Middle Permian (Guadalupian) fusulinid taxonomy and biostratigraphy of the mid-latitude Dalan Basin, Zagros, Iran and their applications in paleoclimate dynamics and paleogeography

Carbonates in the Permian–Triassic Dalan and Kangan formations in the Zagros region in Iran and equivalents in the Arabian Plate are one of the major hydrocarbon reservoirs in the Middle East. Although numerous studies of Permian sequences were performed in the Zagros area (Szabo and Keradpir, 1978; Al-Jallal, 1987; Alsharhan and Nairn, 1995; Sharland et al., 2001; Insalaco et al., 2006) they were mostly focused on sedimentology, tectonics, sequence-stratigraphic and palaeoecological studies of the Upper Dalan Member and Kangan Formation (Upper Permian–Triassic) reservoirs in the Iranian offshore Fars area and their equivalent outcrops in the Zagros region. The biostratigraphy in these studies was quite limited and the major part of the Permian marine succession below the Upper Dalan Member was poorly studied. Recent foraminiferal biostratigraphic work in the area (Gaillot and Vachard, 2007; Kolodka et al., 2012) also dealt, in large part, with the Upper Permian (Lopingian) Upper Dalan (Khuff) Member and the lack of non-oriented fusulinids available from the Lower Dalan Member prevented a proper biostratigraphic documentation of this part of the succession. Overall, the biostratigraphic, chronostratigraphic and paleogeographic studies of the Middle Permian (Guadalupian) Lower Dalan Member in the Zagros area remained quite poor.

The Zagros area itself possesses a unique mid-latitude paleogeographic position within the transition from the margins of Gondwana to Neo-Tethys (Sharland et al., 200l), where climatic changes occur with considerable frequency and amplitude (Burrows et al., 2011). Thus, careful sedimentological and paleontological studies may significantly contribute to our knowledge and understanding of the paleogeography and paleoclimate of this region. The objectives of this paper are (Figures 1 and 2): (1) summarize the stratigraphy of the Dalan Formation in the Il-e Beyk Section; (2) analyze microfacies and paleoenvironments emphasizing bathymetry; (3) present a comprehensive taxonomic study of fusulinids and smaller foraminifers (shown in taxonomic order in Plates 1–7); (4) provide a better biostratigraphic and chronostratigraphic age of the Dalan Formation based on fusulinids; and (5) improve the paleogeographic and paleoclimate reconstruction of Zagros and surrounding areas in Middle Permian (Guadalupian).

Permian–Triassic deposits in Zard-Kuh area (Figure 1) were preliminarily studied as the Khuff Formation equivalent (Sampo, 1969). The lithology and thickness of Permian strata in Zard-Kuh and Kuh-e Dena are rather different. In Zard-Kuh, Permian deposits are 932 m thick and consist of lower limestone and upper dolomite units whereas in Kuh-e Dena a middle sandstone unit rests between lower limestone and upper dolomites and the Permian sediments decrease in thickness to 332 m. Setudehnia (1975) considered the lower fossiliferous limestone to be Middle to Late Permian in age on the basis of fossil content (small foraminifers, few fusulinids, corals and algae). Szabo and Kheradpir (1978) studied the Permian–Triassic deposits in Kuh-e Surmeh and Kuh-e Siah No. 1 Well in the Fars Province and Kuh-e Faraghan in Bandar Abbas, naming the Permian carbonate and evaporite as the Dalan Formation. They subdivided it into three members: (1) Lower Dalan Member (carbonate); (2) predominantly anhydrite middle Nar Member; and (3) Upper Dalan Member consisting of anhydrite interbedded with carbonate.

Two assemblage zones (A and B) contain small foraminifers, fusulinids and algae and were recognized in the lower and upper members of the Dalan Formation. Based on the fossil content, a Middle Permian (Guadalupian) to early Wuchiapingian (early Late Permian) age was proposed for the formation. Zaninetti et al. (1978) studied Permian deposits in Kuh-e Gakhum, where the predominant lithologies of Midian (late Wordian) to early Wuchiapingian deposits are dolomitic with some interbeds of limestones especially in upper parts. Rosen (1979) revised the stratigraphy of Permian–Triassic deposits of the Fars-Gulf area and suggested that the Wuchiapingian fauna in the Upper Dalan and the Scythian (Early Triassic) fauna in the basal beds of the Kangan indicate no major erosional cycle. Moreover, he suggested that only a short period of cessation of carbonate deposition had occurred.

Johnson (1981) remeasured and resampled the Dalan Formation in Kuh-e Surmeh and studied this formation biostratigraphically. He recovered Shanita and several other smaller foraminifers in the Nar Member and assigned the formation as late Murgabian. Baghbani (1997) reported some fusulinids and small foraminifer assemblages in outcrops of the Dalan Formation in Kuh-e Surmeh, Kuh-e Dena and Qali Kuh and assigned middle Kubergandian to Dorashamian for the Dalan Formation. Additional biostratigraphic studies carried out by Insalaco et al. (2006) show the age to be Wuchiapingian–Changhsingian due to the foraminiferal content of the Upper Dalan Member in the Surmeh and Dena sections. Gaillot and Vachard (2007) provided a detailed biostratigraphic and microfacies study of the late Middle Permian to Early Triassic Khuff Formation, undertaken in three reference regions in the Zagros area (Iran), United Arabian Emirates (UAE) and Saudi Arabia.

Recently, Kolodka et al. (2012) published new results from the Dalan Formation of southeastern Zagros, where details of depositional environments of the entire formation and new biostratigraphic biozonation were described. The latter, however, is difficult to evaluate as the taxonomy in the paper is based on non-oriented and/or poorly preserved specimens. The latter are highly interpretive and some of the “species”, in our opinion, cannot be determined even at generic level (for example “Chusenella cf. conycocylindrica”, fig. 9R; or “Ch. cf. brevipola”, fig. 10C etc.). The morphology of the new genus Praedunbarula, in our opinion, is perfectly consistent with long-ranging Schubertella, indicating that the proposed age of the Praedunbarula Biozone is therefore questionable. Besides, in the genus description (in appendix by D. Vachard) the range of this genus is reported as Kungurian-Changhsingian, whereas in biostratigraphy (Kolodka et al., 2012) the age of the Praedunbarula Biozone established as late Wordian (?). We thus consider the proposed “zonation” to be mainly ecostratigraphic.

Weidlich and Bernecker (2007) and Koehrer et al. (2010) introduced integrated facial, stratigraphic and biostraigraphic framework from the Saiq Formation in Oman. Unfortunately, no details of biostratigraphy are provided and both papers are lacking of details of taxonomy. Forke et al. (2012) published a short abstract with some details of smaller foraminifera and fusulinids distribution within the Saiq Formation in the Jabal al-Akhdar area integrated with sequence-stratigraphic framework. The precise correlation with Tethyan Scale is also provided in the work.

The Zagros orogenic belt of Iran, as part of the Alpine-Himalayan mountain chain, extends for about 2,000 km in a NW-SE direction from the East Anatolian Fault of eastern Turkey to the Oman Line in southern Iran (Alavi, 1994). The orogen is a product of Early Pliocene closure of the Neo-Tethys Ocean (Stöcklin, 1968; Hessami et al., 2001). This orogen consists of the Urumieh-Dokhtar Magmatic Arc (Stöcklin, 1968, 1974; Berberian et al., 1982; Alavi, 2007), the Sanandaj-Sirjan Zone (Stöcklin, 1968; Berberian et al., 1982; Alavi, 2007) as the internal part, and Zagros Fold-and-Thrust Belt (ZFTB) (the Zagros “simple folded zone” of Falcon, 1974) as the external part (Alavi, 2007) (Figure 1). The ZFTB can be divided into two structural domains: (1) an imbricate thrust belt or the High Zagros Belt (HZB), and (2) the Simply Folded Belt (SFB) that are separated by the High Zagros Fault (Berberian, 1995).

The studied area is located in the central part of the High Zagros Belt. One of the main features of the High Zagros Belt is the numerous thrust fault complex (Figure 1). Zard-Kuh-e Bakhtiari also includes the transitional zone from the internal (Sanandaj-Sirjan Zone) to the external parts (ZFTB) of the Zagros orogen, and is structurally characterized by several subhorizontal or gently NE-dipping and SW-verging thrusts (Alavi, 2007) (Figure 1). One of the main thrusts in the Zard-Kuh-e Bakhtiari is the Zard-Kuh thrust that has juxtaposed lowermost Paleozoic rock units with the Pleistocene Bakhtiari conglomerates (Alavi, 2007). A number of NW-trending, NE-dipping thrust sheets are the major structures of the Bakhtiari area (Nemati and Yassaghi, 2010).

The Il-e Beyk Section is located on the Zard-Kuh Mountain (270 km west of Shahr-e Kurd city, Figure 1). In this section, the Lower Permian Faraghan Formation is missing and the Dalan Formation disconformably overlies the Ordovician shale of the Zard-Kuh Formation (Figure 2,Table 1). In other Dalan Formation outcrops within the Zagros area (in Tang-e Putak, Kuh-e Dena and Chal-I-Sheh) there is sharp contact between the red and gray shales interbedded with yellow weathering sandstones of the Faraghan Formation and the dark gray, sandy fossiliferous limestones of the Dalan Formation. The siliciclastics of the Faraghan Formation replaced by the carbonates of the Dalan Formation are also in sharp contact. The Dalan Formation is overlain by apparently conformable Triassic dolomites of the Kangan Formation (Insalaco et al., 2006). The Permian/Triassic boundary is not well-recognizable in the Il-e Beyk section (Setudehnia, 1975) since the upper part of the Dalan Formation is composed of dolomite. Because the Triassic Kangan Formation has the same lithology it is difficult to distinguish the clear boundary between the Dalan and Kangan formations. But the changing of bedding from massive dolomitic limestones to thin- to medium-bedded dolomites of the Kangan Formation is clear in the field. Conventionally we considered this boundary just above the sample Z-1 with the last occurrence of the Permian foraminifers (Figure 2).

The Dalan Formation in the section is 927 m thick and can generally be divided into two main parts (Figure 2). The lower part (484.5 m thick, including units 1–18) is composed mainly of gray to light brown, medium- to very thick-bedded, cliff-forming packstones and grainstones with rare horizons of mudstones. The middle and upper portions of this succession contain several fusulinid horizons (Table 2). The limestones within this lower part of the section also include brachiopods, corals, algae, bryozoans, gastropods and crinoid stems. The upper part of the Dalan Formation (442.5 m thick, including units 19–28, Figure 2) is predominantly made of cream-colored, cliff-making, medium- to coarse-grained dolomite with occasional chert nodules interbedded with dolomitic limestones, yielding one fusulinid horizon and remains of brachiopods, algae, crinoid stems, gastropods and small foraminifers.

The total thickness and lithology of the Dalan Formation in the High Zagros varies. It is thickest in the Zard-Kuh area (Il-e Beyk section) and thinnest in Kuh-e Dena (Setudehnia, 1975; Parvizi et al., 2008). In the latter area, a section of middle sandstone adds to the Dalan sequence. It should be noted that the well-developed evaporite deposits within the middle Nar Member of the Dalan Formation in the Zagros simply folded belt in the type section of this formation at Kuh-e Surmeh (Szabo and Kheradpir, 1978), are absent in the studied area.

Forty-six samples from the entire Dalan Formation in the Il-e Beyk section were studied and 66 taxa of fusulinid and smaller foraminifers were recognized (Table 2). About 15% of these taxa are long-ranging smaller foraminifers. The fusulinids are unevenly distributed throughout the succession and were recovered from six stratigraphically restricted horizons (Figure 2). Taking into account the time-span of the entire Dalan Formation (Roadian through the top of the Permian, i.e. about 20 Myr) this characterization is inadequate. Nevertheless, at the moment it is the most comprehensive record of fusulinid biostratigraphy of the Lower Dalan Member in the entire Zagros.

In this paper we are using Tethyan chronostratigraphic terms rather than the global because of complexity in correlation with the Global Time Scale (Figure 3). Our correlation with the latter is the same as in the Geological Time Scale 2004 (Wardlaw et al., 2005), but different from what was proposed in the recently published Geological Time Scale 2012 (Henderson et al., 2012). It should be noted that the correlation between the Global and Tethyan scales in GTS 2012 is shown as provisional (dashed lines) and it was a compromise involving the co-authors between radical suggestions of Henderson and Mei (2003) and data from fusulinids and ammonoids (Chedija et al., 1986; Leven, 1992; Davydov, 1996). Recently, however, the differences between these two correlation models, particularly in Oman region were resolved (Kozur and Wardlaw, 2010; Baud et al., 2012), but still exist in South China (Shen et al., 2007) and thus in the Permian Chapter of GTS 2012.

Five biozones can be distinguished in the studied succession. The lower boundary of each zone can be identified by the presence of the characteristic assemblage. The upper boundary is sometimes conventional and placed at the base of succeeding zone.

The first biozone, Praeskinnerella parviflucta - Chusenella? pseudocompacta, is recognized in the lower 120 m of the succession (Figure 2, Table 2). The assemblage is relatively poor, possessing three species of fusulinids and a few smaller foraminifers. Neofusulinella lantenoisi is a widely distributed fusulinid in the Kubergandian–Murgabian (early Roadian to early Wordian) of the entire Tethys (Deprat, 1913; Toriyama, 1975; Kahler and Kahler, 1979; Leven, 1998, 1997; Xiao et al., 1986; Huang et al., 2009). Praeskinnerella parviflucta was originally described from the Chihsia Formation (Yakhtashian–Kubergandian) in South China (Zhou, 1982) and later recovered in Bolorian (Kungurian) of Darvaz (Leven, 1992) and Kubergandian (early Roadian) in Transcaucasia (Leven, 1998). Chusenella? pseudocompacta is widely known in the lower-middle Maokou in South China (Sheng, 1956, 1963; Zhang and Dong, 1986) and in recent times has been reported in upper Middle Permian (Guadalupian) in the Lhasa Block of Peri-Gondwanan affinity (Zhang et al., 2010). We propose a Kubergandian (early Roadian) age for the Praeskinnerella parviflucta-Chusenella? pseudocompacta Biozone, because of the combination of the Bolorian–Kubergandian (Kungurian–early Roadian) Praeskinnerella parviflucta and the Kubergandian–Midian (early Roadian–Capitanian) Chusenella? pseudocompacta. We also suggest the Kubergandian (early Roadian) age since the biozone’s stratigraphic position is below the beds we assigned to the Murgabian (late Roadian–early Wordian) (Figure 2, Table 2).

The second Eopolydiexodina persica Biozone is established from 120 through 320 meters in the section. Although the upper 90 m of the biozone contain only long-ranging smaller foraminifers and one specimen of Yangchienia iniqua (Figures 2, Table 2), the fusulinids here are more diverse than that in the lower biozone. It includes several long-ranging smaller foraminifera (Table 2) and 12 fusulinid species: Schubertella rara, Schubertella aff. rara, Minojapanella sp., Yangchienia iniqua, Yangchienia sp., Chusenella conicocylindrica, Chusenella urulungaensis, Chusenella dorashamensis, Chusenella aff. tieni, Skinnerella elliptica, Rugososchwagerina ex gr. xanzensis, Eopolydiexodina persica, Eopolydiexodina aff. persica. The most characteristic species in this assemblage is Eopolydiexodina persica, as originally described from an unidentified horizon in Zagros (Kahler, 1933). The species was also reported in Turkey (Köylüoglu and Altiner, 1989), Abadeh (Kobayashi and Ishii, 2003a, b) and Transcaucasia (Kotlyar et al., 1989; Leven, 1998) from deposits correlated with lower–middle Murgabian (upper Roadian–lower Wordian). Eopolydiexodina in all these regions occurs as a dominant faunal element in a relatively narrow horizon (Douglas, 1936; Leven, 1998; Kobayashi and Ishii, 2003a).

Similarly, in Il-e Beyk section two Eopolydiexodina-bearing horizons are 1 –2 m thick. The genus ranges just within 25 m (Figure 2) in the section. The other characteristic species of this biozone are Yangchienia iniqua and several species of Chusenella (Table 2). Yangchienia iniqua has been described from the Maokou Formation in South China (Lee, 1927) and in the Southeast Pamirs. It is known to occur in upper Kubergandian (early Roadian) and lower–middle Murgabian (upper Roadian–lower Wordian) (Leven, 1967; Chedija et al., 1986). The Chusenella species were described from South China and Tibet (Chen, 1956; Wang et al., 1981) and in Transcaucasia (Rosovskaya and Rauser-Chernousova, 1965). In the latter region, these species occur in the upper Asni and lower Gnishik formations that correlate with lower–middle Murgabian (upper Roadian–lower Wordian) (Rauser-Chernousova et al., 1974; Kotlyar et al., 1989; Leven, 1998). The other indicative species are Skinnerella elliptica and Rugososchwagerina ex gr. xanzensis. Although the latter is represented by an atypical form, it is quite characteristic for the lower–middle Murgabian (upper Roadian–lower Wordian) Eopolydiexodina persica Biozone in the Abadeh Region (Kobayashi and Ishii, 2003b) and occurs in the lower Gnishik Formation in Transcaucasia (first author’s personal data). Overall, the age of this biozone is lower–middle Murgabian of the Tethyan Scale or upper Roadian–lower Wordian of the Global Scale (Davydov, 1996; Wardlaw et al., 2005).

The third biozone, Chusenella schwagerinaeformis-Rugososchwagerina zagroensis occupies a relatively narrow horizon within 320–350 m of the succession (Figure 2). The assemblage there is quite diverse. Minojapanella sp. ranged from the previous biozone. The dominant forms are diverse Chusenella (10 species), among which some first appear in the previous biozone and range into the described one. The other species, such as Yangchienia iniqua, Ch. minuta, Ch. solida, Ch. cheni, Ch. sinensis, first occur in this assemblage. Ch. sinensis was originally described from the Maokou Formation in South China (Sheng, 1963), are widely distributed in the analogues of Murgabian to Midian (late Roadian–Capitanian) deposits in Tethys (Sun and Zhang, 1988; Leven, 1998; Kobayashi and Ishii, 2003a, b). Ch. minuta, and Ch. solida were described from upper Midian (Capitanian) of Central Turkey in association with Kahlerina, Dunbarula, an advanced Neoschwagerina and Yabeina (Skinner, 1969).

Also quite abundant in this biozone, Schubertella karasawensis was described in Japan from the upper Nabeyama Formation in association with Reichelina, Dunbarula, advanced Yangchienia, Rauserella, Neoschwagerina, Pseudodoliolina ozawai, and Codonofusiella (Kobayashi, 2006). The formation is of Midian (upper Wordian-Capitanian) age. Two species of Rugososchwagerina that were also recovered in this assemblage closely resemble R. altimurica Leven. The latter species was designated in Afghanistan from undoubtedly Midian (upper Wordian–Capitanian) deposits (Leven, 1997). Rare Codonofusiella also occurring in this biozone have been reported from Middle Permian (Guadalupian) of Transcaucasia (Leven, 1998). Common, yet primitive Dunbarula and a single specimen of Kahlerina ussurica clearly suggest that upper part of this biozone correlates with the Arpa Formation in Transcaucasia of early Midian (late Wordian) age (Kotlyar et al., 1989; Davydov, 1996; Leven, 1998) in the Tethyan Scale. We therefore correlate this biozone with the upper Murgabian and lower Midian of the Tethyan Scale, which is equivalent to the upper Wordian of the Global Scale (Davydov, 1996; Wardlaw et al., 2005).

Kobayashi and Altiner (2011) reported “Dunbarula” from the lower Murgabian (upper Roadian) in Turkey, though in our opinion this form does not belong to the genus. The initial stage of the development of these forms (skewed axis) clearly suggests that they are schubertellids rather than ozawainellids (Dunbarula). Dunbarula and Kahlerina were also reported from Neoschwagerina occidentalis Biozone in Abadeh (Kobayashi and Ishii, 2003a). However, they were neither described nor figured and thus the evaluation of these forms is not possible. Taking into account that Kobayashi (Kobayashi and Altiner, 2011) include some clear schubertellids into Dunbarula we are avoiding using this data.

The fourth upper Middle Permian (Guadalupian) fusulinid biozone Monodiexodina kattaensis in Il-e Beyk section was recovered from 360 to 435 m in the succession. Fusulinids are quite rare throughout the biozone, except for one thin (1–1.5 m) horizon where fusulinids are reasonably diverse (Figure 2). Chusenella cheni and Yangchienia hainanica occur in the lower part of the biozone and ranged from the previous biozone. Yangchienia haydeni, found throughout this biozone, was described from Afghanistan in association with Afghanella schencki (Thompson, 1946). The former species is widely distributed in the upper Murgabian to Midian (Wordian–Capitanian) of the western Tethys (Leven, 1997). Codonofusiella erki, originally described from lower Wuchiapingian in Transcaucasia (Rosovskaya and Rauser-Chernousova, 1965), is also known to occur in the Midian (upper Wordian–Capitanian) of South Afghanistan (Leven 1997) and Turkey (Erk, 1942).

Dunbarula tumida and Pseudodunbarula erki were described from the Midian (upper Wordian–Capitanian) and possibly Wuchiapingian of Turkey (Skinner, 1969); they have also been reported from upper Midian (Capitanian) of Tunisia (first author’s personal data) and Afghanistan (Leven, 1997). The latter species closely resembles Pseudodunbarula arpaensis from the lower Wuchiapingian of Transcaucasia (Kotlyar et al., 1983) but is definitely a more primitive representative of the genus. Kahlerina is common in sample Z33. This genus is quite characteristic of the Midian (upper Wordian–Capitanian) in the entire Tethys (Sosnina, 1968; Sheng, 1963; Leven, 1997, 1998; Han, 1980; Kobayashi, 2006, 2009). It is especially abundant in the upper Midian (Capitanian) and occurs only at this level in the type area of the Capitanian (Nestell and Nestell, 2006). The most intriguing aspect of this biozone is the occurrence of Monodiexodina kattaensis (new subspecies), the most abundant fossil in bed Z33 (Table 2). The species has been described from the Peri-Gondwanan cool-water Amb Formation in the Salt Range, Pakistan (Schwager, 1887; Dunbar, 1933; Douglass, 1970) together with Codonofusiella laxa. According to the conodonts, the age of this formation is Wordian (Wardlaw and Pogue, 1995).

The fifth and uppermost Middle Permian (Guadalupian) Afghanella schencki Biozone extends from 438 through 687 m, although the lower 200 m of this biozone possess only rare smaller foraminifers (Table 2). Afghanella schencki and Sumatrina brevis recovered in this biozone are the only verbeekinids from this section (Plate 6, Figures 6–8). The first species is widely known from the upper Murgabian through the Midian (Wordian–Capitanian) throughout the Tethys (Thompson, 1946; Sheng, 1963; Xiao et al., 1986; Leven, 1997; Kobayashi, 2003). However, in Zagros Afghanella possesses a higher position in the upper Midian, i.e. Capitanian (Figures 2 and 3). Sumatrina brevis has been described in Southeast Pamirs from the upper Gan Formation (Leven, 1967) but has never been recovered elsewhere.

The foraminifers in the succession above 687 m in Il-e Beyk section are quite rare due to strong dolomitization (Figure 2). We recovered a very poorly preserved assemblage that includes Paraglobivalvulina sp. and Dagmarita sp. in the uppermost part of the Dalan Formation, suggesting a Late Permian (Lopingian) age for at least the upper 30 m of the section. The interval from 687 through 920 m can, at the moment, be assigned either a Middle Permian (Guadalupian) or Late Permian (Lopingian) age. The entire Dalan Formation (952 m thick) and the Lower Dalan Member in this section (lower 687 m) are the thickest within entire Zagros.

In summary, the lower 120 m in Il-e Beyk Section belong to the Kubergandian (early Roadian) Stage and correlates with the Asni Formation of Transcaucasia; the succession from 120 through 320 m correlates with the Gnishik Formation of Transcaucasia and can be assigned to the lower–middle Murgabian (upper Roadian–lower Wordian); the beds from 320 through 350 m correlate with upper Gznishik and lower Arpa formations of Transcaucasia and can be assigned to the upper Murgabian and lower Midian (Wordian) of the Tethyan Scale. The sequence from 350 through 687 m, in our opinion, correlates with Khachik Formation of Transcaucasia and can be assigned to the upper Midian Stage (Capitanian) (Figure 3).

In order to better understand the environments and bathymetry of Middle Permian (Guadalupian) strata in the studied section, detailed petrographic analysis and microfacies examination of the Dalan carbonates in the Il-e Beyk Section were performed which led to record eight facies types and their subtypes. A total of 180 thin sections were prepared from collected samples in the field and studied for microfacies analyses.

This facies is composed of thin to medium-bedded, dark gray beds. Quartz grains and skeletal fragments are absent or represent minor components. No bioturbation is visible. This dolomudstone has no evidence of subaerial exposure and shows that this kind of mudstones formed under subtidal condition. The rare skeletal grains (crinoid stems, ostracodes, brachiopods) and their low diversity indicate a landward position of this facies. This facies is present in the middle part of unit 4, unit 9, the lowermost portion of unit 12 and the lower part of unit 15 (Figure 4a).

Textural changes, internal structure, and crystal shape indicate two kinds of dolostones in the Dalan Formation. These include very fine dolostones, or, dolomicrite, and medium crystalline dolostones, or, dolosparite. Dolomicrites yield no porosity and demonstrate xenotopic fabric. The dolosparites of the Dalan Formation do not retain any of the the primary sedimentary texture due to pervasive dolomitization. The dolosparite is formed at very high temperatures and replaced the limestones (Gregg and Shelton, 1990). The preservasion of the original sedimentary structures such as fenestral fabric, evaporite molds and relics of primary allochems (including crinoid stems, brachiopods, small forams) in the dolomicrites of the Dalan Formation suggest that these dolomites were formed under near-surface low temperature conditions (Gregg and Shelton, 1990) and in arid upper intertidal settings (Shinn 1986). This facies was encountered in units 21, 25 and 27 (Figures 4b, 5a).

This facies is typified by light to dark gray, thin to medium thick beds of skeletal debris cemented mostly by micrite. Sedimentary structures include low-angle to wavy lamination and muddy (bioturbated) tops. Sorting is generally poor. Sparse skeletal fragments include crinoid stems, smaller foraminifera, echinoderms, ostracodes and calcareous algae. Minor components include fenestrate bryozoa fronds, ostracodes and rare staffellid foraminifera, forming between 10% to 15% of the rock. Quartz is a minor constituent (< 3%). This facies occurs in association with peloids (about 10%) and defines the peloidal bioclastic wackestone. Dolomitization has locally caused micrite to be partially altered to dolomicrite to dolosparite and therefore form dolo-bioclastic wackestones. The bioclastic wackestone indicates a restricted shelf region. The micritic groundmass and the low diversity of skeletal grains suggest deposition under low energy, very restricted lagoon. In some cases this facies is burrowed. Partial micritization of skeletal fragments indicate that the deposition occurred in a quiet water and lagoonal environment (Nichols, 1999; Flügel, 2010). Rare wavy laminations at the top suggest the influence of locally wave activity. This facies encompasses the greater percentage of unit 13, lower and upper parts of unit 17, and units 15, 18 and 23 of the studied section (Figure 4c).

This facies contains dark gray, thin-bedded, poorly sorted limestones with skeletal grains including echinoderms, small foraminifers, gymnocodiaceans, ostracodes, and brachiopods within a micritic mud matrix. Bioturbation is a common feature in this facies. Bedding characteristics and rock texture point to a low rate of sedimentation and a low energy environment. This facies is interpreted as open-marine deposits in an outer-ramp setting. This facies is not represented extensively in the Dalan Formation and is rarely observed in units 13 and 14 (Figures 4d, 5e).

Light gray to yellowish gray, medium to thick beds are characteristic for this facies in the field. The skeletal grains consist of fusulinids, small foraminifers, brachiopods, green and red algae, corals (in growth position), Tubiphytes, gastropods, ostracodes, bivalves, crinoid stems in a micritic background. Minor allochems occur as intraclasts (3%). The skeletal grains have micritized boundaries. In some cases, skeletal grains are compacted in this facies and create a bioclastic packstone. Poorly to moderately sorted, angular to subrounded peloids (about 10%) form peloidal bioclastic packstones. The microfacies indicate that these sediments were deposited in a low energy lagoonal environment. This facies occurs in units 2, 3, 6, 8, 12, the upper part of unit 13 and unit 16 of the Il-e Beyk Section (Figures 4e, 5c).

This facies consists of medium- to thick-bedded grayish brown to dark gray beds and include relatively abundant and diverse fauna dominated by fusulinids, brachiopods, echinoderms, unbroken fenestrate bryozoa fronds, crinoid stems, dacyclad algae, Tubiphytes, red algae and smaller foraminifers. This facies shows scour faces, low-angle laminations and normal grading and is moderately to poorly sorted. Ostracodes and peloids represent minor constituents. The presence of intraclasts (15%) in some cases modifies this facies to intraclast bioclastic packstone-grainstone. Chert nodules are common in some beds of this facies (only in lowermost portion of unit 1 of the studied section) and can be termed cherty bioclastic packstone-grainstone.

The bioclast composition is variable, and this facies includes the following lithologies: a dacycladacean staffellid packstone-grainstone, a foraminifer bioclastic packstone-grainstone, an algal bioclastic packstone-grainstone, and a fusulinid packstone-grainstone. Some beds indicate the occurrence of dolomitization. The degree of dolomitization ranges from dolomite cement growth in inter- and intraparticle pore space to pervasive dolomitization with dolomite rhombs entirely replacing the original rock fabric. The skeletal grains have commonly micritized boundaries that show evidence of boring by endolithic algae (Flügel, 2010). The coarse grain texture and high biodiversity in this facies imply that they were deposited in a moderate energy environment most likely a near-shoal subtidal setting, either foreshoal or backshoal. Cross-bedding represents storm reworking. This facies is identifiable in units 1, 14, 28, the lower and upper part of the unit 4, and the middle part of unit 17 (Figure 4f).

This facies is composed of light to dark gray medium beds of poorly to moderately sorted peloids. Hummucky cross-stratification, low-angle lamination, and faint normal grading are among common features in this facies. Peloids are the main components and minor constituents are dominanted by brachiopods, bivalves, crinoids and gastropods and small foraminifers. Hummucky cross-stratification indicates a storm-dominated environment. This facies is observed in the middle part of unit 14 of the studied section and shows a moderate energy, open-marine setting representing a transition between mid- to outer ramp (Figure 4g).

The shoal sedimentary deposits are composed of an ooid dolograinstone with oomoldic porosity. This facies forms brown to cream, thick-bedded, massive-appearing, fine to coarse-beds and shows cross-bedding. In some beds intensive dolomitization obscures the original character of the microfacies. Where the original structure is preserved, ooids are recognizable as round grains with relics of internal concentric layering. Ooid moulds are filled with calcite cement (granular cement). In some cases, oomolds remain empty producing oomoldic porosity. Some oomolds have been filled with secondary dolomite and remain as ooid ghosts. Isopachous and blocky cements are common in this facies. Peloids and bioclasts are rarely observed. This facies indicates high-energy mid ramp shoal. Well-sorted oolitic grainstones are interpreted as shoal or bar complex deposits (Koehrer et al., 2010). This facies is observed in the middle part of unit 4 and in units 19 and 20 of the studied section (Figures 4h, 5b).

The Dalan Formation is predominantly a shallow-marine carbonate succession of Middle to Late Permian age, formed during several third-order transgressive-regressive depositional sequences (Lotfpour et al., 2005). On the basis of facies distribution and associated microfaunas, a homoclinal ramp (Read, 1985; Burchette and Wright, 1992; Ahr, 1998) is proposed for the Dalan Formation in the studied section, especially for its lower limestone part. The upper portion of the formation is predominantly dolomitic and makes the microfacies analyses and the interpretation of the depositiotinal environment exceedingly difficult. The formation consists of mud-dominated and grain-dominated facies encompassing eight lithofacies types. The encountered facies range from more protected backshoal to storm-dominated foreshoal to shoal deposits.

Deeper-water shelf facies, for instance in the Surmeh Section (Insalaco et al., 2006), were very rare. In the Dalan Formation, deposits are composed mostly of bioclastic, peloid, ooid allochems. The scarcity of faunal components indicates the existence of a restricted environment during deposition of dolomudstones. The bioclastic wackestone-packstone possesses the highest biodiversity and represents most viable taxa. This facies was deposited in leeward shoals. The bioclastic wackestone and bioclastic wackestone-packstone suggest the existence of an open lagoon behind the windward shoals and was the habitat for fusulinids, dasycladacean algae, crinoids, brachiopods, bivalve and gastropods. Ooid dolograinstone indicates a period of significant shoal development located on the windward side of inner ramp. Rare, open-marine facies are suggestive of deposition in calmer water of the mid ramp to possibly outer-ramp setting (Flügel, 2010).

It is noteworthy that staffellids are present in all these aforementioned lagoonal facies but their quantity varies. Below the Eopolydiexodina-bearing beds within the studied stratigraphic section, staffellids are abundant but their number decreases above the Eopolydiexodina persica Biozone and they become rare in facies between the Eopolydiexodina persica Biozone and Chusenella schwagerinaeformis-Rugososchwagerina zagroensis Biozone. Toward the Afghanella schencki Biozone they become absent altogether. Staffellids reappear in this facies upsection in the Upper Permian (Lopingian) towards the top of the Permian.

We studied Middle Permian Dalan Formation in only one Il-e Beyk Section in the Zard-Kuh area where the formation provisionally can be divided into two third-order depositional sequences that shallow upwards (Figure 2). Sequence 1 is Kubergandian to middle Murgabian (Roadian), consists of units 1 to 14, and is 320 m thick. It was deposited during a transgression over the Carboniferous? Faraghan Formation siliciclastics (Figure 5f). Although the Faraghan Formation is covered in the studied section, it is present below the Permian carbonates in other outcrops within the High and Folded belts.

The transgressive systems tract (TST) deposits of sequence 1 occur in units 1 to 13. They are medium to thick-bedded and composed mainly of bioclastic wackestone, bioclastic wackeston-packstone, and bioclastic packstone-grainstone. The deposits contain aboundant and diverse marine fauna such as fusulinids, small foraminifers, brachiopods, corals, crinoids, and gastropods. Bioturbation is also distinguishable in some microfacies. Higher up in the section, thin- to medium bedded burrowed mudstones and wackestones appear and show low-energy deeper water Zoophycus ichnofabrics. This part of the sequence shows a transition from shallow-marine to the open-marine carbonates.

The maximum flooding surface (MFS) occurs in the upper part of unit 13. It is represented by massive bioturbated mudstone to wackestone with large stylolites (Figure 5d) that shows lowest energy condition and sedimentation below storm wave base. The highstand systems tract (HST) occurs in unit 14, which consists of peloidal cross-bedded wackestone-packstone. Bioclasts decrease and include a low percent of small foraminifers and some rare shell fragments. It shows the shallowing upward deposits in which peloidal packstone to grainstone set over open-marine massive mudstones with bioturbation. The lower boundary of sequence 1 is a type 1 but its upper boundary is a type 2 as there is no evidence of a long-term withdrawal of sea level in the studies area.

Sequence 2 is late Murgabian to Midian (Wordian-Capitanian), includes units 15 to 23 and is 400 m thick. Its lower part consists of limestone and its upper part is dolomitized. The TST occurs in units 15 to 17 and is characterized by medium-bedded bioclastic packstone-grainstone with abundant brachiopod shells that further up in the section are replaced by bioclastic mudstone to wackestone with fusulinids, small foraminifer, brachiopods and crinoids. The maximum flooding surface occurs in the upper of part unit 17. It is thin and rarely preserved, and represented by dark beds containing bioturbation and open-marine skeletal fragments.

The HST of sequence 2 occurs in units 18 to 23 and is dominated by oolitic dolograinstones that pass upward into beige bioclastic dolomitized packstone/grainstone containing fusulinids Afghanella and Sumatrina. These facies show shore and back shore settings. The overlying beds are dolomitized. The upper part of this sequence coincides with the lack of distinct Middle Permian fusulinids. Because of dolomitization, the sequence-stratigraphic analysis of the upper 170 m of the Dalan Formation is unresolved. The upper boundary of sequence 2 is a type 1.

The Middle Permian to Early Triassic Khuff Formation is a widespread rock unit in the Arabian Basin. It represents a second-order transgressive-regressive sequence, which is composed of several third-order composite sequences (Sharland et al., 2001; Strohmenger et al., 2002; Alsharhan, 2006; Insalaco et al., 2006). In a sequence-stratigraphic study of Khuff-equivalent deposits (Saiq and Mahil formations) in Oman, Koehrer et al. (2010, 2012) interpreted six third-order sequences, named KS6 to KS1. KS6 and KS5 are interpreted to correspond to the Kubergandian (?) to Midian (early Roadian–Capitanian). Sequences KS4, KS3 and lower KS2 correspond to the Late Permian (Wuchiapingian and Changhsingian).

Based on the overall lithostratigraphic, microfacies features and microfossils contents (Kubergandian fusulinids) of the middle part of the Dalan Formation (see biostratigraphy section) in the studied section, we suggest that this part of the formation can be correlated with the upper part of KS6 and KS5. This suggestion requires further sequence-stratigraphic studies, especially for the Lower Dalan and Nar members in other outcrops within the Folded and High Zagros belts. Moreover, the Kubergandian (early Roadian) age of the lower part of Sequence KS6 has to be confirmed elsewhere in Zagros, Oman and surrounding areas. Further investigation of facies, depositional environments and sequence stratigraphy in the Zagros area is essential in order to understand the overall Permian stratigraphy in the region.

Studies of the Middle Permian (Guadalupian) succession in Zagros reveal a particular problem that was briefly mentioned but not explained by the previous studies (Altiner et al., 2000, and discussions in Ueno, 2003; Kolodka et al., 2012): namely, the exceptionally rare occurrences of fusulinids within the Middle Permian (Guadalupian) in the region. Although extinct, fusulinid paleobiology and their environmental constraints are relatively well-understood (Stevens, 1971; Ross, 1982, 1995; Vachard et al., 2004; Leppig et al., 2005). Fusulinids were advanced, single-celled, symbiont-carrying organisms distributed in late Paleozoic tropical-subtropical belts (up to 35– 40 degree south/north latitude) within carbonate to mixed carbonate-siliciclastic shallow-water settings (Ross, 1995; Kobayashi and Ishii, 2003a). This paleogeographic distribution, their known sensitivity to paleoenvironments (Figures 6 and 7), coupled with their high-resolution spatial and temporal framework provides the basis for the study presented here.

Most benthic foraminifera (including all larger foraminifera such as fusulinids) became extinct at the end of the Permian. As such, their paleobiology and paleoecology can only be inferred from studies of recent benthic foraminifera or other marine microfaunas with living symbionts. It is assumed that the majority of the fusulinids hosted photosynthetic symbionts, as is true for many living tropical larger foraminifera (Ross, 1982; Vachard et al., 2004; BouDagher-Fadel, 2008). The symbiosis could explain the rapid growth of fusulinids and their enormous single-cell size (up to 15 cm), as well as their rapid evolution, lower latitudinal geographical distribution, and shallow-water ecological patterns.

Recent symbiont-bearing benthic foraminifera are restricted to the euphotic zone of tropical and warm-temperate seas (Hohenegger, 2004; BouDagher-Fadel, 2008). Their species distribution is controlled by their primary limiting factors: temperature, light, water movement, substrate and nutrients (Beavington-Penney and Racey, 2004). Temperature determines geographic distribution and depth distribution of larger foraminifera by the development of a shallow thermocline that truncates the distribution of shallower water species, thus excluding species adapted to the deepest euphotic zone (Figures 6 and 7). Within these constraints, light is the other most important primary factor because larger foraminifera are at least partly dependent upon photosynthesis of their algal endosymbionts for growth and calcification (Hohenegger, 2004). Shallow-water assemblages of recent larger foraminifera with optimal water temperatures (20–30°C) are generally much more diverse than those with temperatures greater than 30°C and/or less than 20°C. (Beavington-Penney and Racey, 2004; Hohenegger, 2004; Murray, 2006).

Benthic foraminifera are poikilothermic, meaning that their body temperature is very close to that of the surrounding water, and they respond very quickly to the slightest changes in the ambient water temperature (Beavington-Penney and Racey, 2004). Temperature is generally considered to be the major physical factor influencing the distribution of species or assemblages (Hohenegger, 2004; Murray, 2006; BouDagher-Fadel, 2008). Fuhrman et al. (2008) clearly determined that the richness of shallow-marine bacteria with living symbionts (warm-water, shallow thermoclinal micro-organisms environmentally analogous to fusulinids) strongly and positively correlates with water temperature at the time of sampling, and even more strongly with average annual sea-surface temperature. Equally strong correlations of marine bacterial diversity with latitude and temperature suggest that latitudinal gradient directly reflects the effects of solar energy flux on sea-surface temperature. These data and analyses (Fuhrman et al., 2008) support the hypothesis that bacterial diversity in a given habitat is largely generated and maintained by the effects of temperature on the kinetics of metabolism.

The metabolic rate, which increases exponentially with increasing temperature, sets the pace of life and hence the rates of nearly all biological activities. Thus, we can hypothesize that during warming episodes warm-water foraminifera for given latitude and depth reached their maximum diversity and world-wide distribution, with tropical-subtropical taxa migrating to higher latitudes.

By contrast, cooling led to migration of foraminiferal fauna from higher to lower latitudes, their stepwise extinction, increasing provincialism and the preferential survival of euryfacial faunas (Stanley, 1984; Kalvoda, 2002; Fuhrman et al., 2008). We believe that fusulinid faunas provide a particularly sensitive index of climate changes within marginal conditions at the transition from subtropics to temperate mid-latitudes, such as existed in Zagros, Oman, Abadeh Transcaucasia, Salt-Range, Southeast Pamirs and other regions of the Neo-Tethys.

Mid-latitudes are areas where climatic fluctuations and the rapidity of climate change are most significant as it was shown in modern world (Burrows et al., 2011) (Figure 8) and in the Permian models (Winguth et al., 2002) (Figure 9). Consequently, the greatest variations in foraminiferal diversity occur at mid-latitudes as has has been demonstrated in modern brachiopods (Powell, 2009) and already noted for the record of Middle Permian (Guadalupian) foraminifers (Bond and Wignall, 2009). Recent development in Permian climate simulation models (Winguth et al., 2002) suggest a significant latitudinal shift in sea-surface temperature and associated climatic belts under hypothetical atmospheric pCO₂ concentrations attributable to glacial versus non-glacial climate states (Figure 9). This model proposes the most significant changes in temperature at mid-latitudes (approximately between 30–60° north and south from equator). The isotherms in mid-latitudes are much tighter and thus change more frequently than in polar regions (approximately between 60–90° in the south and the north) or tropical belt (approximately between 30°S and 30°N). Modeled fluctuations in atmospheric pCO₂ of up to 8x the threshold for permanent polar ice are apparently warranted by recent compilations of its variability through the Early Permian from proxy records (Montañez et al., 2007).

Three climatically-determined assemblages can be distinguished in Zagros and surrounding areas (Figure 8). The first one, characteristic for the cool-water environments includes inhabiting cool-water environment smaller foraminifera. The surface water in Zagros and surrounding areas of the habitats of this climatically-determined assemblage perhaps was temperate. This assemblage includes only benthic smaller foraminifera and may indicate cool climate episode or strong cold-water upwelling. The facies where this assemblage occurs usually is considered as heterozoan (James and Clarke, 1997; Weidlich, 2007).

Fusulinids occur in the region at warmer water environments. This second climatically-determined assemblage is dominated by staffellids, schubertellids and Chusenella (Table 2, Figure 2). It seems that Eopolydiexodina also belongs to this assemblage, which is transitional from cool/cold to warm water. Occurrence of fusulinids in limestone is usually interpreted to be a photozoan carbonates (James and Clarke, 1997; Weidlich, 2007). However, the fusulinids mentioned above generally occur at higher latitudes in (Kobayashi and Ishii, 2003b; Gaillot and Vachard, 2007).

Typical warm-water tropical condition in Zagros and surrounding areas (Figures 2 and 10) appears where surface-water temperature exceeded a level optimal for the recent warm-water larger foraminifera (Beavington-Penney and Racey, 2004). This third assemblage is characterized in Zagros by verbeekinids Afghanella and Sumatrina. However, the low taxonomic diversity in the assemblage (two verbeekinids genera represented by one species each) indicates that this warming episode in the region was short. The climate possibly shifted back to cool surface water near the end of the Capitanian or at the Capitanian/Wuchiapingian boundary (Figure 2) once the carbonate factory changed from protozoan to heterozoan type (Weidlich, 2007).

The beginning of Permian sedimentation in Zagros is associated with a Middle Permian–Early Jurassic Arabian Plate (AP) tectonostratigraphic megasequence 6 (AP6) that is related to the onset of rapid thermal subsidence of the early Neo-Tethys passive margin in Arabia and Iran (Sharland et al., 2001). The sedimentation is also associated with the continental rifting and spreading of the Sanandaj-Sirjan and Central Iran terrains away from the Arabian Plate. Because of poor age constraints of the Faraghan Formation the exact time of the beginning of the thermal subsidence and continental rifting is not clear. Kubergandian (early Roadian) fusulinids near the bottom of the Dalan Formation and the moderate thickness of the Faraghan Formation suggest that the thermal event is probably Bolorian–Yakhtashian (Kungurian–Artinskian); this is in agreement with the recent proposal by Gaillot and Vachard, (2007). At the same time, the most recent discovery of diverse assemblages of fusulinids in Moscovian–Gzhelian and in the Permian in Sanandaj-Sirjan Zone (Leven and Gorgij, 2008a, b, 2011) suggests the paleo-position of the region during this time was located within the tropics. Thus, we cannot exclude the possibility that the rifting of the Sanandaj-Sirjan region and Middle Permian (Guadalupian) thermal subsidence of the passive margin of Arabian Plate including Zagros occurred at different times.

The following paleogeographic and paleotectonic observations can be made from analyses of spatial and temporal distribution of foraminiferal assemblages from region to region within the Neo-Tethys (Figure 10).

The beginning of Middle Permian (Guadalupian) sedimentation in the Arabian Plate consists of siliciclastics characterized by the Gharif Formation in Oman, from which diverse plant fossils of Cathaysian, Euramerican and Gondwanan affinities were recovered by Berthelin et al. (2006). They recognized two assemblages in units A and B of the Gharif Formation. Plants from the lower unit A do not provide a precise age, whereas those from unit B indicate an early Kubergandian (early Roadian) to Murghabian transition. The siliciclastics of the Gharif Formation correlate to the lower part of Dalan Formation in Zagros.

Within the eastern shelves of the passive margin of the Arabian Plate (Khuff Basin, Figure 10) Middle–Late Permian (Guadalupian–Lopingian) fusulinids are extremely rare (Table 3). Very rare Monodiexodina kattaensis and Reichelina (?) were recovered from the lower Khuff Formation (Khuff D Member in subsurface near the Saudi-Oman border) and no fusulinids, except staffellids or dasycladacean algae, were observed in the upper part of the Khuff (Vaslet et al., 2005; Hughes, 2009). Bivalves, brachiopods, crinoids and bryozoans are the most abundant skeletal grains in the Khuff Formation at least in the Haushi area and the siliciclastic input is generally low. The biotic content is typical of the heterozoan association (Weidlich, 2007). Only rare staffellids, schubertellids and ozawainellids occur in the Upper Permian (Lopingian) (Gaillot and Vachard, 2007; Hughes, 2009) suggesting the existence of abnormal environments within the region (Table 2).

The rare occurrence of fusulinids in the Khuff Formation is explained either by strong terrigenous input (Angiolini et al., 2003) or by the onset of the end-Permian anoxia and/or warming related to a rapid change in the plate movements (Gaillot and Vachard, 2007). However, the occurrence of siliciclastics in the Khuff Formation is not significant (Vaslet et al., 2005) and thus the dispersal of fusulinids is possible. The siliciclastic component itself does not greatly influence the occurrence of fusulinids. Besides, it is quite high in fusulinid- and dasyclads-bearing carbonates in Abadeh, Transcaucasia (Kotlyar et al., 1989; Iranian-Japanese Research Group, 1981; Leven, 1998) and other regions. Therefore the siliciclastics are not responsible for the reduction of fusulinid occurrences.

The anoxia could be responsible for the diversity reduction in Khuff Basin, but no sedimentological evidence of anoxia (highly-organic black shales and/or framboidal pyrite) has been reported in the region (Vaslet et al., 2005; Insalaco et al., 2006). In our opinion the most probable explanation for the extremely poor taxonomic diversity of fusulinids, corals and dasyclads algae in the Khuff Basin are temperate (cool) shallow-water climate conditions that are consistent with carbonate sedimentology analyses (Vaslet et al., 2005). The temperate climate of the area could be attributed to the higher latitude position, around 40–45°S, of the Arabian Plate during Wordian–Capitanian time (Figure 9). Additionally, we cannot exclude the occurrence of strong and permanent upwelling currents that might have influenced the warm-water biotas.

The Permian in north and northeast of Oman consist of tectonic nappes formed from the relicts of Middle Permian–Late Triassic carbonate platforms include the (para-) autochtonous (?) Saiq Formation. The formation consists of a lower, thin siliciclastic unit and an upper thick carbonate unit. The lower siliciclastics may correlate to the Gharif Formation of the Khuff Basin, although it is quite thin (> 20 m) (Köhrer et al., 2010). Shallow-water, warm to temperate carbonate sedimentation of the Upper Saiq Member (over 600 m thick) started in the Wordian as indicated by the occurrence of typical warm-water verbeekinds Afhanella, Verbeekina and Pseudodoliolina (Forke et al., 2012). The diversity of fusulinids there are consistent with early Midian (Wordian–early Capitanian) assemblages in Transcaucasia and Abadeh.

The Capitanian part of the Saiq Formation is characterized by smaller foraminifera only, including the prominent Shanita amosi and Sphairionia sikuoides (Forke et al., 2012). The post-Wordian part of the Saiq Formation is completely dolomitized and similar in lithology to the Upper Dalan Member in Zagros (Figures 2 and 3). The content of fusulinid fauna in the lower part of the Upper Saiq Member suggests its paleotectonic and paleogeographic position within a warmer climate, close to Zagros rather than the internal part of the Arabian Plate. This interpretation contradicts the commonly accepted view on these areas as autochtonous units at the eastern rim of the Arabian Shield (Glennie et al., 1973; Weidlich and Bernecker, 2007). This autochthon, however, could be a nappe as well (Blendinger, 1988; Chauvet et al., 2009).

Each of these tectonic nappes of the Oman Mountains possesses a different stratigraphy and may represent two different basins. The Hawasina Nappes appear as a part of the Neo-Tethys breakup that is marked by pillow lavas up to 300 m thick (Glennie et al., 1973; Blendinger, 1988). The volcanoclastics that conformably covered the pillow lavas contain reefal conglomerates and boulders. Unfortunately, the reef debris and boulders, some of which are enormous (up to 200 m), are basically unstudied. Just Neoschwagerina margaritae and Verbeekina sp. were listed in the literature (De Wever et al., 1988; Weidlich and Bernecker, 2007). Within the lavas and above the volcanoclastics, a thin neritic limestone contains abundant Wordian ammonids (Blendinger et al., 1992), including evolute Timorites of Capitanian age (Baud et al., 2012). Fusulinids Kahlerina sp., Dunbarulla sp., and Codonofusiella sp. were recovered from the ammonoid matrix (first author’s personal collection). A late Capitanian Clarkina postbitteri hongshinensis conodont was found near the top of neritic limestone. A condensed, deep-water succession with radiolarian and rare conodonts extended into the Triassic (Baud et al., 2012). The existing data suggest the presence of a carbonate microplatform during the Wordian-Capitanian that later became the Hawasina Nappes; however, the nature, size and geometry of this platform remain uncertain due to very poor data.

The Batain Nappes consist of Lower Permian to Mesozoic sedimentary and volcanic rocks of the Batain Group that were obducted on top of the autochthonous sedimentary cover of the northeast Oman margin (Vachard et al., 2002; Weidlich and Bernecker, 2007). During the Permian the Batain Nappes may have been an isolated microplatform (sea-mounts or plateau?) that were located north from the Arabian Plate within the Neo-Tethys (Béchennec et al., 1990). The record of the platform is preserved as limestone conglomerates, boulders and blocks within the Mesozoic (Permian lost carbonate platform or Oman “exotics”). The record from these limestone blocks suggests the existence of the platform at least since Yakhtashian (Artinskian) through Capitanian (Vachard et al., 2002). The fusulinid diversity from the carbonates indicates the occurrence of the platform within the normal tropical belt. The record from the Batain Nappes, however, is still quite poor and additional studies are essential.

The Zagros region was essentially part of the passive margin of the Arabian Plate, but located approximately at 35–40°S (Figure 10). The frequent occurrence of Roadian–Capitanian fusulinids and their significantly greater taxonomic diversity than in the Khuff Basin (Table 3) suggests that the surface-water temperature there was warmer than in Khuff Basin. Nevertheless, the region was located within the temperate or warm-to-temperate transitional climate zone. Most of the time the surface water was cool, and fusulinid dispersal was restricted to relatively short periods, when the shallow-water temperature exceeded acceptable levels for conditions suitable for fusulinids. Schubertellids, staffellids, and chusenellids dominate the Roadian–early Wordian fusulinid assemblages without verbeekinids. The fusulinid-bearing horizons are quite thin and thus the occurrence of fusulinids in the area was relatively short.

In general fusulinids occur in the very thin (1–3 m thick packages within the 700 m succession (Figure 2). The upper Wordian thin horizon (about one meter thick) with typical temperate Monodiexodina kattaensis contains a reasonably diverse assemblage of fusulinids (Figure 2); these are taxonomically quite similar to faunas from Sibumasu and Baoshan blocks (Ueno, 2003), such as Rugososchwagerina, Kahlerina, Dunbarula, Yangchienia and Codonofusiella (Tables 1 and 2). About 200 m between this horizon and the next fusulinid-bearing horizon, a succession of dolomites and dolomitic limestones occurs with very rare smaller foraminifera or lacking any preserved fossils. The presence of dolomites precludes making a statement with confidence regarding the environment.

Two very thin horizons (less than 1 m thick) occurring at 660 and at 690 m contain the first and the only verbeekinids in the entire succession (Table 2, Figure 2). Their taxonomic diversity however, is relatively low. Only two genera of verbeekinids with a single species each and one species of schubertellid Yangchienia were recovered. With confidence we interpret this as an indication that a definite warm but short event occurred in the region; further it suggests that the surface temperature was not extremely elevated. Only smaller foraminifera were recovered from the Upper Permian (Lopingian) in the section and thus the climate and the environment in Zagros can be interpreted to have been somewhat similar to that of the Upper Permian (Lopingian) in the Khuff Basin.

The prevailing interpretations regarding the paleo-latitudinal position of the Zagros area during the Permian are quite controversial. It is commonly accepted that during pre-Wordian time Zagros, as a part of Arabian Plate, moved rapidly from 40°S to 25°S (Sharland et al., 2001) and by Wordian time it was located within the tropics (Kobayashi and Ishii, 2003a; Angiolini et al., 2003). More radical interpretations propose a position of the Arabian Plate during the Lopingian as far as at the Equator (Vachard et al., 2005; Gaillot and Vachard, 2007). The data from Zagros in our opinion suggest a more stable position of this part of Arabian Plate during the Middle Permian (Guadalupian). The changes and shifts of biota from non-fusulinid to fusulinid-bearing most probably related to the climate fluctuations rather than to tectonic movement of the Arabian Plate. Starting from Roadian through at least mid-Capitanian, Zagros was located within the temperate climatic belt, i.e. 35–40°S and only by the late Capitanian time could it have been closer to the Equator, still, however, within the tropical/subtropical transition (30–35°S).

Both regions were located within the eastern carbonate shelves of the Neo-Tethys Cimmerian terrains (Figure 10) and within a climatic zone that was warmer than Zagros. Artinskian– Kungurian thermal subsidence of the Arabian Plate was coeval with the beginning of Permian marine sedimentation in these regions (Figure 9). The Kungurian (Bolorian) assemblages in Transcaucasia are quite diverse and consist of typical warm-water taxa including the verbeekinid Misellina. Similarly, during Kubergandian and Murgabian (early Roadian–early Wordian) the verbeekinids Armenina, Cancellina, Pseudodoliolina and Presumatrina were common (Leven, 1998). Nevertheless, the diversity and abundance as well the frequency of occurrences of verbeekinids and other fusulinids are not as great as in typical tropical sections (Skinner and Wilde, 1966; Leven, 1997).

Only 6 to 8 horizons of fusulinids in Transcaucasia and about 12 horizons in Abadeh were reported within the entire Kuberganian and Murgabian (early Roadian to early Wordian), although the carbonate succession there exceeds 300 m (Leven, 1998; Kobayashi and Ishii, 2003b). Among these horizons in Transcaucasia only four contain verbeekinids. Five verbeekinid-bearing horizons occur in Abadeh. In the lower Midian (upper Wordian) of both regions the fusulinid assemblages are Chusenella-dominated. Very few and quite rare verbeekinids Verbeekina, Sumatrina and Neoschwagerina schubertellids and ozawainellids were recovered there (Leven, 1998; Kotlyar et al., 1989; Kobayashi and Ishii, 2003b).

The upper Midian (Capitanian), Chusenella-dominated assemblage were supplemented with one specimen of advanced Neoschwagerina (N. pinguis) recovered from the Khachik Formation in Transcaucasia (Kotlyar et al., 1989). Yabeina and Metadoliolina were reported from the Hambast section in Abadeh (Baghbani, 1993) from about upper Midian (Capitanian) strata. However, this occurrence has not been confirmed by subsequent studies (Kobayashi and Ishii, 2003b, and personal study of the authors). The present fusulinid data from Abadeh and Transcaucasia suggest that during Roadian–Wordian time the regions were located within the subtropics with annual temperature favorable but marginal for the appearance for verbeekinids. The limited diversity and low frequency of these taxa are evidence that the surface water in both regions was not so high (perhaps 20–25°C) possibly due to permanent upwelling currents that kept surface water temperature low.

According the tectonic data the region was located within the west Qiangtang Microcontinent (Hutchison, 1993). Very thick Carboniferous–Sakmarian siliciclastics with cold-water Conularia and Eurydesma clearly express the position of the region at that time within the Gondwana margins. Post-Sakmarian–pre-Kungurian (pre-Bolorian) pillow lava basalts, the first temperate fusulinids Monodiexodina and cool-water Gondwanan conodonts Vjalovognathus (Leven, 1967; Kozur, 1978) suggest that the Artinskian rifting of this microcontinent towards the tropics was rather rapid. Because in the mid-Kungurian (Bolorian) Pamirs and surrounding areas of Qiangtang microcontinent were already located within the tropics, fusulinid assemblages of the region were dominated by verbeekinids through the entire Middle Permian (Guadalupian) (Nie and Song, 1983a, b; Leven, 1997). Chusenella, Rugososchwagerina and Yangchienia that are the most common fauna in Zagros are quite rare in the Southeast Pamirs (Leven, 1967). In the composite section of the Southeast Pamirs we recognized over 30 verbeekinid-dominated horizons within the Kungurian–Capitanian, suggesting very warm-water environments for most its duration (Table 3). Volcanoclastic facies developed in Southeast Pamirs during the Upper Permian (Lopingian). Although these facies are not favorable for the fusulinids and corals, the assemblages are typical of diverse warm-water tropical biota (Kotlyar et al., 1983, 1989).

The region belongs to the northern part of Greater India and underwent separation and early rifting phase during the late Paleozoic. The Middle and Late Permian sedimentary development is characterized by two transgressive-regressive (T-R) second-order cycles. In the Salt Range area, these two T-R cycles have been subdivided into five third-order sequences (Baud et al., 1996). The late Paleozoic succession in the Salt Range is quite similar to that in the Zagros area, except that the marine deposits began later in the Wordian Amb Formation. The formation possesses temperate-water Monodiexodina kattaensis and Codonofusiella laxa in the middle part (Douglass, 1970). It seems that the occurrence of Monodiexodina kattaensis in middle Amb Formation is coeval with that of Monodiexodina kattaensis dalanensis in Zagros (late Wordian).

Above the Amb Formation, the Wargal Formation is divided into a lower thin and an upper thick part (Mertmann, 2003) that are bounded by sequence boundaries. The lower part of the Wargal Formation contains abundant staffellids and rare Chusenella and schubertellids. A single specimen of warm-water verbeekinid Neoschwagerina margaritae has been recovered in only one locality (Pakistani-Japanese Research Group, 1985). The lower Wargal with Neoschwagerina margaritae corresponds to the Capitanian (upper Midian) Khachik Formation with Neoschwagerina in Transcaucasia (Kotlyar et al., 1989) and with Afghanella schencki Biozone in Zagros (Figure 3). According to conodonts (Wardlaw and Pogue, 1995) the lower Wargal is Capitanian. Thick upper Wargal and Chidru formations possess Codonofusiella, Reichelina, Nanlingella simplex and several biostratigraphically significant smaller foraminifers (Colaniella, Paradagmarita etc.) of Lopingian age (Pakistani-Japanese Research Group, 1985) suggesting an age consistent with conodont ages (Wardlaw and Pogue, 1995).

The distribution of fusulinids and other foraminifers in the succession suggest that by late Wordian time the Salt Range area was positioned within the mid-latitude with temperate water conditions similar to that of Zagros, i.e within the subtropical-tropical transition (40–35°S). The appearance of Neoschwagerina in the Capitanian of the Salt Range is coeval with the appearance of verbeekinids Neoschwagerina in Transcaucasia, and Sumatrina and Afghanella in Zagros, and is consistent with the proposed middle Capitanian warming event. The Late Permian (Lopingian) assemblages of the Salt Range area seem more diverse than those in Zagros, indicating that the former region was located at slightly lower latitude and experienced warmer-water climate than the latter.

In conclusion, it appears clear to us that the proposed model of climate–biota interaction developed on the basis of the Zard-Kuh data represents the initial step towards a better understanding of Middle Permian (Guadalupian) climate in the area. This model requires further consideration and must be tested in other Zagros sections and elsewhere in the Neo-Tethys.

Order FUSULINIDA Fursenko, 1958

Family SCHWAGERINIDAE Dunbar and Henbest, 1930

Subfamily MONODIEXODININAE Kanmera, Ishii and Toriyama, 1976

Genus Monodiexodina Sosnina in Kiparisova et al., 1956

Monodiexodina kattaensis dalanensis, new subspecies

Plate 7, Figures 10–19 

Holotype:Plate 7, Figure 12, SUI 132123, axial section; sample Z-33, Il-e Beyk section, Dalan Formation, unit 17, 433 m above the base of the section; Zard-Kuh area, Zagros, southwest Iran; Middle Permian (Guadalupian), Midian (upper Wordian–Capitanian).

Synonymy:Monodiexodina kattaensis (Schwager, 1887) in Douglass, 1970, G2-G5, his Plate 4, Figures 1–8 (only).

Material examined: 10 axial sections and many paraxial sections.

Etymology: After Dalan Formation where this subspecies has been found.

Type locality: Il-e Beyk section in Zard-Kuh area, Zagros, southwest Iran.

Type level: The nominative species has been described from the Wordian in Salt Range, Pakistan and is also probably found in Oman from Wordian (?). It has been recovered for the first time in Zagros and due to some morphological differences described as a new subspecies. It characterizes a narrow horizon within upper Wordian.

Description: Test medium to large, subcylindrical with bluntly pointed to rounded poles. Coiling is moderate and rather uniform throughout the growth. Mature individuals have 4–7 volutions and are 7–11.4 mm in length and 2–2.9 mm in diameter, giving form ratios of 3.2–4.4. The first volution is 1.3–2.9 mm long and 0.4–0.6 mm wide with form ratios of 1–4.4. The length and diameter of the fourth volution are 5.1–8 mm and 1.3–1.88 mm form ratios of 3.6–4.5. The proloculus is medium to large in size, spherical and commonly distorted in shape; its outside diameter varies from 275–575 microns. The proloculus/final volution diameters ratio is 5–10. Spirotheca is composed of tectum and coarsely alveolar keriotheca, and thickens gradually from third volution. Septa are strongly and regularly fluted from pole to pole. Septal folds are narrow and low and reaching to half as high as chambers but particularly extending to tops of chambers in polar regions. Low cuniculi are observed in two outermost volutions (Plate 7, Figure 14). Tunnel is low and narrow with rather regular path. Chomata are absent or only present on proloculus and first volution. Weak or moderate axial fillings are present in middle and outer volutions.

Comparison: Superficially this subspecies resembles Monodiexodina kattaensis (Schwager, 1887) from Salt Range, northeast Pakistan. Douglass (1970), who re-described the species from many locations in the middle Amb Formation of Salt Range and Khisor Range, noted two morphologically separate groups that can be seen as “different and statistically separable species” (Douglass, 1970, p. G2). Particularly they differ in proloculus and overall test size, wall thickness, volution height and tunnel angle (p. G8-9, Figs 5 and 6). One group described from Kattha location in Salt Range (Douglass, 1970, his Plate 2) is characterized by medium sized test and proloculus, medium wall thickness, relatively tight coiling, wide tunnel and medium to well-developed axial fillings. The Kattha location is the topotype of the species according Schwager (1987) and Dunbar (1933) and therefore there are no doubts the forms from this group resemble the lectotype of the species designated by Douglass (1970). It is also obviously close to the forms described by Dunbar (1933) from Warchha location.

The second group of Monodiexodina from Kafirkat location in Khisor Range is characterized by a large test and proloculus, medium coiling, relatively thick wall, narrow tunnel and weakly developed axial fillings. This group closely resembles forms of Monodiexodina we recovered in Zagros. Douglass, however, did not separate these two groups and consider these differences as morphological variations because of the same growth curve (the diameter/length, height of the volution and wall thickness). Unfortunately no information has been provided for the geographic distribution of these clines and on the microfacies. At the same time, the forms with large proloculus, greater test and weakly developed axial fillings are described from one Kafirkat location in the Khisor Range, whereas forms with medium size test and proloculus seems widely distributed in the Salt Range (Douglass, 1970, his Plate 4). It is also not clear if the forms with large proloculus and test occur within the population of forms with medium proloculus and vice versa. Douglass noted, however, that no correlation of morphology with stratigraphy and/or geography was recognized. Monodiexodina from Zagros possesses strong similarities with those described from Kafirkat location and designated here as a new subspecies of the Monodiexodina kattaensis. It differs from the nominative subspecies in having twice as large proloculus, generally larger overall test size, thicker wall, more regularly fluted septa both in central and polar ends of the test, narrow tunnel, and weakly developed axial fillings or sometimes no fillings at all. It might be that after additional study of forms from Kafirkat location in Salt Range the status of this subspecies could be raised to species level.

Stratigraphic range: Unit 17, Sample Z-33, Dalan Formation, Middle Permian (Guadalupian), Murgabian (early Roadian–early Wordian).

Geographic occurrence: Il-e Beyk section in Zard-Kuh area, Zagros, southwest Iran.

Monodiexodina hendersoni, new species

Plate 7, Figures 1–9 

Holotype:Plate 7, Figure 2, axial section; sample Z-33, Il-e Beyk section, unit 17, 433 m above the base of the section, Zard-Kuh area, Zagros, southwest Iran; Middle Permian, Guadalupian, Murgabian (early Roadian–early Wordian).

Material examined: 11 axial sections and many paraxial sections.

Etymology: Named after Professor Charles Henderson, the former chair of Permian Subcomission of the ISC and well-known Permian conodont worker.

Type locality: Il-e Beyk section in Zard-Kuh area, Zagros, southwest Iran.

Type level: lower Midian (upper Wordian).

Description: Shell medium, short elongate-fusiform with weakly convex median area and lateral slopes and bluntly rounded poles, the expansion of volutions is gradual during the growth. Coiling is tight in early volutions but uniform and gradual in later ones. Mature tests of 5–6 volutions are 6–9 mm long and 1.9–2.7 mm wide, giving form ratios of 2.9–3.8 The first volution is 0.7–1.7 long and 0.5–0.7 mm wide with form ratios of 1–2.4. The length and diameter of the fourth volution are 4.2–6.8 mm and 1.3–2 mm, respectively, with form ratios of 3.1–4.2. The proloculus large in size, spherical, ellipsoidal and often distorted in shape with outside diameter of 337–600 microns averaging 437 microns for 7 specimens. Spirotheca is relatively thin in early two volutions and gradually thickens at the later ones and is composed of tectum and coarsely alveolar keriotheca. The thickness of the spirotheca in the first, fourth and last volutions of 7 specimens are 15–25, 50–100, and 75–100 microns, respectively. Septa are very regularly and intensely fluted throughout the length of the shell. Fluting is extending to half the height of the chambers. Septa are less than twice as thin as the wall in respective volutions. Tunnel is low with a rather irregular path. Chomata are lacking or only recognizable on proloculus. Axial fillings are absent or only present on middle or outer volutions.

Comparison:Monodiexodina hendersoni can be distinguished from the comparable Monodiexodina kattaensis (Schwager, 1887) from Salt Range, northeast Pakistan, in its short, elongate-fusiform outline in all volutions as opposed to subcylindrical outline on the latter; it has more regularly fluted septa especially in central part of the shell, a larger proloculus and narrow tunnel. The described species possesses strong similarities with Monodiexodina kattaensis dalanensis subsp. nov., but differs from it in elongate-fusiform outline in all volutions as opposed to subcylindrical outline throughout the growth in the latter.

Stratigraphic range: Unit 17, Dalan Formation, sample Z-33, Middle Permian, (Guadalupian), Murgabian (early Roadian–early Wordian).

Geographic occurrence: Il-e Beyk section in the Zard-Kuh area, Zagros, southwest Iran.

Genus RUGOSOSCHWAGERINA A.D. Miklokho-Maklay, 1959

Rugososchwagerina zagroensis, new species

Plate 4, Figures 5-12

Holotype:Plate 4, Figure 8, axial section; sample Z-55, Il-e Beyk section, Dalan Formation, unit 15, 330 m above the base of the section, Zard-Kuh area, Zagros, southwest Iran.

Material examined: 15 axial sections and 7 paraxial sections.

Etymology: Named after Zagros region.

Type locality: Il-e Beyk section in Zard-Kuh area, Zagros, southwest Iran.

Type level: upper Murgabian–lower Midian (Wordian).

Description: Test is medium in size, inflated fusiform to subglobular with slightly extended and pointed poles. The coiling is straight. First two to four volutions constitute tightly coiled juvenarium, which is followed by sharp inflation into loosely coiled adult stage. Fully grown specimens of 5–7 volutions are 5.9–7.4 mm long and 2.6–3.8 mm wide with form ratios of 1.25–2.5. The first volution is 0.3–1.1 mm long and 0.2–0.4 mm wide, giving form ratios of 1.1–3.4. The length and diameter of fourth volution are 2.1–3.6 mm and 0.8–1.6 mm with form ratios of 1.8–2.9. Proloculus is spherical, small to medium in size with an outside diameter of 150–250 microns. Spirotheca is composed of tectum and rather coarsely alveolar keriotheca with thickness that varies from thin in juvenarium to abruptly thick in later volutions. The thickness of the spirotheca in the first, fourth and last volutions of 12 specimens are 7–12, 50–80, and 100–130 microns, respectively. Septa are nearly plane or slightly fluted initially, becoming intensely and deeply fluted from pole to pole in loosely outer volutions. Septal folds of outer whorls are high, subsquare to triangular and rather irregular in shape commonly reach tops of chambers. Tunnel is very narrow and low in early volutions and rather narrow with irregular path outwards. Weak chomata are present in the juvenarium and absent in later whorls. Prominent axial fillings along the axis well are developed in tightly coiled volutions; they nearly absent in 2–3 outermost volutions.

Stratigraphic range: Unit 15, sample Z55-Z57, Dalan Formation, Middle Permian (Guadalupian), Murgabian (early Roadian–early Wordian).

Geographic occurrence: Il-e Beyk section in Zard-Kuh area, Zagros, southwest Iran.

The thick (950 m) Middle–Upper Permian (Guadalupian–Lopingian) carbonate succession of the Dalan Formation has been studied in the Zard-Kuh Mountains of Zagros in Iran. Careful documentation of microfacies throughout the section clarified the environments and particularly the bathymetry of the foraminifera within the succession. Foraminiferal study of 46 samples from the Dalan Formation provided precise chronostratigraphic constraints and taxonomy. Five fusulinid biozones in the Middle Permian (Guadalupian) have been established for the first time in the Zagros.

The Lower Dalan Member has been correlated with Tethyan and Global scales. Particularly, the analogues of Kubergandian (early Roadian), Murgabian (late Roadian–early Wordian) and Midian (late Wordian and Capitanian) stages were recognized. The Kubergandian (early Roadian) age of the basal part of the Dalan Formation was documented for the first time in the region. In the studied material 66 species and 38 genera of foraminifera were recognized, two new species and one subspecies are established and described. Temperate water Monodiexodina was recorded for the first time in the area. The model of interaction of biota and climate during the Middle Permian (Guadalupian) in the Zagros region is proposed. The unique position of Zagros, within mid-latitudes where climatic fluctuations and velocity of the climate change and consequently the variations in foraminiferal diversity are most significant, open the opportunity to better understand the Middle–Late Permian (Guadalupian–Lopingian) climate in this and surrounding areas of Neo-Tethys. Three climatically-determined assemblages were distinguished in the Zagros and surrounding areas. The developed model of climate fluctuations and paleogeography in the Neo-Tethys is based on analyses of temporal and spatial distribution of fusulinids.

This study was partially supported by NSF grants EAR-0746042 and EAR-1004079 (VID). We thank Shuzhong Shen and Yichun Zhang for providing some rare Chinese literature. We warmly regard the help of Claude Spinosa with English improvements. The notes and suggestions of two anonymous reviewers improved the manuscript. We appreciate the suggestions by GeoArabia’s Editor-in-Chief, Moujahed Al-Husseini, the proofreading of the manuscript by GeoArabia’s Assistant Editor Kathy Breining and the design of the paper for press by GeoArabia’s Designer Arnoldo Egdane.

Vladimir I. Davydov is a Research Professor at the Department of Geosciences of Boise State Univeristy, Idaho, USA. He studied geology and paleontology at Tadzhik State University, Dushanbe and his Master diploma thesis dealt with sedimentology and foraminiferal biostratigraphy of southern Tian-Shan. His PhD dissertation at All Russian Geological Research Institute (VSEGEI) in St. Petersburg, Russia, focused on Pennsylvanian stratigraphy and foraminiferal biostratigraphy of North Pamirs, Tadzhikistan. Vladimir started his carreer as a Mapping Geologist at Tadzhik Geological Survey where he mapped the southern part of North Pamirs. After completing a PhD he worked at Department of Stratigraphy and Paleontology in VSEGEI and was part of the team responsible for the Carboniferous of the USSR. Since 1995 he is working at Permian Research Institute of Boise State University. Besides Russia and CIS, Vladimir worked in many regions global wide, including USA, Norway, Japan, China, Bolivia, and Iran. Currently he is involved in several projects dealing with paleoclimate. Chronostratigraphy and quantitative biostratigraphy integrated with high-resolution ID-TIMS is another topic that is his great interest.

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Sakineh Arefifard is a Lecturer at the Department of Geology of Lorestan University, Khorramabad, Iran. She studied stratigraphy and paleontology at Shahid Beheshti University, Tehran, Iran. Her MSc was on the biostratigraphy of Permian strata in Alborz (north Iran) and her PhD on sequence stratigraphy, biostratigraphy and geochemistry of the Permian in the east-central Iran (Push-t Badam block). Recently she has also worked on sequence stratigraphy and depositional environment of Permian deposits in east-central Iran in cooperation with Professor Peter Isaccson from University of Idaho, USA.

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