area (Fig. 3). Some evidences show that this sedimentary unit was deposited in a submarine environment. Paleontological studies by Ghavidel-Syooki et al. (2011) suggest Late Ordovician age for it. The presence of a lava flow with pillow structure located just beneath this sedimentary layer supports submarine condition of accumulation (Figs. 3 and 6a). In addition, a conglomerate unit with moderate thickness of 7 m containing enormous rounded pink granitic clasts with clasts-size of up to about 50 cm in diameter have been spotted in the upper part of the complex (Figs. 3 and 6b). Properties of the conglomerate are totally similar to conglomerate in the Cheshmeh-Seyed Valley. Ghavidel-Syooki et al. (2011) carried out U–Pb dating on the zircons of these granitic clasts and obtained ages of 434.4 ± 6.4 Ma (late early Silurian). They emphasize that plutons with this age are entirely unknown throughout Iran, and host conglomerate of these granitic clasts is stratigraphically located at the base of the complex. It is worth to note that paleontological study of a thin shale layer located over this conglomerate unit by Ghavidel-Syooki et al. (2011) shows Silurian age for this part of SMC. In addition, they argue that granite emplacement, cooling, exhumation, erosion, transportation, and deposition of its clasts took place in a time interval between 5 and 10 Ma years. However, the presence of similar granites with this age is not completely unknown throughout Iran and it seems that they are similar to granitoids which was reported by Ghasemi and Khanalizadeh (2011) in the Touyeh-Darvar area, about 120 km farther from the study area.
Fig. 6

Close up view of a pillow lava structure and b a conglomerate with rounded pink granitic clasts within the SMC in the Khoshyeilaq area

Petrography

A total of 270 samples from volcanic and sedimentary rocks of Kaludar and Cheshmeh-Seyed valleys and Khoshyeilaq area were studied by polarizing microscope. The volcanic rocks have the same mineralogical and textural features and mainly consist of weakly to strongly altered basalt to basaltic-andesite rocks. Tuffaceous rocks are also present but they are not considerable in area. Basaltic rocks have aphanitic, porphyritic, seriate, intersertal, trachytic, ophitic, sub-ophitic, poikilitic, glomeroporphyritic, and amygdaloidal textures (Fig. 7).
Fig. 7

Thin-section images of SMC. a Seriate texture in the basaltic rocks. As shown phenocrysts are plagioclase and clinopyroxene. b Cluster of large clinopyroxene phenocrysts included in a fine-grained microlitic and glassy groundmass. c Sub-hedral phenocrysts of olivine, entirely replaced by calcite and Fe-oxides. Also, this image showing some amygdules, filled by epidote and quartz. d An image of crystal tuff with plagioclase, clinopyroxene and opaque clast fragments from lower part of Cheshmeh-Seyed Valley section. e Amygdaloidal texture in a basaltic sample with a fine-grained microlitic and glassy groundmass. Amygdules are filled by quartz and chlorite. Also, all glasses are converted into chlorite. f A sed-arenite with sedimentary clasts of siltstone, sandstone and chert from sedimentary layer located in the upper part of SMC in the Cheshmeh-Seyed Valley section

The volcanic rocks consist dominantly of phenocrysts of plagioclase (up to 25 %), clinopyroxene (up to 10 %), and subordinate olivine and/or orthopyroxene (up to 3 %) and opaque minerals (mainly Fe–Ti oxides, up to 5 %; Fig. 7a–c). Plagioclase and clinopyroxene constitute phenocryst and microcryst in the groundmass (Fig. 7a). Pyroxene and also olivine (that exists in some lava flows) are primary mafic phases in the rocks. Also, tuffaceous rocks consist of plagioclase, clinopyroxene, and opaque crystal fragments (Fig. 7d).

The extent of alteration in the studied rocks varies from slight to severe, where phenocrysts remain only as relict forms or hand specimens appear green due to chloritization. Almost in all samples, plagioclase has undergone high alteration and it has been partly or completely replaced by the assemblage of sericite, calcite, clay minerals and sometimes epidote.

Clinopyroxene is the second common mineral phase in these rocks and is relatively fresh than the other minerals (Fig. 7). It consists of both augite and Ti-rich augite. Ti-rich augite is abundant and can be easily identified by the pinkish to pinkish-brown colors. It represents occasionally hourglass zoning with slight pleochroism which indicates that it has high-Ti contents. Olivine, the most altered mineral in the samples, is completely replaced (pseudomorphically) by secondary minerals (e.g., iddingsite, serpentine, calcite, chlorite, quartz, epidote and iron oxides). Consequently, olivine is recognized via pseudomorphs, replaced entirely by secondary minerals.

The secondary minerals in the studied rocks include chlorite, sericite, epidote (Fig. 7c), calcite, quartz, titanite, Fe-Ti oxides, and clay minerals. Chlorite is the most predominant secondary mineral (Fig. 7e). Fe–Ti oxides are sometime replaced by a light gray aggregate of cryptocrystalline titanite (leucoxene).

The groundmass of volcanic rocks is microlitic or glassy. The glassy groundmass has often been devitrified into more stable crystalline phases such as chlorite (Fig. 7e).

Alkali elements (Na, K, Rb) are the most sensitive elements to alteration and their concentrations can be affected even in slightly altered samples (Rollinson 1993). Their mobilities are influenced by breakdown of interstitial glass and alkali bearing phases which develop secondary minerals.

Amygdaloidal texture is very distinctive for all lava flows, both in hand sample and thin-section scale (Fig. 7c, e). Amygdales have been filled by secondary minerals such as calcite, chlorite (Fig. 7e), quartz (Fig. 7e), Fe-oxides and epidote (Fig. 7c), and clay minerals.

Interlayer sedimentary rocks in Cheshmeh-Seyed Valley and Khoshyeilaq area include predominantly sed-arenite with clasts of shale, siltstone, sandstone, and chert (Fig. 7f).

Analytical methods

Some 19, 15, and 7 samples from Kaludar Valley, Cheshmeh-Seyed Valley and Khoshyeilaq area, respectively, were analyzed for major and trace elements after fusion of 0.2 g of rock powder with 1.5 g LiBO2 and dissolved in 100 mL 5 % HNO3 by inductively coupled plasma spectroscopy (ICP) at SGS Analytical Laboratories, Toronto, Canada. Loss on ignition was determined by drying the samples heated at 1,000 °C. Major elements and some trace elements including Ba, Sr, Y, Zn, and Zr were analyzed via Inductively Coupled Plasma–Atomic Emission Spectrometry (ICP–AES). A further suite of trace elements include Co, Cs, Cu, Ga, Hf, Nb, Ni, Rb, Ta, Th, U, V, and rare earth elements (REEs) were analyzed by Coupled Plasma–Mass Spectrometry (ICP–MS). Major elements analyzed with detection limit of 0.01 % and trace elements analyzed with detection limits between 0.05 to 10 ppm (U, Pr, Eu, Gd, Tb, Dy, Ho, Er, Tm, Lu 0.05 ppm, Cs, Th, Y, La, Ce, Nd, Sm, Yb 0.1 ppm, Rb 0.2 ppm, Ta, Co 0.5 ppm, Hf, Nb, Ga 1 ppm, Ni, V, Zn, Cu 5 ppm, Ba, Sr 10 ppm). Major and trace element compositions of the samples are given in Table 1.
Table 1

Analytical results of major (%) and trace elements (ppm) of the basaltic rocks of the SMC

Location

Kaludar Valley

Kaludar Valley

Khoshyeilaq area

Cheshmeh-Seyed Valley

Cheshmeh-Seyed Valley

Sample

K 1A

K 2B

KH 1A

K 5A

K 7B

K 10A

K 11B

K 13A

K 17A

K 19A

T.A.7.8

T.B.1.4

T.B.1.7

T.C.2.1

T.D.1.8

T.D.2.1

T.D.2.3

T.D.3.3

T.D.3.4

KH 2

KH 3B

KH 4

KH 5C

KH 9B

KH G4

M 1A

M 2D

M 2G

M 5B

M 7B

M 10C

M 12A

M 16A

M 18A

M 21B

M 27

M 30

M 35

M 42A

M44A

SiO2

51.2

50.4

49.9

48.0

51.4

47.5

51.5

50.8

48.5

55.2

48.4

48.4

48.2

49.5

45.1

49.7

47.0

47.1

49.4

48.5

50.3

49.5

49.4

50.6

46.7

45.9

48.1

49.5

48.4

48.8

49.2

49.5

50.2

50.9

49.9

49.4

55.0

51.6

48.0

52.4

TiO2

2.95

2.89

2.99

2.02

2.46

1.61

1.68

2.1

2.07

2.8

3.04

3.04

2.38

2.67

3.12

2.95

2.63

3.30

2.07

3.34

2.81

2.33

2.24

2.53

2.56

2.11

1.86

2.8

2.31

2.03

1.87

1.89

1.77

1.95

2.1

1.83

2.07

2.93

3.16

1.74

Al2O3

13.3

13.3

12.4

15.6

14

12.7

14.5

13.6

13.5

14.2

13.60

13.70

15.80

13.20

13.80

12.80

15.10

15.30

17.00

12.8

16.3

12.3

13.3

12.8

13.3

15.4

14.2

13

14.3

13.5

13.5

13.5

14.1

14.8

13.2

13.4

13.4

13.6

12.6

14.4

Fe2O3(t)

13.7

12.5

14.2

10.2

11.9

12.6

9.68

11.9

12.5

8.38

13.30

13.10

11.20

13.80

13.90

13.80

12.00

12.30

9.83

14.9

11.1

12.6

12

13.5

13.6

10.7

11.5

13.6

11.7

11.7

11.4

11.3

11.1

10.9

12.6

11.3

11

10.2

14.7

11

MnO

0.23

0.17

0.26

0.27

0.26

0.24

0.19

0.23

0.22

0.16

0.33

0.28

0.30

0.25

0.28

0.21

0.21

0.24

0.19

0.19

0.11

0.27

0.2

0.3

0.25

0.17

0.22

0.2

0.25

0.18

0.18

0.17

0.18

0.19

0.21

0.19

0.16

0.21

0.22

0.19

MgO

4.2

4.24

3.98

5.57

5.2

10.1

6.07

6.37

6.1

4.23

4.28

4.41

5.61

4.40

4.71

4.46

6.45

4.48

5.15

3.96

2.8

5.96

5.75

5.34

6.02

6.93

6.32

4.5

5.84

6.63

6.4

6.31

5.55

4.88

5.72

5.87

3.64

5.35

4.78

4.37

CaO

7.88

8.69

7.78

7.47

9.43

8.27

6.25

8.09

8.03

4.01

6.37

8.15

7.82

7.27

5.13

7.09

6.72

4.26

4.35

7.07

4.41

5.6

7.46

6.16

7.52

4.6

7.3

6.76

7.9

9.18

8.11

8.2

8.98

7.14

8.22

7.76

6.69

4.2

9.03

6.67

Na2O

2.7

2.5

2.9

4.1

2.5

1.8

3.5

3

2.6

5.9

3.80

3.20

4.90

3.20

4.80

2.90

4.40

5.30

5.80

2.7

6.5

3.5

3.3

3.4

3.7

4.3

3.6

3.9

3.5

2.7

3.2

3.3

2.4

2.8

2.6

4.3

2.9

4.4

2.3

4.4

K2O

0.86

0.81

0.43

0.55

0.95

0.99

1.11

0.83

1.26

0.18

1.18

0.91

0.16

0.97

0.60

0.12

0.30

1.26

0.68

1.24

0.04

0.66

0.92

1

0.36

0.43

0.79

0.66

1.1

0.77

1.05

1.17

0.93

1.5

0.82

0.34

1.64

0.07

0.13

0.95

P2O5

0.48

0.45

0.48

0.23

0.3

0.16

0.19

0.21

0.19

0.29

0.47

0.47

0.32

0.32

0.50

0.36

0.32

0.49

0.25

0.48

0.36

0.25

0.25

0.26

0.25

0.25

0.2

0.38

0.27

0.21

0.17

0.2

0.17

0.22

0.21

0.16

0.23

0.35

0.41

0.25

LOI

1.88

2.11

2.56

3.35

1.4

3.39

3.72

2.66

2.61

2.92

2.17

1.92

2.75

1.94

3.14

3.37

3.68

3.57

3.12

2.44

3.53

3

2.34

2.36

3.38

7.24

3.29

2.73

2.88

1.99

3.01

2.53

1.74

2.38

2.23

2.84

1.7

4.65

2.46

2.16

Total

99.38

98.06

97.88

97.36

99.8

99.36

98.39

99.79

97.58

98.27

96.94

97.58

99.44

97.52

95.08

97.76

98.81

97.60

97.84

97.62

98.26

95.97

97.16

98.25

97.64

98.03

97.38

98.03

98.45

97.69

98.09

98.07

97.12

97.66

97.81

97.39

98.43

97.56

97.79

98.53

Ba

350

310

270

110

250

280

370

240

640

60

370

510

90

200

200

120

160

450

370

300

40

130

250

290

160

120

190

160

310

210

500

310

220

340

220

90

390

60

110

450

Sr

370

360

350

500

350

270

440

410

340

300

350

410

680

330

290

280

280

190

410

280

150

250

320

280

440

90

320

210

440

310

360

350

350

320

290

240

300

220

270

340

Cs

0.6

0.3

0.3

0.4

0.2

1.3

0.9

0.2

0.7

0.1

0.6

0.5

0.7

0.2

0.4

1.3

0.3

0.1

0.5

0.1

0.1

0.4

0.2

0.2

0.1

0.5

0.4

0.2

1

0.2

0.2

0.5

0.3

0.4

0.2

0.2

0.4

0.2

0.8

0.2

Rb

10.5

10.3

7.1

12.2

13.8

14.9

22.1

18.2

22.7

3

16.9

12.4

3.9

18.0

13.6

3.6

4.1

17.7

13.7

16.6

0.7

13.1

19.8

19.7

5.5

15.4

22.9

14.7

26.7

13.5

26.2

20.6

18

26.6

13

5.8

38.1

1.5

1.1

15

Hf

6

6

5

3

4

2

3

3

3

5

6

6

2

2

4

3

2

5

1

5

4

4

4

4

4

4

3

7

4

3

2

3

2

4

3

3

4

5

5

5

Th

5

5

4

2.3

2.1

1.2

3.3

1.6

1.7

2.3

4.5

3.8

3.1

1.6

4.8

3.1

2.2

5.5

2.3

3.9

2.1

2

2

2.1

1.8

2.4

3

4.8

1.7

1.2

1.1

1.4

1.5

3

1.4

1.5

4.9

3.3

2.6

2.4

U

1.13

1.13

1.04

0.42

0.51

0.31

0.85

0.51

0.3

0.62

1.01

0.92

0.80

0.41

1.14

0.84

0.49

1.24

0.33

0.87

0.34

0.63

0.58

0.59

0.45

0.48

0.79

1.26

0.49

0.32

0.24

0.46

0.35

0.81

0.35

0.38

1.29

0.9

0.62

0.63

Zr

267

276

254

145

184

115

157

153

151

204

231

234

175

163

252

187

178

273

133

235

184

186

177

188

167

173

162

288

176

148

127

136

132

171

152

129

190

235

212

204

Nb

33

34

30

20

19

12

14

14

14

20

28

28

24

16

36

23

20

39

21

29

22

18

17

19

21

22

20

38

20

13

12

13

13

18

14

12

17

24

29

20

Ta

2

2.1

2.1

1.2

1.1

0.7

0.9

0.8

0.9

1.2

1.7

1.6

1.3

0.8

1.9

1.3

1.1

2.1

1.0

1.9

1.4

1

1.1

1.1

1.2

1.3

1.1

2.3

1.1

0.7

0.7

0.7

0.7

1

0.8

0.6

1.1

1.3

1.8

1.2

Ga

22

22

24

20

21

19

19

21

21

23

20

20

25

25

24

26

24

30

20

24

24

22

21

23

21

20

19

24

20

19

20

19

21

22

20

20

21

26

23

22

Y

37.1

37.5

36.1

22.4

27.1

20.7

26

26.8

28.1

34.5

30.3

31.3

28.4

33.4

39.8

32.8

29.9

44.2

23.9

36.6

32.1

30.2

28.7

30.4

27.2

29.5

24.2

37.5

25.6

24.3

22.5

23.5

23.5

28.4

26.1

23.8

31.1

37

34.4

31.3

Ni

24

124

14

41

37

218

55

68

38

70

27

24

61

43

34

44

46

35

49

6

24

44

67

39

31

59

75

42

68

90

75

70

62

36

53

41

27

26

27

40

V

315

318

268

265

264

221

212

253

258

312

270

276

327

348

375

393

312

404

311

311

277

300

274

307

329

269

253

295

284

246

251

246

235

241

267

233

259

341

325

174

Co

36.7

30.3

34.6

37

38

55.3

33.8

41.8

43.9

68

32.0

29.1

53.2

51.5

44.6

44.2

48.0

42.4

56.5

39

35.3

41

40.7

43.2

50.7

37.8

42.6

38

41.6

44.1

40.2

39.7

40.9

37.9

42.5

38.5

32

46.2

37.5

28.8

Zn

170

140

215

206

194

252

176

259

214

221

161

213

276

235

196

151

242

129

197

152

121

146

136

170

168

212

635

775

400

178

231

108

215

147

218

158

182

224

106

211

Cu

35

53

41

56

83

10

62

18

22

65

31

41

159

46

105

86

19

28

14

16

8

71

29

222

61

10

102

84

44

28

78

42

30

69

44

85

82

209

37

56

La

35.8

34.8

32.8

17.7

18.2

12.8

17.3

14.2

14

17.2

28.7

25.4

21.2

15.0

23.0

24.0

17.6

24.2

16.6

30.6

18

19

18.8

20.1

18.2

23.9

22.6

40.9

18.7

14.7

12.7

14.4

14.7

21.3

16.5

13.6

28.1

26

27.9

23.3

Ce

76.5

76.7

71

41.5

42.7

28.4

40.8

32.4

32.2

43.8

60.5

55.5

48.1

35.7

53.6

50.4

40.6

58.1

34.0

66.5

40.4

42.5

43.4

44

43.2

55.1

45.8

86.8

42.7

33.9

29.4

32.9

33.2

47

37.1

30

57.4

58.7

62.2

51.3

Pr

9.79

9.82

9.18

5.39

5.93

3.84

5.28

4.46

4.49

5.97

7.73

7.20

6.59

5.07

7.68

7.06

5.69

8.16

4.59

8.68

5.25

5.86

5.81

6.21

5.98

7.57

5.73

10.8

5.61

4.87

4.05

4.69

4.67

6.28

5.12

4.25

7.44

7.72

8.17

6.9

Nd

37.4

37.1

35.9

20.8

24.3

16

22.2

18.5

18.8

23.5

31.1

29.6

26.8

22.3

32.5

28.9

24.8

34.3

18.6

34.2

21.7

23.8

24.2

25

24.5

30.8

22.2

40.3

23

20

17.1

18.8

18.6

24.3

21.1

17.5

27.7

32.3

33.3

26.8

Sm

8.7

8.5

8.3

4.9

6.3

4

5.5

4.9

5.7

7

6.6

6.4

6.1

5.9

7.8

6.8

6.1

8.0

4.2

8.5

5.7

6.5

6

6.5

6

8

5.6

9.2

5.6

5.1

4.4

5

4.7

6.5

5.5

4.7

6.9

7.7

8.2

7.3

Eu

2.53

2.64

3.12

1.76

2.2

1.47

1.79

1.8

1.74

2.5

2.04

2.10

1.98

2.17

2.23

2.52

2.18

2.41

1.30

2.75

1.88

2.04

2.11

2.19

2.02

2.97

1.66

3.02

1.85

1.8

1.7

1.68

1.76

2.07

1.89

1.7

2.04

2.7

2.78

2.32

Gd

8.8

8.6

8.92

5.78

6.8

4.45

5.81

5.77

6.07

7.19

6.84

6.94

6.01

6.61

8.01

6.95

6.43

8.05

4.45

8.06

6.03

6.31

6.53

6.79

6.11

7.81

5.49

9.32

6.14

5.43

5.21

4.91

5.06

6.21

5.71

5.1

6.84

8.71

8.63

7.96

Tb

1.31

1.38

1.38

0.8

1

0.72

0.89

0.94

0.96

1.18

1.10

1.04

0.91

1.09

1.23

1.11

1.01

1.37

0.69

1.32

1

1.03

1.04

1.11

1.04

1.17

0.85

1.3

0.94

0.86

0.8

0.81

0.78

0.96

0.92

0.76

1.06

1.34

1.29

1.14

Dy

7.65

7.61

7.36

4.55

5.59

4.53

5.46

5.62

5.62

6.87

6.10

6.12

5.38

6.42

7.48

6.38

5.58

8.09

4.33

7.6

6.17

6.15

6.13

6.46

5.79

6.21

4.92

7.56

5.19

5.18

4.55

4.92

4.93

5.71

5.28

5.06

6.3

7.37

7.21

6.93

Ho

1.48

1.49

1.46

0.85

1.07

0.85

1.07

1.03

1.1

1.38

1.23

1.17

1.06

1.25

1.47

1.22

1.11

1.60

0.88

1.48

1.33

1.16

1.13

1.23

1.14

1.08

0.92

1.54

0.98

1.07

0.9

0.93

0.91

1.17

1.07

0.91

1.21

1.47

1.43

1.34

Er

4.09

4.08

3.79

2.56

3

2.19

2.77

2.8

2.85

3.29

3.19

3.27

2.84

3.24

3.84

3.28

2.95

4.36

2.41

3.71

3.71

3.22

3.14

3.3

2.7

2.96

2.46

4.15

2.8

2.65

2.45

2.73

2.31

2.96

2.48

2.4

3.47

3.9

3.84

3.38

Tm

0.55

0.54

0.48

0.32

0.4

0.25

0.38

0.34

0.41

0.48

0.49

0.44

0.39

0.48

0.57

0.47

0.41

0.63

0.33

0.54

0.43

0.4

0.41

0.45

0.38

0.4

0.35

0.56

0.39

0.34

0.33

0.33

0.37

0.45

0.36

0.34

0.5

0.5

0.47

0.52

Yb

3.5

3.4

3

2

2.3

1.7

2.3

2.1

2.4

2.6

2.80

2.80

2.40

2.80

3.30

2.70

2.40

3.80

1.90

3.2

2.9

2.5

2.5

2.7

2.2

2.3

2.1

3.4

2.2

1.9

2.1

2

2.1

2.5

2.2

2.1

2.7

3.1

3.1

2.9

Lu

0.46

0.46

0.45

0.3

0.31

0.26

0.31

0.3

0.33

0.38

0.48

0.42

0.34

0.39

0.53

0.43

0.33

0.53

0.32

0.49

0.43

0.32

0.31

0.34

0.3

0.3

0.34

0.47

0.27

0.32

0.29

0.29

0.29

0.31

0.33

0.28

0.39

0.45

0.46

0.37

∑REE

198.56

197.12

187.14

109.21

120.1

81.46

111.86

95.16

96.67

123.34

158.9

148.4

130.1

108.42

153.24

142.22

117.19

163.6

94.6

177.63

114.93

120.79

121.51

126.38

119.56

150.57

121.02

219.32

116.37

98.12

85.98

94.39

94.38

127.72

105.56

88.7

152.05

161.96

168.98

142.46

The samples have been altered slightly. Thus, trace elements such as Zr, Y, Nb, and rare earth elements are usually considered for petrological study, because they are immobile and appear to remain largely unaffected even during hydrothermal alteration and low-grade metamorphism (e.g., Pearce and Cann 1973; Pearce 1975; Winchester and Floyd 1976; Floyd and Winchester 1978; Wood 1980; Rollinson 1993; Kerrich et al. 1998, Hollanda et al. 2006). Moreover, all major element data are recalculated based on free volatile.

Geochemistry

Soltan Maidan volcanic rocks contain SiO2 contents ranging from 49.0 to 56.9 wt% (average = 52.07 wt%). They have relatively wide ranges of MgO = 2.96–10.52 wt%, Fe2O3(t) = 8.8–15.7 wt%, TiO2 = 1.68–3.51 wt%, P2O5 = 0.17–0.54 wt%, CaO (4.21 and 9.59 wt%) and alkali contents (i.e., K2O + Na2O = 2.55–6.97 wt%).

As shown on SiO2-Nb/Y diagram (Fig. 8), volcanic rocks of Soltan Maidan plot in both subalkaline and alkaline fields. Also, in order to classify all volcanic rocks, Zr/TiO2–Nb/Y classification diagram was used (Fig. 9). The elements used in this diagram are immobile under most hydrothermal conditions, and gives a more robust classification for variably altered samples (Winchester and Floyd 1977). Based on this diagram, Soltan Maidan volcanic samples are plotted in both subalkaline and alkaline basalts regions, mostly near the line separating subalkaline from alkaline basalts (Fig. 9). Position of samples near the subalkaline–alkaline line and also their petrographic composition are indication of transitional to mildly alkaline nature of primary magma.
Fig. 8

SiO2-Nb/Y diagram (after Wood et al. 1979) for Soltan Maidan volcanic rocks. As shown samples plot near the line separating subalkaline–alkaline fields. Symbols are as: (blue open circle) Kaludar Valley, (red open upright triangle) Cheshmeh-Seyed Valley, (green open inverted triangle) Khoshyeilaq area

Fig. 9

Classification of Soltan Maidan volcanic rocks into subalkaline and alkaline on the basis of Zr/TiO2 and Nb/Y contents (after Winchester and Floyd 1977). As shown, samples mostly plot near the line separating subalkaline from alkaline basalts. Symbols as in Fig. 8

The positive correlation of TiO2 with Fe2O3(t) indicate the extraction of titaniferous magnetite and/or ilmenite in the most differentiated terms. Low MgO, Ni (6–218 ppm) and Co (28.8–68 ppm) contents of the studied rocks (Table 1) show low-pressure fractional crystallization and suggest that these rocks are fractionated products which were not in equilibrium with a mantle peridotite source.

Petrographic composition, together with major and trace elements trends on divariant diagrams (Fig. 10) indicate that all of these basaltic rocks originate from a common source and fractional crystallization played main role in controlling the compositional variation and their evolution.
Fig. 10

Variation diagrams of selected major and trace elements vs. Zr for the SMC. Symbols as in Fig. 8

Zirconium is a trace element used as a differentiation indicator for identifying those processes which are responsible for chemical variations in mafic rocks. The negative liner trend of MgO and positive trends of incompatible trace elements (e.g., Nb, Y, and Ta) with increasing Zr, imply progressive removal of small volumes of olivine and/or clinopyroxene. TiO2, Fe2O3(t) and P2O5 are compatible elements for Fe–Ti oxides and apatite. As shown in Fig. 10, their concentrations increase with increasing Zr which might represent fractionation of Fe–Ti oxides and apatite.

Figure 11 illustrates the variation of chondrite-normalized REEs patterns of volcanic rocks of the Kaludar Valley, Cheshmeh-Seyed Valley, and Khoshyeilaq area. They show similar REE patterns with LREE enrichment and weakly fractionated HREE. The similar patterns confirm that all samples might be cogenetic with a common source and related by low-pressure fractional crystallization with each other. This implies that the studied rocks are petrogenetically related and may have been originated via partial melting of the same source and that they underwent similar differentiation processes. The absence of negative Eu-anomaly reveals that plagioclase fractionation was negligible. As low-pressure fractional crystallization is unable to modify LREE/HREE ratios significantly, relatively constant variation and parallel pattern of REEs in these diagrams confirm that all samples are cogenetic and formed from a common source by low-pressure fractional crystallization. A positive correlation between lanthanum contents and concentrations of other incompatible elements suggests that all rocks were likely derived from magmas which came from a similar source, and the observed variations were most likely produced by variation in degrees of partial melting.
Fig. 11

Chondrite-normalized rare earth element plots of Soltan Maidan volcanic rocks in Kaludar Valley, Cheshmeh-Seyed Valley and Khoshyeilaq area. Normalization values after Nakamura (1974)

Variable degrees of partial melting can be explained by variation in highly incompatible to not-so-highly incompatible element ratios (Aldanmaz et al. 2006), as samples representing smaller degrees of partial melting are expected to show higher incompatible element concentrations and highly incompatible to not-so-highly incompatible element ratios. Zirconium and Y are incompatible in the main fractionating phases in basaltic magmas (such as olivine, pyroxene, and plagioclase); Zr/Y ratio is not influenced by moderate degrees of fractional crystallization (Abdel-Rahman 2002, Abdel-Rahman and Nassar 2004). Moreover, the incompatibility of Zr is higher than Y in the mantle phases (e.g., Pearce 1980; Nicholson and Latin 1992); Zr/Y ratio can be sensitive to degree of melting. Thus, the relatively low variations of Zr/Y ratios (Zr/Y = 4.9–7.9 ppm) in the studied samples and a relatively linear trend of Zr vs. Y is mostly due to the role of fractional crystallization. Some scattering in the samples is probably due to variations in partial melting. As shown in Fig. 12, the Zr/Y ratio increase slightly with increasing Zr in the Soltan Maidan volcanic rocks is an indication of their generation by some variable degrees of partial melting (discussed in the following section).
Fig. 12

Variation diagram of Zr/Y vs. Zr for studied samples. Trend showing some variation of partial melting degrees. Symbols as in Fig. 8

One of the essential points is whether the studied samples have suffered crustal contamination. It is known that basaltic rocks influenced by crustal contamination are characterized by La/Nb > 1.5 and La/Ta > 22 (e.g., Hart et al. 1989). Low elemental ratios of these elements in most samples (Fig. 13), suggest that the effect of crustal contamination has probably been negligible during their ascent. In addition, in the primitive mantle-normalized trace element patterns, Soltan Maidan samples display weakly negative anomalies of Nb and (somewhat) Ta (not shown), and also no samples show negative Ta or Nb anomalies on MORB-normalized plots (not shown). Also, because Lu and Yb have similar geochemical behavior, Lu/Yb ratios are not significantly modified by partial melting or fractional crystallization (Dai et al. 2011). Mantle-derived magmas are characterized by low Lu/Yb ratios (0.14–0.15) (Sun and McDonough 1989), whereas continental crust has relatively higher Lu/Yb ratios (0.16–0.18) (Rudnick and Gao 2003). All of the Soltan Maidan volcanic rocks display Lu/Yb ratios in the range of 0.12 to 0.17 (with average 0.14), lower than continental crust, suggesting that they were derived from mantle source, without significant continental crust contamination. However, isotopic data are required to identify the precise role of crustal contamination.
Fig. 13

Diagram of La/Ta ratios vs. La/Nb ratios. Almost all samples plot in the La/Nb < 1.5 and La/Ta < 22 field and showing insignificant crustal contamination. Boundary line values are from Hart et al. (1989). Symbols as in Fig. 8

Wilson (1989) mentioned that if partial melting is fairly extensive (>10 %), the REE should not be fractionated during melting, and therefore ratios of REEs (and spatially very light REE ratios such as La/Ce) should reflect the ratios of the source. In addition, ratios of highly incompatible elements (such as Nb/Ta, Zr/Sm, Ho/Y, and Y/Tb), which don’t vary during fractional crystallization, can be used for diagnosing source composition (Rollinson 1993). As showed in Fig. 14, linear array of incompatible element ratios in binary diagrams illustrate that Soltan Maidan volcanic rocks generated from a common source and fractional crystallization have not generally affected these ratios.
Fig. 14

Variation diagrams of the study area samples based on incompatible-immobile trace elements. Symbols as in Fig. 8

In spite of the presence of a relatively linear correlation between Zr and Nb (Fig. 10), the variations of Nb/Zr ratios (0.088–0.158), cannot possibly be explained by fractional crystallization because the ratio of Zr/Nb is not affected significantly by fractionation of mantle phases (Aldanmaz et al. 2006). Because HREE’s are compatible in garnet and not in spinel (Kuang et al. 2012), REE abundances and their ratios are widely used to determine the degrees and variations of mantle melting (Aldanmaz et al. 2000; Zhao and Zhou 2007). REEs such as La and Yb are particularly useful, because their relative abundances depend strongly on the degree of partial melting and the nature of aluminous phase (garnet or spinel) in the mantle source (Zongfeng et al. 2009; Xia et al. 2012). The Sm/Yb versus La/Yb diagram can be usually used to distinguish the degree of melting and field stability between spinel and garnet peridotite in the source (Lai et al. 2011). Because in the garnet peridotites, Yb is compatible in garnet and La and Sm are incompatible, then La/Yb and Sm/Yb will be strongly fractionated during low degrees of melting (Lai et al. 2011). Moreover, Sm/Yb is nearly unfractionated and La/Yb is only slightly fractionated during the melting in the spinel stability field (White and McKenzie 1995; Yaxley 2000; Xu et al. 2005). Sm/Yb versus La/Yb diagram (after Lai et al. 2011) (Fig. 15) indicates that the parental melt of Soltan Maidan volcanic rocks was probably formed by relatively high degree partial melting, i.e., 14–20 % batch melting of a garnet peridotite mantle source. Furthermore, because the transition from spinel to garnet lherzolite is located at depths of about 70–80 km (e.g., Frey et al. 1991; McKenzie and O’Nions 1991) in the upper mantle, we can assume these depths as the minimum depths of magma generation.
Fig. 15

Sm/Yb vs. La/Yb diagram (after Lai et al. 2011) for the Soltan Maidan volcanic rocks. The samples show a good fit with garnet peridotite stability field and also relatively high degrees of partial melting. Symbols as in Fig. 8

Studied samples have higher Nb (12–39 ppm) and Zr (115–288 ppm) concentrations than those of the N-MORB (Nb = 2.33 ppm, Zr = 74 ppm), and lower than OIB (Nb = 48 ppm, Zr = 280 ppm) (after Sun and McDonough, 1989) suggesting that they could be derived from an enriched mantle source. As ratios of Th/La for the SMC range from 0.08 to 0.23 with an average of 0.13 and Zr/Nb ratios range from 7 to 11.38 with an average of 9.11, between OIB-type magmas (Th/La of 0.11 and Zr/Nb of 5.8) and EMI mantle compositions (Th/La of 0.16 and Zr/Nb of 14.8, Weaver 1991), thus incompatible element ratios for the SMC are mixing between an OIB-like magma with EMI-type mantle source.

The tectonic setting of studied rocks was determined by using a few tectonic discrimination diagrams based on the concentrations of immobile elements. The studied samples are plotted in the within-plate field on the Zr/Y versus Zr diagram (Pearce and Norry 1979) and Ti/100–Zr–Y*3 diagram (Pearce and Cann 1973; Fig. 16a, b). This also confirms by plotting samples on the DF1-DF2 diagrams (after Agrawal et al. 2008; Fig. 16c, d). As shown in the mentioned diagrams, all of the samples indicate a within-plate (continental rift) tectonic setting.
Fig. 16

Tectonomagmatic discrimination diagrams for Soltan Maidan volcanic samples. These diagrams showing a within-plate and continental rift setting for the studied samples. a Zr/Y–Zr diagram is after Pearce and Norry (1979), b Ti–Zr–Y diagram is after Pearce and Cann (1973), c and d La, Sm, Yb, Nb, and Th diagram is after Agrawal et al. (2008). Symbols as in Fig. 8. MORB mid-ocean ridge basalt, IAB island arc basalt, CRB continental rift basalt, OIB ocean–island basalt

Geochemistry of the Soltan Maidan volcanic rocks suggests that the primary magma was derived by relatively high degrees of (14–20 %) partial melting of an enriched mantle source in the garnet stability field. They also suffered minor fractional crystallization and crustal contamination. In addition, the volcanic activity has occurred in a continental rift geotectonic setting.

Discussion

SMC lies unconformably on the Ghelli Formation and is covered by the Padeha Formation (Fig. 3). This complex was accumulated in both subaerial and submarine environments. The marine environment is supported by pillow lavas and intercalated green shale-siltstone layers (Fig. 3). However, columnar joints, red sandstones, and existence of intercalated granitic conglomerates in different outcrops of the SMC show some parts of it, formed in a subaerial environment (Fig. 3). These evidences show instability and vertical movements of the study area at this time. Because of high volumes of magmatism in the SMC, it is possible that surface topography has been affected by outpourings of the voluminous lava flows.

Ghavidel-Syooki et al. (2011) suggested a relatively shallow marine and platform depositional environment for the Upper Ordovician shale beds of the SMC and Ghelli Formations. They pointed out that Alborz zone was located in the northern continental margin of Gondwanaland during Ordovician and Silurian times. Aharipour et al. (2010) also showed that Padeha Formation was deposited in a terrestrial environment during Early–Middle Devonian and its sedimentary fill pattern is similar to an intracratonic rift basin. The shallow marine depositional environment of Ghelli Formation and continental environment of Padeha Formation deposition are all indications of an uplift event during Late Ordovician to Middle Devonian in the area.

The lava flows of the westernmost outcrop (i.e., Kaludar Valley) show columnar joints without any pillow structure or marine sedimentary interlayers (Figs. 3 and 4b). Furthermore, less general thickness of lava flows here indicate that during Late Ordovician to Silurian times the western parts of the study area might be located in the higher elevations in comparison to eastern parts. Indeed, evidences indicate an eastward tilting of the whole area.

Granitic rocks of the continental high elevations were eroded and transported to the local lava flows accumulation basin in Cheshmeh-Seyed Valley and Khoshyeilaq area. The presence of a polygenic conglomerate with enormous large granitic clasts in the upper parts of the SMC in the Cheshmeh-Seyed Valley and Khoshyeilaq area provide a good signature for uplifting and subsequent eroding of the elevations by unroofing these granites during accumulation of SMC. This indicates that: (1) the study area was tectonically very active and the normal listric faults controlled the depth of this syn-rift basin and, (2) the development of vertical movements in this part of the Alborz zone during Late Ordovician to Silurian, probably related to formation of paleotethys rift basin at this time. These activities might also cause high partial melting of upwelling mantle in the Soltan Maidan area. The position of the studied samples in the 14–20 % partial melting field in Fig. 15 is an indication of this. However, the transitional to alkaline nature of samples also shows intense crustal thinning of the study area. These characteristics could be considered as the evidence of maximum crustal thinning and beginning of the break-off of Alborz continent at this time. Continental rift systems are frequently associated with abundant volcanism and consequently, the thinning of the continental lithosphere during the rifting process is believed to be an important cause of magma formation (Sengör and Burke 1978; McKenzie and Bickle 1988). According to Wilson (1989), continental basalts are dominantly alkaline in the early stages of continental rifting; however, in regions which large amounts of crustal extension have occurred, transitional and tholeiitic types may be common.

It is probable that the upwelling mantle plume(s) was/were responsible for exhumation and crustal thinning in this part of Alborz zone and possibly some variation of partial melting (as discussed in the “Analytical methods” section) happened by decompression partial melting of ascending mantle.

During the Paleozoic, the Iranian and Arabian plates formed a coherent unit and were separated from the Turan plate by Paleotethys (Karimpour et al. 2010). According to Alavi (1996), the Alborz continental crust was a site of development of a very shallow epicontinental sea in the time interval between the latest Precambrian to Ordovician. From Late Ordovician to Late Devonian, the Iranian terranes, including Alborz, Central Iran, Sanandaj-Sirjan, and Zagros had peri-Gondwana/Gondwana position and located at 15–30°S palaeolatitude (Cocks and Torsvik 2002). Alavi (1996) pointed out that Ordovician was the time of initiation of extension and breakup of Alborz continent, and rifting event lasted up to Early–Middle Devonian times. Also, the unusually long Ordovician to Early–Middle Devonian magmatism in different parts of Iran is interpreted by many authors as rift-related magmatic events (e.g., Jenny 1977; Stampfli 1978; Berberian and King 1981; Boulin 1991; Stampfli et al. 1991; Alavi 1996; Lasemi 2001; Bagheri and Stamppfli 2008; Ghasemi and Derakhshi 2008; Aharipour et al. 2010). The tectonic evolution of this rift system led to the formation of the Paleotethys Ocean in north of Iran. The Paleotethys Ocean opened in Silurian (Karimpour et al. 2010) and Central Iran zone was located along the passive margin coast of the Paleotethys during this time (Nowrouzi et al. 2013). Paleotethys life ended when the Alborz continent lithosphere collided with the Turan plate along a north-dipping subduction zone in the late Triassic time (Davoudzadeh and Schmidt 1984; Stampfli and Pillevuit 1993; Alavi 1996; Stampfli 1996; Karimpour et al. 2010). The subduction of Paleotethys Ocean under the Turan plate also started in the Late Devonian (Karimpour et al. 2010).

The results of our investigation indicate apparent synchrony of the extension and high voluminous magma generation in a continental rift setting during Late Ordovician to Silurian. It was also found that Soltan Maidan volcanism is the most important and voluminous magmatic event in the eastern Alborz zone that is related to the initiation of Paleotethys development.

Conclusions

The combination of field, petrological and geochemical studies together with tectonic constraints have enabled us to propose a model for magma genesis of the Soltan Maidan Complex. The main aspects of this model are briefly summarized as follows:
  1. 1.

    Upper Ordovician to Middle Devonian magmatism in the form of mafic dykes, sills and basaltic lava flows is widespread in the eastern part of Alborz zone, northern Iran. SMC with Late Ordovician–Silurian age is the main phase of magmatic activity in the study area. It consists of 436 to 1,286 m thick basaltic lava flows, agglomerate and tuff accompanied with some sedimentary intercalations.

     
  2. 2.

    Soltan Maidan volcanic rocks indicate that the magma is transitional to mildly alkaline in nature, derived by fairly high degrees, 14–20 %, partial melting of an enriched mantle source in the garnet stability field. These volcanic rocks are related to a common basaltic magma, which was evolved via fractional crystallization and insignificant crustal contamination. The magmas underwent fractional crystallization, possibly in shallow magma chambers. The magma chambers were then tapped, and the magmas erupted at surface.

     
  3. 3.

    Field evidence show that SMC accumulated in both subaerial and submarine environments, and the western parts of the study area were located in the higher elevations in comparison to eastern parts during formation of this complex, which indicates an eastward tilting of the whole area. Examination of rock successions in the study area show a long period of uplifting event, accompanied by vertical movements which lasted from Upper Ordovician to Middle Devonian.

     
  4. 4.

    Study of Soltan Maidan volcanic rocks indicate that these rocks were generated from partial melting of an enriched mantle source in a continental within-plate geotectonic setting. In addition, incompatible element ratios for the SMC are indications of mixing between an OIB-like magma with an EMI-type mantle source.

     
  5. 5.

    The results also show that magmatic activity which led to the formation of the SMC in the eastern Alborz zone is relevant to a cycle of rifting which were happened during the opening of the Paleotethys Ocean.

     

Notes

Acknowledgments

This paper is part of the PhD thesis by M.D. at Shahrood University. This study was supported by the Shahrood University and Iran National Science Foundation, Presidential Office, Deputy of Science and Technology, Grant no. 90004893. We thank all of them for financial supports. We are also thankful Dr. G. H. Kazemi and Dr. M. Rezaei from Shahrood University and Dr. F. T. Koksal from Middle East Technical University, Turkey, for their valuable comments and suggestions on manuscript. Journal editors are thanked for handling the manuscript.

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Copyright information

© Saudi Society for Geosciences 2013

Authors and Affiliations

  1. 1.Department of GeosciencesShahrood UniversityShahroodIran

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