The Talesh Mountains at the NW margin of the Iranian Plateau curve around the southwestern corner of the South Caspian Block and developed in response to the collision of the Arabian-Eurasian Plates. The timing, rates, and regional changes in late Cenozoic deformation of the Talesh Mountains are not fully understood. In this study, we integrate 23 new apatite and zircon bedrock U-Th/He ages and structurally restored geologic cross sections with previously published detrital apatite fission track data to reconstruct the deformation history of the Talesh Mountains. Our results reveal that slow rock exhumation initiated during the late Oligocene (~27–23 Ma) and then accelerated in the middle Miocene (~12 Ma). These events resulted in the present-day high-elevation and curved geometry of the mountains. The spatial and temporal distribution of cooling ages suggest that the Oligocene bending of the Talesh Mountains was earlier than in the eastern Alborz, Kopeh Dagh, and central Alborz Mountains that initiated during the late Cenozoic. Late Oligocene and middle Miocene deformation episodes recorded in the Talesh Mountains can be related to the collisional phases of the Arabian and Eurasian Plates. The lower rate of exhumation recorded in the Talesh Mountains occurred during the initial soft collision of the Arabian-Eurasian Plates in the late Oligocene. The accelerated exhumation that occurred during final collision since the middle Miocene resulted from collision of the harder continental margin.

Key Points

Late Cenozoic deformation in the Talesh Mountains initiated at late Oligocene (27–23 Ma) then accelerated during middle Miocene (12 Ma)
Bending of the Talesh Mountains around the South Caspian Block starting in the late Oligocene and accelerated during middle Miocene
Episodes of deformation recorded here for the Talesh Mountains attributed to the collisional phases of the Arabian and Eurasian Plates

1 Introduction

The continental collision between the Arabian and Eurasian Plates formed a large orogenic system that comprises the Lesser and Greater Caucasus, Alborz, Talesh, Kopeh Dagh, and Zagros mountain ranges at the northern and southern boundaries of the Iranian Plateau (Figure 1). The spatial pattern of deformation, uplift, and exhumation related to this complex orogenic system is well constrained by structural, stratigraphic, and geochronologic and thermochronologic analyses in the central Alborz and Caucasus Mountains located at the northern margin of the collision zone [Axen et al., 2001; Guest et al., 2006a; Vincent et al., 2007, 2011; Avdeev and Niemi, 2011; Rezaeian et al., 2012; Ballato et al., 2013]. However, the temporal evolution of the continental collision is a matter of debate. The onset of collision has been suggested to have occurred in the late Cretaceous-middle Paleocene (80–60 Ma) [e.g., Haynes and McQuillan, 1974; Stöcklin, 1974; Berberian and King, 1981; Alavi, 1994; Mohajjel et al., 2003; Ghasemi and Talbot, 2003], in the late Eocene-Oligocene (35–25 Ma) [e.g., Hempton, 1985; Jolivet and Faccenna, 2000; Hessami et al., 2001; McQuarrie et al., 2003; Agard et al., 2005; Horton et al., 2008; Ballato et al., 2011], or in the Miocene (20–5 Ma) [e.g., Sengör and Kidd, 1979; Philip et al., 1989; Axen et al., 2001; McQuarrie et al., 2003; Guest et al., 2006b; Mouthereau, 2011; Mouthereau et al., 2012; Madanipour et al., 2013]. Conceptual models for soft and hard collision during continental convergence are relevant to the Arabian-Eurasia collision. Soft continental collision means that most of the plate convergence is absorbed by subduction processes associated with a partial decoupling of sediments from the subducting continental crust and possible delamination of subducted continental crust at depth [e.g., Chemenda et al., 1996, 2000; Regard et al., 2003; Toussaint et al., 2004]. During soft collision, a small fraction of plate convergence is accommodated within the colliding plates by deformation in a fold and thrust belt on the lower plate and formation of an intracontinental mountain belt on the upper plate [e.g., Regard et al., 2003]. During soft collision, a lower rate of deformation, uplift, and exhumation is predicted for mountain belts forming on the upper plate. However, during hard collision, deformation is accommodated by upper and lower plate crustal shortening, crustal thickening, and related uplift and exhumation coupled with lateral extrusion of crustal blocks [e.g., Regard et al., 2003; Toussaint et al., 2004]. In this study we aim to quantify the timing of initial collision, when the stretched continental crust of the Arabian and Eurasian Plates commenced. In the process of quantifying this, we also evaluate the temporal evolution of soft versus hard collision in the region.

Details are in the caption following the image — **Figure 1**
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(a) General structural features in the Arabia-Eurasia continental collision zone. The brown shaded areas represent the high-elevation regions of the Arabia-Eurasia collision zone. The GPS velocity vectors (black arrows in mm/yr) show the velocity of the Arabian Plate relative to the fixed Eurasian Plate [*Vernant et al*., 2004]. Red shaded area in the box represents the location of the Talesh Mountains. White shaded boxes represent the location of the swath profiles. (b) Swath profile showing minimum, maximum, and mean elevation along the Alborz to Zagros transect and (c) along the Talesh to Zagros transects.

Our study is located in the Talesh Mountains of the NW Iranian Plateau that borders the western edge of the South Caspian Sea (Figure 1). The curved geometry of the mountains is a result of the interaction with the rigid South Caspian Block during the Arabia-Eurasia continental collision [Berberian and King, 1981; Allen et al., 2003; Brunet et al., 2003; Allen, 2010; Madanipour et al., 2013]. Cenozoic deformation in the Talesh Mountains has previously been investigated through detrital fission track analysis in conjunction with tectonostratigraphic studies [Vincent et al., 2005; Madanipour et al., 2013] that reveal three major deformation phases that occurred in the early Oligocene (27–23 Ma), early to middle Miocene (18–12 Ma), and early Pliocene (~5 Ma). These phases are attributed to the initial and final Arabia-Eurasia collision and subsequent structural reorganization [Madanipour et al., 2013]. However, little is known about the magnitude and rates of deformation and rock exhumation related to these phases, as well as their spatial variations along the orogen strike.

In this study, we combine structurally restored cross sections with apatite and zircon U-Th/He (AHe and ZHe) data of bedrock samples to understand the prominent deformation events and related exhumation in the Talesh Mountains. Our new data document three major phases of exhumation, including the late Oligocene (~27–23 Ma), middle Miocene (~18–12 Ma), and early Pliocene (~5 Ma). A distinct acceleration in exhumation is observed in the middle Miocene (~12 Ma). The spatial distribution of cooling ages suggests that oroclinal bending occurred during the late Cenozoic. Our new results are integrated with published time constraints on the oroclinal bending in the central and eastern Alborz, Lesser Caucasus, and Kopeh Dagh Mountains, to regionally constrain the deformation at the entire northern boundary of the Arabia-Eurasia collision zone.

2 Geological Setting of the Talesh Mountains

The Talesh Mountains are part of the Arabia-Eurasia collision zone and connect the Alborz Mountains with the Lesser Caucasus located at the southwestern corner of the South Caspian Basin (SCB) (Figures 1 and 2a). The Talesh Mountains are composed of rock units ranging from the lower Paleozoic to Quaternary (Figure 2). Lower Paleozoic slightly metamorphosed carbonate and clastic rocks are the oldest rock units exposed in the hanging wall of the main thrust faults in the southern Talesh Mountains (Masuleh Dagh Fault (MD) and Boghrov Dagh Fault (BD)) [Davies et al., 1972; Clark et al., 1975; Alavi, 1996; Zancheta et al., 2009; Zanchi et al., 2009; Madanipour et al., 2013]. The lower to middle Paleozoic strata of the Talesh Mountains comprise volcanic, clastic, and carbonate sequences that were possibly deposited in an extensional basin [e.g., Berberian and King, 1981; Allen et al., 2003; Guest et al., 2006a, 2006b; Zancheta et al., 2009] (Figure 3). The Mesozoic sequence of the Talesh Mountains, similar to other parts of the northern Iranian Plateau, is generally composed of Triassic to Early Jurassic carbonate and clastic rock units that are unconformably covered by clastic rocks of the Jurassic Shemshak Formation that resulted from the Cimmerian orogeny (Figure 3) [Zanchi et al., 2009; Zancheta et al., 2009].

Late Paleozoic Variscan metamorphism and the Cimmerian orogeny (~230–200 Ma) are the main pre-Cenozoic deformation events in the Talesh Mountains (Figure 3) [Davies et al., 1972; Clark et al., 1975; Alavi, 1996; Zancheta et al., 2009; Zanchi et al., 2009]. ⁴⁰Ar/³⁹Ar cooling ages of white micas suggest a Late Carboniferous age (315 ± 9 Ma) for the Variscan nappes [Zancheta et al., 2009]. The emplacement of the Variscan metamorphic rocks on the Talesh Mountains occurred in the Early Jurassic [Zancheta et al., 2009] (Figure 3). New Petrological investigations on these metamorphic units integrated with whole-rock geochemistry and ⁴⁰Ar/³⁹Ar phengite geochronology represent subduction of a branch of the Paleo-Tethyan ocean, with peak metamorphism equilibrated under blueschist to eclogite facies conditions during the Early Carboniferous (~350 Ma) [Rossetti et al., 2017]. The Cimmerian orogeny-related deformation in the Talesh Mountains is covered by base conglomeratic sequences of the Shemshak Formation. This conglomeratic unit has been prograded over the Paleozoic metasedimentary rocks that were deformed during Cimmerian orogeny [Zanchi et al., 2009]. The Late Paleozoic and Cimmerian events are strongly obscured by the Cenozoic deformation [Alavi, 1996; Zanchi et al., 2009; Madanipour et al., 2013]. A late Cretaceous to Paleocene compressional event that occurred along the Neotethyan subduction zone [Guest et al., 2006a, 2006b; Yassaghi and Madanipour, 2008] is evident in the Talesh Mountains by a sequence of conglomerate and sandstone, which are Late Cretaceous-early Paleocene in age (Figure 3). This conglomeratic unit is unconformably covered with the lower Eocene carbonate units (Figure 3).

The ~8 km thick Cenozoic strata in the Talesh Mountains mainly comprise volcanic and volcaniclastic rock units, which were rapidly deposited (~2 Myr) in an extensional sedimentary basin (Figure 3) [Vincent et al., 2005]. Cenozoic sedimentary and volcanic sequences are exposed in the central and northern Talesh Mountains, and older rock units are only locally exposed in the core of some anticlines (Figure 2b) [Allen et al., 2003; Vincent et al., 2005; Madanipour et al., 2013]. Whole-rock ⁴⁰Ar/³⁹Ar dating of lava flows in these volcaniclastic rocks yields an age of ~39 Ma [Vincent et al., 2005]. Most of these volcanosedimentary units are proposed to have formed in a back-arc extension/transtension basin associated with north dipping Neotethyan subduction zone in the northern Talesh Mountains [Vincent et al., 2005]. Lower Cenozoic volcanic and volcaniclastic rock units are unconformably covered by dominantly coarse- to medium-grained clastic rocks of upper Cenozoic sequences (Figure 3). This angular unconformity documents that the early-middle Eocene extension was followed by compression in the late Eocene-early Oligocene. This event is interpreted as the beginning of the late Cenozoic orogenic evolution comprising three main deformation phases (Figure 3) [Vincent et al., 2005; Madanipour et al., 2013].

The curved geometry of the Talesh Mountains is shown by the southern NW trending segment, the central north trending segment, and the northern NW trending segment of the mountain range (Figure 2). The southern segment has stratigraphic and structural affinities with the neighboring west and central Alborz Mountains. The NW trending Talesh Fault forms a topographic barrier and constitutes the northeastern boundary of the southern Talesh Mountains to the coastal plains of the Caspian Sea (Figure 2b). The BD and MD faults are the main structures that juxtapose pre-Cenozoic rocks over Neogene stratigraphic sequences of the Shahroud and Qezel Owzan intramountainous basins (Figure 2b) [Madanipour et al., 2013]. Whether the BD and MD faults involve basement units remains unknown in most places due to lack of exposures. Nevertheless, the upper Paleozoic rocks appear to be affected in the southern parts of the BD and MD faults (Figure 2b). The NE trending central part of the Talesh Mountains is structurally controlled by the northern continuation of the Talesh, MD, and BD faults (Figure 2). Seismicity and field-based kinematic analysis indicate right-lateral strike-slip kinematics for the MD and BD faults in the central part of the Talesh Mountains [Berberian and Yeats, 1999; Allen, 2010; Madanipour et al., 2013].

The northern Talesh Mountains comprise a shallow NE verging fold and thrust sequence that merges into a main detachment fault at depth [Allen et al., 2003; Vincent et al., 2005; Madanipour et al., 2013]. Earthquake focal mechanisms show that the present-day active faulting and deformation in the western flank of the Talesh Mountains are localized along a low-angle thrust at 15–20 km depth with a general slip toward the Caspian Basin [Jackson et al., 2002; Allen et al., 2003; Aziz Zanjan et al., 2013]. The Ojagh Gheshlagh Fault is one of the major structural features in the northern Talesh Mountains that juxtaposes the lower Eocene volcaniclastic rock units against the upper Miocene clastic rocks (Figure 2b).

The basement of the SCB is slightly thicker than other oceanic basins, with a mean thickness of ~10 km (Figure 3) [Berberian, 1983; Knapp and Connor, 2004] that is overlain by ~20 km of Cenozoic strata [Berberian, 1983; Zonenshain and Le Pichon, 1986]. During the late Cenozoic this oceanic lithosphere was subducted northward along the Apsheron-Balkahn sill (Figure 1) at a low angle below the Eurasian Plate [Berberian, 1983; Jackson et al., 2002; Allen et al., 2002]. Seismic and gravity data suggest southward underthrusting of this oceanic lithosphere beneath the continental crust of the Talesh Mountains [Berberian, 1983; Jackson et al., 2002; Allen et al., 2003; Granath et al., 2007; Aziz Zanjan et al., 2013].

At present, the Talesh Mountains are influenced by the compressional stresses transferred from the Arabia-Eurasia collision zone (Figure 1). GPS network measurements suggest 5–6 mm/yr of shortening and 4–5 mm/yr of right-lateral strike-slip motion along the eastern boundary of the southern and central Talesh Mountains [Djamour et al., 2010, 2011]. In contrast, NNE extension perpendicular to the mountains at a rate of ~4 mm/yr has been documented for the northern Talesh Mountains [Masson et al., 2006]. Based on new GPS data, this extension rate has been revised to <2 mm/yr and a present-day clockwise rotation of the SCB with respect to the Talesh Mountains has been suggested (Figure 2a) [Djamour et al., 2010, 2011]. Patterns of seismicity indicate a partitioning of oblique strain into strike-slip and dip-slip faults at the western corner of the SCB and the Talesh Mountains [e.g., Jackson et al., 2002; Aziz Zanjan et al., 2013].

3 Materials and Methods

We collected 23 bedrock samples from clastic and volcanosedimentary units of the upper Paleozoic to lower Eocene rock units across the Talesh Mountains. Apatite and zircon grains were separated by crushing, sieving, and standard heavy liquid and magnetic separation. Euhedral apatites and zircons were picked using a cross-polarized binocular microscope. Most of the grains had a minimum diameter of 90 μm and were inclusion free to avoid effects of He implantation from inclusions or excess loss of He during decay due to a large surface/volume ratio [Farley, 2000]. The grain dimensions were measured for the calculation of the alpha-ejection (Ft) correction factor after Farley et al. [1996], and single grains were packed in Nb tubes for U-Th-Sm/He measurement. The ⁴He concentration was measured in a Patterson Instruments gas extraction line at the Tübingen University, Germany. Apatite samples were heated for 5 min at 11 A with a 960 nm diode laser for degassing, and zircon grains were heated for 20 min at 20 A. Each sample was reheated and measured to ensure that all gas was extracted in the first run. After degassing, the U, Th, and Sm were measured after dissolution in HNO₃ and HF. U, Th, and Sm measurements were conducted using an inductively coupled plasma mass spectrometer at the University of Arizona, USA. In general, we measured three single-grain aliquots per sample and reported the mean age and standard deviation as the sample age (Tables 1 and 2).

Table 1. Summary of the Apatite U-Th/He Cooling Agesa

Transect	Sample	Latitude N Longitude E	Rock Unit-Age	Elevation (m)	Single Grain	He₄ (mol)	U₂₃₈ (mol)	U₂₃₅ (mol)	Th₂₃₂ (mol)	Sm₁₄₇ (mol)	Ft	r (μm)	Single Grain Age (Ma)	Mean Age (Ma)	SD (Ma)
Southern	10TAL	N37.73922 E48.54901	Volcaniclastics-Eocene	1897	1	7.72E−14	1.17E−12	8.67E−15	4.03E−12	2.69E−12	0.844	170	33.61	34.69	1.44
					2	8.31E−14	1.20E−12	8.89E−15	4.39E−12	2.95E−12	0.849	176	34.14
					3	6.98E−14	1.01E−12	7.46E−15	3.30E−12	2.34E−12	0.839	172	36.32
	32TAL	N37.27673 E48.85893	Sandstone-Jurassic	2274	1	9.15E−16	1.05E−13	7.76E−16	3.58E−13	4.09E−13	0.75	106	5.01	4.38	0.55
					2	6.40E−16	8.63E−14	6.39E−16	3.47E−13	3.40E−13	0.728	103	4.08
					3	5.38E−16	7.56E−14	5.59E−16	2.59E−13	3.67E−13	0.753	110	4.06
	37TAL	N37.43521 E48.95963	Volcaniclastics-Cretaceous	628	1	8.65E−15	1.89E−13	1.40E−15	5.37E−13	3.67E−12	0.827	154	24.49	25.06	4.32
					2	1.72E−14	3.10E−13	2.29E−15	8.28E−13	5.28E−12	0.854	189	29.64
					3	6.34E−15	1.14E−13	8.46E−16	6.66E−13	1.94E−12	0.844	179	21.05
	39TAL	N37.60636 E48.67712	Sandstone-Cretaceous	2432	1	2.13E−15	1.32E−13	9.77E−16	1.04E−12	6.82E−13	0.720	110	6.14	6.47	1.65
					2	5.60E−16	6.60E−14	4.88E−16	2.32E−13	2.48E−13	0.720	93	5.01
					3	3.12E−15	1.59E−13	1.17E−15	9.90E−13	9.02E−13	0.753	113	8.26
	42TAL	N37.79523 E48.58476	Sandstone-Cretaceous	2371	1	5.17E−15	2.87E−13	2.12E−15	8.03E−13	1.07E−12	0.707	90	11.90	11.73	0.24
					2	4.98E−14	1.09E−12	8.07E−15	3.14E−12	2.52E−12	0.76	105	27.84
					3	5.97E−15	3.16E−13	2.33E−15	9.18E−13	1.18E−12	0.753	107	11.56
	45TAL	N37.62331 E48.76861	Volcaniclastics-Cretaceous	1856	1	3.26E−16	4.24E−14	3.13E−16	6.55E−14	5.54E−13	0.766	113	5.48	9.74	3.77
					2	3.78E−15	1.72E−13	1.27E−15	6.98E−13	8.62E−13	0.786	129	11.10
					3	3.31E−15	1.60E−13	1.18E−15	4.84E−13	8.61E−13	0.739	97	12.64
Central	5TAL	N38.03329 E48.69861	Volcaniclastics-Eocene	2303	1	3.54E−14	4.29E−13	3.17E−15	1.52E−12	1.45E−12	0.819	14	42.64	34.27	7.79
					2	1.05E−14	2.69E−13	1.99E−15	4.85E−13	4.95E−13	0.783	129	27.22
					3	1.53E−14	3.08E−13	2.28E−15	6.63E−13	6.36E−13	0.779	115	32.95
	11TAL	N38.04627 E48.66199	Volcaniclastics-Eocene	2625	1	7.46E−14	9.00E−13	6.66E−15	3.08E−12	2.86E−12	0.852	173	41.81	42.14	2.04
					2	6.26E−14	7.68E−13	5.68E−15	2.69E−12	2.45E−12	0.862	194	40.28
					3	5.14E−14	5.96E−13	4.41E−15	2.11E−12	1.91E−12	0.826	151	44.33
	14TAL	N38.10679 E48.64787	Volcaniclastics-Eocene	2537	1	1.36E−14	2.11E−13	1.56E−15	6.78E−13	5.55E−13	0.749	102	38.18	37.94	0.22
					2	1.16E−14	1.87E−13	1.39E−15	5.88E−13	4.82E−13	0.735	96	37.75
					3	1.32E−14	2.03E−13	1.50E−15	6.45E−13	5.89E−13	0.764	115	37.88
	29TAL	N38.02462 E48.81997	Volcaniclastics-Cretaceous	450	1	1.54E−14	6.63E−13	4.90E−15	2.94E−12	3.20E−12	0.841	166	10.54	8.35	3.09
					2	3.53E−15	2.79E−13	2.06E−15	1.15E−12	1.36E−12	0.810	149	6.17
					3	-------------	-------------	-------------	-------------	-------------	-------------	-------------	-------------
	35TAL	N37.96492 E48.59136	Volcaniclastics-Eocene	3151	1	3.09E−14	3.37E−13	2.49E−15	1.32E−12	1.13E−12	0.815	143	45.50	43.36	2.0
					2	2.48E−14	3.30E−13	2.44E−15	1.09E−12	9.62E−13	0.792	121	41.55
					3	2.26E−14	2.84E−13	2.10E−15	1.01E−12	8.41E−13	0.784	121	43.03
	44TAL	N37.97598 E48.77746	Volcaniclastics-Cretaceous	652	1	1.55E−15	5.75E−14	4.26E−16	1.25E−13	3.18E−13	0.697	95	19.68	18.03	1.48
					2	2.27E−15	4.51E−14	3.33E−16	4.77E−13	3.58E−13	0.673	90	16.82
					3	7.90E−16	3.28E−14	2.42E−16	7.27E−14	4.59E−13	0.673	83	17.59
Northern	4TAL	N39.09839 E47.73321	Volcaniclastics-Oligocene	792	1	2.22E−14	8.26E−13	6.11E−15	4.85E−12	2.76E−12	0.82	155	10.78	9.51	1.23
					2	1.05E−14	4.98E−13	3.68E−15	2.82E−12	1.66E−12	0.847	171	8.33
					3	1.18E−14	5.17E−13	3.82E−15	3.03E−12	1.80E−12	0.8	138	9.41
	8TAL	N38.96929 E47.80195	Sandstone-Eocene	1850	1	1.27E−14	2.48E−13	1.83E−15	8.86E−13	8.31E−13	0.804	136	26.98	29.02	2.25
					2	1.61E−14	3.08E−13	2.28E−15	1.02E−12	8.68E−13	0.793	147	28.81
					3	1.69E−14	2.61E−13	1.93E−15	1.11E−12	9.63E−13	0.808	137	31.27
	17TAL	N38.96392 E48.03696	Volcaniclastics-Eocene	1721	1	8.43E−15	2.70E−13	2.00E−15	2.83E−12	2.13E−12	0.745	115	9.48	14.44	5.06
					2	9.47E−15	2.11E−13	1.56E−15	2.11E−12	2.11E−12	0.734	104	14.24
					3	9.93E−15	1.99E−13	1.47E−15	1.43E−12	9.43E−13	0.742	106	19.60
	18TAL	N39.00624 E48.01198	Sandstone-Paleocene	1137	1	7.32E−15	3.61E−13	2.67E−15	2.92E−12	4.11E−12	0.817	147	6.63	12.17	5.50
					2	1.41E−14	3.48E−13	2.57E−15	1.83E−12	1.85E−12	0.801	144	17.62
					3	1.23E−14	4.61E−13	3.41E−15	2.35E−12	1.45E−12	0.776	125	12.28
	26TAL	N39.01882 E48.06858	Sandstone-Oligocene	904	1	9.73E−15	6.29E−13	4.65E−15	2.93E−12	1.86E−12	0.791	128	7.29	7.29	0.80
					2	1.37E−14	7.61E−13	5.63E−15	3.68E−12	2.65E−12	0.812	145	8.10
					3	8.61E−15	4.52E−13	3.35E−15	3.70E−12	3.34E−12	0.783	128	6.50
	27TAL	N39.18098 E48.06548	Volcaniclastics-Oligocene	521	1	6.10E−16	1.42E−13	1.05E−15	3.35E−13	4.52E−13	0.737	104	2.90	4.04	1.14
					2	1.43E−15	1.67E−13	1.23E−15	4.70E−13	7.23E−13	0.770	114	5.17
					3	1.27E−15	1.93E−13	1.43E−15	5.38E−13	7.18E−13	0.761	114	4.05
	28TAL	N39.01959 E47.92275	Sandstone-Oligocene	995	1	6.94E−15	4.00E−13	2.96E−15	1.81E−12	1.96E−12	0.811	144	8.07	9.49	3.88
					2	2.95E−14	1.05E−12	7.78E−15	4.31E−12	2.42E−12	0.803	137	13.88
					3	6.94E−15	4.90E−13	3.62E−15	2.14E−12	2.35E−12	0.834	172	6.53

^a Corrected age is corrected for alpha ejection after Farley et al. [1996]. Ft: alpha correction factor, r: crystal width. The U and Th concentrations in ppm were estimated by assuming a 3.2 g/cm³ apatite density, cylinder-shaped grain geometry, and using the measured grain dimensions.

Table 2. Summary of the Zircon U-Th/He Cooling Agesa

Transect	Sample	Latitude N	Rock Unit-Age	Elevation	Single Grain	He₄	U₂₃₈	U₂₃₅	Th₂₃₂	Ft	r (μm)	Single Grain Age (Ma)	Mean Age (Ma)	SD (Ma)
Transect	Sample	Latitude N	Rock Unit-Age	(m)	Single Grain	(mol)	(mol)	(mol)	(mol)	Ft	r (μm)	Single Grain Age (Ma)	Mean Age (Ma)	SD (Ma)
Southern	23TAL	N37.60977 E48.6295	Sandstone-Cretaceous	2353	1	9.6955E−12	4.5277E−11	3.3488E−13	1.4455E−11	0.829	128.7	183.37	189.69	7.85
					2	4.531E−12	1.924E−11	1.4231E−13	1.5523E−11	0.81	109.8	187.22
					3	6.3272E−12	2.8085E−11	2.0773E−13	1.2509E−11	0.784	93.5	198.48
	42TAL	N37.79523 E48.58476	Sandstone-Cretaceous	2371	1	2.9087E−12	1.0016E−11	7.4079E−14	5.2083E−12	0.785	99.9	250.57	197.95	49.90
					2	5.9255E−12	2.6366E−11	1.9502E−13	1.4508E−11	0.792	102	192.00
					3	5.4686E−12	3.2743E−11	2.4218E−13	1.2043E−11	0.778	94.9	151.29
Central	20TAL	N37.99300 E48.77917	Granite-Jurassic	1320	1	9.0728E−13	6.4505E−12	4.771E−14	4.2787E−12	0.813	125.2	115.23	123.77	9.82
					2	7.8374E−13	4.773E−12	3.5303E−14	4.1916E−12	0.779	94.9	134.50
					3	9.8187E−13	6.6557E−12	4.9228E−14	4.9612E−12	0.795	94.6	121.57
	7TAL	N37.96120 E48.76178	Granite-Jurassic	871	1	8.0433E−13	4.5937E−12	3.3977E−14	4.5078E−12	0.765	90.7	143.17	160.10	28.01
					2	1.6493E−12	6.8956E−12	5.1003E−14	5.5218E−12	0.801	106.9	192.42
					3	8.2908E−13	4.7071E−12	3.4816E−14	4.1695E−12	0.775	92.8	144.69

^a Ft: alpha correction factor, r: crystal width.

Our geologic maps and cross sections are based on detailed structural and stratigraphic field analyses and existing reconnaissance maps and reports. All new mapping was conducted at a scale of 1:50,000. The kinematics of major faults was investigated through a detailed analysis of the fault zone structures, including S-C structures, drag folds, and slickensides. The restored cross sections were constructed using the sinuous bed method, in which bed line lengths and areas are conserved. Five cross sections are constructed located in the southern (A-A´ and B-B´ transects), central (C-C´ transect), and wnorthern (D-D´ transect) Talesh Mountains (Figures 2b and 4-7). The bedrock cooling ages are interpreted in the context of our geologic maps and balanced cross sections that allow estimation of the timing, rates, and spatial pattern of rock exhumation associated with specific phases of deformation. The stratigraphic sequences of these cross sections provide independent estimates on the magnitude of exhumation. In order to determine changes in the exhumation history of each transect, we plotted our sample ages against their structural depths, which was determined through the restoration of the cross sections (Figures 5-8). The apparent exhumation rates were calculated for each region based on the slope or changes in the slope of a regression through the age versus structural depth relationships (Figure 9). We note that apparent exhumation rates from bedrock thermochronometer data do not always reflect true rates [e.g., Ehlers et al., 2001; Ehlers, 2005]. However, a focused modeling study of exhumation rates on these sections [e.g., McQuarrie and Ehlers, 2015] is beyond the scope of this study. Removal of the material along strike-slip faults is one of the main points that could affect shortening estimates in the balanced cross sections. The BD and MD faults represent strike-slip movement that could affect our estimate on shortening and related exhumation in the central part of the Talesh Mountains. We note that we do not account for any variations in stratigraphic thickness. Thickness variations may have existed in the sections due to erosion. This is especially the case for syntectonic Paleogene and Eocene deposits that vary in thickness.

4 Results

We present the AHe and ZHe cooling ages from three major transects that cross the southern, central, and northern parts of the Talesh Mountains (Figures 2b and 5-8). Generally, our single-grain data reproduce well and the mean age of the samples is used for our interpretation of cooling during exhumation (Tables 1 and 2). Some of our cooling ages, especially in the case of ZHe, are unreset or partially reset and are not considered for constraining the timing of exhumation but instead used to infer maximum burial and exhumation magnitudes.

4.1 Southern Transect

The NE trending MD and BD faults are the main structural features in the southern part of the Talesh Mountains (Figure 2b). The BD and MD faults, measured in three main transects, are steeply dipping (50°–60°) to the northeast (Figures 4a and 4b). Fault plane striation analysis shows westward thrusting of Paleozoic rock units over Mesozoic-Cenozoic deposits with a right-lateral strike-slip component (Figures 4a, 4d, and 4e).

Based on detailed field measurements and previously published maps and reports, the A-A′ and B-B′ cross sections are constructed and restored along the southern transect and represent a total of 20–27 km (25–30%) shortening in the late Cenozoic (Figures 5 and 6) [Madanipour et al., 2013]. We sampled the upper Paleozoic to lower Cenozoic rock units from the deepest structural level exposed near the cross-section lines in the hanging wall of the MD and BD faults (Figure 2).

The structural geometry along the A-A′ section was reconstructed above the topographic level to estimate the amount of material removed by erosion (Figure 5). The restored balanced section suggests at least ~4–10 km of exhumation in this part of the Talesh Mountains (Figure 5). The Neogene apatite helium partial retention zone (PRZ; ~55–75°C) is also estimated in the eroded part of the cross section by assuming 25°C/km of geothermal gradient (gray bar in Figure 5). Samples 32TAL and 37TAL were collected from Jurassic sandstone and Cretaceous volcaniclastic rocks in the hanging walls of the MD and BD faults, respectively, yielding 4.38 ± 0.55 Ma and 25.06 ± 4.32 Ma AHe cooling ages, respectively (Figure 5).

Most of the samples in the southern Talesh Mountains are located along the B-B′ cross section and in the hanging wall of the BD Fault. These samples suggest at least ~3–11 km of exhumation (Figure 6) [Madanipour et al., 2013]. The apatite and zircon PRZ was constructed in the exhumed part of the cross section and indicates reset AHe ages but unreset or partially reset ZHe ages (Figure 6). Samples 23TAL and 42TAL are from Cretaceous sandstones and conglomerates, yielding single-grain ZHe ages that range from 151 Ma to 250 Ma and are thus older than their stratigraphic ages (Figure 6 and Table 2). Samples 10TAL, 39TAL, 42TAL, and 45TAL are from Cretaceous and Eocene clastic and volcaniclastic rock units and yield AHe ages of 34.7 ± 1.44 Ma, 6.47 ± 1.65 Ma, 11.73 ± 0.24 Ma, and 9.74 ± 3.37 Ma, respectively (Figure 6 and Table 1).

4.2 Central Transect

In the north trending central part of the Talesh Mountains, the trend in the MD and BD faults changes to N-S and then terminates at the Talesh Fault in the Caspian coastal plains (Figure 2). Fault plane striation analysis on MD and BD faults indicates right-lateral strike-slip kinematics (Figures 4c and 4f). Lower Cretaceous rock units are the deepest exposed structural levels in the hanging wall of the MD and BD faults in the central Talesh Mountains (Figure 2). However, in the southern Talesh Mountains, lower Paleozoic units are the deepest exposed structural levels in the hanging wall. This might also reflect a general kinematic change of the faults from dominantly thrust motion in south to strike-slip motion (with lower denudation) in the central part of the mountains.

The central part of the Talesh Mountains has experienced a total ~6–11 km (~13%–16%) of late Cenozoic shortening based on the restored C-C′ cross section. The MD and BD faults accommodated a lower amount of the vertical separation. Lower Cretaceous rocks are the deepest exposed structural level in the hanging wall of the BD and MD faults and were sampled for thermochronometric analyses (Figures 2b and 7). All samples located along the C-C′ cross section used to reconstruct the structural depth suggest at least ~2–9 km of exhumation (Figure 7). The constructed apatite and zircon Neogene PRZ in the exhumed part of the cross section predicts reset AHe ages and unreset or partially reset ZHe ages (Figure 7).

Samples 7TAL and 20TAL were collected from the Lisar granitic body located south of the C-C′ cross section (Figure 2). Clasts of this granitic body are observed in the Cretaceous conglomerate and yielded a zircon U-Pb age of 179 ± 18 Ma. Samples 7TAL and 20TAL have ZHe cooling ages of 160.1 ± 28 Ma and 123.8 ± 9.3 Ma, respectively (Figure 7 and Table 2). Proper apatite crystals for AHe dating were not found in this granitic body.

Six samples from Middle-Late Cretaceous to early-late Eocene clastic and volcaniclastic sedimentary rocks were used for AHe analysis. Samples 35TAL and 11TAL are from the late Eocene andesitic volcanics in the hanging wall of the MD Fault, yielding AHe cooling ages that are equal to or slightly younger than crystallization ages (43.36 ± 2.0 Ma and 42.14 ± 2.04 Ma AHe ages). These results are in agreement with the restored cross section that predicts the least amount of exhumation in the western part of the cross section (Figure 7). Samples 14TAL and 5TAL with the stratigraphic age of the early-middle Eocene yield AHe cooling ages of 37.94 ± 0.22 Ma and 34.27 ± 7.79 Ma, respectively (Figure 7 and Table 1). Samples 44TAL and 29TAL are from Middle-Late Cretaceous vocaniclastic rocks in the footwall and hanging wall of the BD and Talesh faults, respectively (Figure 7). Their 18.03 ± 1.48 Ma and 8.35 ± 3.39 Ma AHe ages record the youngest cooling in the central transect (Figure 7 and Table 1). In general, the AHe cooling ages become older from low to higher elevation and from east to west along the central transect (Figure 7).

4.3 Northern Transect

Kinematic data from the fault zone, in addition to interpreted 2-D seismic sections, indicate normal sense kinematics during Paleocene-Eocene regional extension. Following this, the fault was reactivated as a high-angle reverse fault during late Cenozoic compression [Madanipour et al., 2013]. The east trending northern part of the Talesh Mountains is mainly composed of Cenozoic rocks and experienced a total of ~20 km (~22%) shortening in the late Cenozoic based on the restored D-D′ cross section (Figures 2b and 8). The south dipping Angut Fault thrusts the Mesozoic and lower Cenozoic (Paleocene-Eocene) volcanic rocks over the younger (Oligocene-Miocene) sedimentary sequences (Figures 2b and 8). The Oligocene-Miocene stratigraphic sequences generally thin toward the footwall of the Angut Fault [Geological Survey of Iran, 1996]. The deformation along this part of the mountains was accommodated along two different horizons as the Oligocene rock units on the top and the lower to upper Eocene rocks at the bottom (Figure 8). The folding style in the rock units above the upper Eocene Eml unit is different from that in the underlying lower Eocene and older sequences (Figure 8). All the samples along the D-D′ cross section were collected from the lower-middle Cenozoic clastic and volcaniclastic rock units (Figures 2b and 8). The balanced section predicts at least ~2–7 km of late Cenozoic exhumation along this section, and the constructed apatite PRZ in the exhumed part of the cross section predicts reset and partially reset AHe cooling ages (Figure 8).

Sample 27TAL from early Eocene volcaniclastic rocks in the core of Ojagh Gheshlagh Anticline revealed the youngest (4.04 ± 1.14 Ma) AHe cooling age from our data set (Figure 8 and Table 1). Samples 4TAL, 26TAL, 18TAL, and 28TAL were collected from late Eocene to early Oligocene clastic rocks between the Angut Fault and the Ojagh Gheshlagh Anticline and yield AHe ages of 9.51 ± 1.23 Ma, 7.29 ± 0.80 Ma, 12.17 ± 5.5 Ma, and 9.49 ± 3.38 Ma, respectively (Figure 8 and Table 1). The ages located in the footwall of the Angut Fault are all younger than their stratigraphic age, whereby the two southern samples (28TAL and 18TAL) show a large variation in the single-grain ages in contrast to the two northern samples (4TAL and 26TAL) (Figure 8). Samples 8TAL and 17TAL are from the lower to middle Eocene volcanic and volcaniclastic rocks in the hanging wall of the Angut Fault and yield 29.02 ± 2.15 Ma and 14.44 ± 5.06 Ma AHe ages, respectively (Figures 2b and 8). Early Oligocene to late Miocene stratigraphic units have not been deposited on the hanging wall of the Angut Fault suggesting a lower structural level for the 8TAL and 17TAL samples (Figure 8) [Geological Survey of Iran, 1996].

5 Discussion

Our new thermochronometer data, in combination with previously published detrital apatite fission track analyses [Madanipour et al., 2013], structural geometry, and kinematic data allow us to quantify spatial variations in the deformation and exhumation history of the Talesh Mountains.

5.1 Late Oligocene (~27–23 Ma) Slow Cooling and Deformation

The last phase of significant extension at the northern margin of the Arabia-Eurasia collision zone occurred during the Eocene and was associated with widespread volcanic activity and volcaniclastic deposition in a rapidly subsiding basin [Vincent et al., 2005; Ballato et al., 2011; Verdel et al., 2011]. Previous studies of the basin suggested that a transition from extension to compression occurred at the Eocene-Oligocene boundary [Vincent et al., 2005, 2007; Guest et al., 2007; Ballato et al., 2008]. On the Azeri side of the northern Talesh Mountains, the timing of the transition from rifting and volcanism to compression is uncertain. Fine-grained turbidity current and hemipelagic sediments with slope instability features recorded in upper Eocene-lower Oligocene strata attributed this regional tectonic regime change. However, this might also result from other agents like climate change, increased subsidence rate, and change in depositional environment rather than a change in regional tectonics [Vincent et al., 2005].

Detrital apatite fission track ages reveal dominant age populations in the entire Talesh Mountains that peak in the late Oligocene (~27–23 Ma, >50% of grains; Figure 9). While the age distributions from the northern and southern Talesh Mountains yield also younger (late Miocene) and older age populations, all 554 grains dated from the central Talesh Mountains fall into the late Oligocene age population (Figure 9d). We suggest that cooling and exhumation in the entire Talesh Mountains initiated in the late Oligocene and thus later than the previously proposed Eocene-Oligocene boundary. Apparent exhumation rates are relatively slow in the late Oligocene and ~0.15 ± 0.03 km/Myr in the central Talesh and even lower and at ~0.03 ± 0.01 km/Myr and ~0.07 ± 0.01 km/Myr in the northern and southern Talesh, respectively (Figure 9).

Late Oligocene cooling has been documented in other parts of the Arabia-Eurasia collision zone. Thermal modeling of the AHe, ZHe, K-feldspar, and biotite ⁴⁰Ar/³⁹Ar data from granitic bodies in the central Alborz Mountains suggests an onset of slow cooling at ~25 Ma [Axen et al., 2001]. The cooling history of the intrusive bodies and Paleozoic rock units in the Zagros Mountains, where the Arabian and Eurasian Plates collided, also record an onset of slow cooling during the late Oligocene at ~25–23 Ma [Okay et al., 2010; Gavillot et al., 2010; Francois et al., 2014]. In the Greater Caucasus Mountains, located at the northwesternmost margin of the Arabia-Eurasia collision zone, AFT, AHe, and ZHe thermochronometry studies record late Oligocene ages (~30–25 Ma) and the initiation of slow cooling [Vincent et al., 2007, 2010; Avdeev and Niemi, 2011].

Different thermochronometric cooling ages observed in the hanging wall of thrust faults might relate to displacement variations and corresponding erosional exhumation variations along the length of a fault [e.g., McQuarrie and Ehlers, 2015]. As a result, different samples from a certain stratigraphic depth might show differential rates of exhumation along hanging wall of thrust faults. The exhumation and cooling ages of the rock units in the hanging wall of thrust faults can be used as direct evidence for fault movement and reconstruction of the deformation rate and history in thrust belts [McQuarrie and Ehlers, 2015].

The MD, BD, and Angut faults are the main structural features in the Talesh Mountains (Figures 2b and 5-7). The lower-middle Paleozoic rock units have been juxtaposed over Mesozoic and lower Cenozoic rocks along the MD and BD faults in the southern Talesh Mountains. However, their terminations show dominantly strike-slip movement with a minor reverse component in the central Talesh Mountains (Figures 2b and 5-7). The pattern of rock units exposed in the hanging wall and footwall of the MD and BD faults indicates that these faults could have acted as bounding normal faults during the early Mesozoic or late Paleocene-early Eocene extensional phases [Zancheta et al., 2009; Zanchi et al., 2009]. Evidence of early Eocene normal displacement has previously been observed on seismic images along the Angut Fault and Ojagh Gheshlagh pop-up anticline in the northern Talesh Mountains (Figures 2b, 8, and 10). We suggest Eocene extension on the Angut Fault in Figure 10 (top) based on the slow cooling and exhumation rates recorded across the Talesh Mountains (Figures 9a, 9c, and 9e). Following this, slow inversion of the MD, BD, and Angut faults occurred during the late Oligocene. The timing of accelerated exhumation in the middle Miocene represents a minimum age of fault reactivation because accelerated exhumation rates may have lagged behind fault activation. The late Oligocene inversion of the Angut Fault provided space for the late Oligocene-Miocene strata in the footwall of the fault in northern Talesh Mountains (Figure 10, top).

5.2 Rapid Middle Miocene (~12 Ma) Exhumation and Regional Kinematic Change

The middle Miocene (~12 Ma) age population peak found in the northern and southern Talesh Mountains correlates well with the inflection point in the bedrock U-Th/He ages versus structural depths plots (Figures 9a and 9e). Apparent exhumation rates accelerated above previous rates by a factor of 4 at ~12 Ma to 0.27 ± 0.07 km/Myr in the north and 0.31 ± 0.08 km/Myr in the south (Figures 9a, 9c, and 9e). This phase of rapid cooling is not recorded by the detrital or bedrock ages from samples in the central Talesh Mountains (Figures 9b, 9d, and 9f) that is dominated by strike-slip deformation and limited rock exhumation (Figures 8c and 8d) [Madanipour et al., 2013].

However, the middle Miocene cooling is a regional deformation event that probably affected most of the Arabia-Eurasia collision zone [Guest et al., 2007; Gavillot et al., 2010; Ballato et al., 2016]. Thermal history data from a Cretaceous granitic pluton in the central Alborz Mountains (AHe, ZHe, and K-feldspar ⁴⁰Ar/³⁹Ar) reveal slow denudation (~0.1 km/Myr) as late as ~12 Ma with more accelerated exhumation (~0.45 km/Myr) that likely began shortly after ~12 Ma [Guest et al., 2006a, 2007]. Paleozoic rock units in the hanging wall of the High Zagros Fault in the southeastern part of the High Zagros Mountains also cooled rapidly in the early-middle Miocene [Gavillot et al., 2010]. We suggest that the middle Miocene phase of deformation and cooling was widespread across the northern and southern boundaries of the Iranian Plateau and constitutes the collision between the Arabian-Eurasian Plates.

The middle Miocene phase of deformation is also documented in sedimentary sequences of other parts of the Arabia-Eurasia collision zone including the foreland of the Alborz and Talesh mountains and central, northeast, and northwest Iranian basins [Ballato et al., 2008; Morley et al., 2009; Madanipour et al., 2013; Ballato et al., 2016]. Recent 2-D and 3-D seismic data document that in the central Iranian basin some of the major massif-bounding thrusts active in the late Oligocene as normal or transtensional faults subsequently inverted during the middle Miocene [Morley et al., 2009]. The sediment accumulation rate of the Upper Red Formation in the foreland of the Alborz Mountains increased around 13–10 Ma [Ballato et al., 2008]. In the Talesh Mountains the observed unconformity between the lower Miocene and the middle Miocene sequences provides evidence for uplift and deformation. This is also supported by the interpretation of 2-D seismic lines that indicate middle and upper Miocene growth strata in the flanks of major folds in the northern Talesh Mountains [Madanipour et al., 2013]. In northeast Iran, an ~18 Ma faulting/exhumation episode is chiefly recorded by the structure and depositional architecture of Neogene deposits along major strike-slip faults [Calzolari et al., 2016]. Structural, geochronological, and sedimentological data suggest that shortening and thickening led to the outward and vertical growth of the northern sectors of the Iranian Plateau starting from the middle Miocene [Ballato et al., 2016].

In a compressional tectonic setting, an increase in apparent exhumation rate is generally assumed to be an acceleration of rock uplift and erosion related to the upward motion of the hanging wall on a thrust fault [Wagner and Reimer, 1972; Wagner et al., 1977; Fitzgerald et al., 1995; Ehlers, 2005; Reiners and Brandon, 2006]. Since most of our samples are collected from the hanging wall of the MD and BD and Ojagh Gheshlagh faults, we can infer that an increase in the rate of thrusting occurred during the middle to late Miocene deformation in the southern and northern Talesh Mountains, respectively. In the northern Talesh Mountains, the increase in the rate of thrusting might have led to a reactivation of the Ojagh Gheshlagh Fault through an upper Eocene-lower Oligocene detachment. Motion on this detachment resulted in the forward propagation of the deformation into the Kura Basin (Figure 10, bottom). The reactivation of the Ojagh Gheshlagh Fault exposed the lower Eocene clastic and volcaniclastic rock units in a pop-up structure in the para-Tethys side of the Talesh Mountain foreland (Figure 10, bottom). However, in the central Talesh Mountains, strike-slip kinematics of the faults led to very limited amount of exhumation.

Along-strike variation in the regional compressional directions in collision zones is one of the major processes that explain kinematic variations in orogenic belts [e.g., Tapponnier et al., 1982, 1986]. The direction of Arabia-Eurasia convergence is thought to be northeast between ~56 and ~25 Ma and changed to north directed since ~25 Ma when collision initiated [McQuarrie et al., 2003; McQuarrie and van Hinsbergen, 2013]. McQuarrie et al. [2003] proposed that the rate of convergence decreased at ~10 Ma, and reorganization in the compression direction is suggested to have occurred at ~7–2 Ma in the Alborz Mountains [Allen et al., 2003; Allen, 2010; Ballato et al., 2013]. The spatial and temporal pattern of detrital apatite fission track cooling ages across the Talesh Mountains has led to the suggestion of a regional change in the convergence direction during the middle Miocene [Madanipour et al., 2013]. The NE directed early Oligocene stress field resulted in oblique thrust deformation and exhumation that is recorded across the Talesh Mountain, whereas the north directed stress field favored thrust and reverse faulting in the northern and southern Talesh Mountains, and dextral strike-slip faulting in the central Talesh Mountains since ~12 Ma. This change in fault kinematics is recorded in the thermochronology data that reveal accelerated exhumation rates in the northern and southern Talesh since 12 Ma and very limited exhumation in the central part.

5.3 Late Cenozoic Oroclinal Bending of the Talesh Mountains

Along-strike variations in the structural and topographic trend are common in compressional orogens and referred to as oroclines. These features form in response to the combination of different tectonic and sedimentary processes [Marshak and Flöttmann, 1996; Paulsen and Marshak, 1999; Marshak, 2004; Yonkee and Weil, 2010; Weil et al., 2001, 2010]. One common cause for oroclinal bending is warping of an orogen around the corner of a rigid basement block [Marshak and Flöttmann, 1996; Marshak, 2004]. The Talesh, Lesser Caucasus, central and eastern Alborz, and Kopeh Dagh mountains form together a continuous orogenic belt around the southern margin of the South Caspian Block (Figure 11). This orogenic belt might have formed through the collision of the rigid central Iranian Block with the smaller rigid South Caspian Block which then resulted in the wrapping of the Talesh and Alborz mountain ranges around the South Caspian Block (Figure 11) [e.g., Berberian, 1983; Allen et al., 2003]. The Alborz have been suggested as a linear mountain range during the early-middle Oligocene, which subsequently deformed during the Miocene when uplift of the Kopeh Dagh Basin formed the Kopeh Dagh Mountains and continuous shortening resulted in the northward deflection of the eastern Alborz Mountains [Hollingsworth et al., 2010]. A new study by Cifelli et al. [2015] suggests a straight E-W structural trend of the central Alborz Mountains from ~17 Ma to ~9 Ma. Paleomagnetic data from the middle to late Miocene foreland strata suggest that the WNW trending western Alborz Mountains is caused by clockwise rotation, whereas the ENE trending eastern Alborz Mountains is caused by counterclockwise rotation that occurred since ~7.6 Ma [Cifelli et al., 2015]. Paleomagnetic studies also suggest that ~50% of the curvature of the Lesser Caucasus Mountains formed after the Eocene and before the late Miocene, probably as a result of Arabia-Eurasia collision [Meijers et al., 2015].

Previous studies suggest that the oroclinal bending of the Talesh Mountains resulted from the interaction with the South Caspian Block or the effect of right-lateral strike-slip kinematics of the Talesh and West Caspian faults at the western border of the South Caspian Block (Figures 1 and 2b) [Berberian, 1983; Jackson et al., 2002; Allen et al., 2003, Hollingsworth et al., 2010]. Limited paleomagnetic data from basalts within the northern Talesh Mountains suggest that ~26° counterclockwise rotation took place during the early middle Eocene and 18–20° of clockwise rotation has also occurred since the late Eocene [Khalafly, 2001]. This supports a period of major bending of the Talesh Mountains during the late Cenozoic.

The spatial distribution of detrital and bedrock cooling ages constrains the temporal pattern of oroclinal bending in the Talesh Mountains. The observed late Oligocene (27–23 Ma) cooling phase is recorded across the entire Talesh Mountains, but younger exhumation phases (of middle Miocene and early Pliocene 5 Ma) are only documented in the southern and northern parts of the Talesh Mountains (Figures 9a, 9c, and 9e). During the early Oligocene, the entire Talesh Mountains were east trending under a NE directed compressional stress field. The subsequent gradual bending of the Talesh Mountains caused a change in the trend of the central Talesh to an approximately N-S orientation and division of structural trends into the northern, central, and southern Talesh Mountain components. This change led to post middle Miocene enhanced uplift and exhumation in the southern and northern Talesh that are oriented perpendicular to the north directed compressional direction and strike-slip motion in the central Talesh. Based on these observations, we propose a post late Oligocene onset and middle Miocene acceleration for oroclinal bending of the Talesh Mountains (Figure 11), which is consistent with the proposed bending for the eastern Alborz and Kopeh Dagh Mountains [Hollingsworth et al., 2010] and central Alborz [Cifelli et al., 2015]. The final bending of the central Alborz Mountains during the late Miocene formed the present configuration of the curved northern Iranian Plateau margin around the South Caspian Block (Figure 11).

5.4 Implications for the Arabia-Eurasia Collision

Our results of the cooling history and deformation of the Talesh Mountains have implications for lithospheric evolution and deformation processes across the Arabia-Eurasia collision zone. The Arabia-Eurasia collision is generally thought to have initiated with a transition from extension to contraction and the termination of arc magmatism at the Eocene-Oligocene boundary at ~36 Ma [Berberian, 1983; Allen et al., 2003; Guest et al., 2006a, 2006b, 2007; Allen and Armstrong, 2008; Yassaghi and Madanipour, 2008; Ballato et al., 2011, Karagaranbafghi et al., 2011; Mouthereau et al., 2012; Rezaeian et al., 2012; Madanipour et al., 2013]. This initial stage of the continental collision attributed to a soft collision in which (passive margin) continental Arabian lithosphere stretched beneath the central Iranian Block [Ballato et al., 2011]. However, plate kinematic reconstructions combined with documented shortening across the collision zone suggest that the initial collision occurred after ~27 Ma [McQuarrie and Van Hinsbergen, 2013]. In the Talesh Mountains our bedrock and detrital cooling ages suggest ~27–23 Ma as the initiation of deformation (Figure 9). Although the transition from extension to compression occurred around ~36 Ma, this deformation was not strong enough to exhume deeper crustal rocks but resulted in reactivation and inversion of normal faults under the new compressional regime [Morley et al., 2009]. The slow rate of the post late Oligocene deformation phase gradually shortened the northern and southern margins of the collision zone. However, the central Iranian Plateau experienced overall subsidence related to the postrift cooling that led to the formation of the Qom seaway in which the Qom Formation was deposited [Guest et al., 2007; Morley et al., 2009].

The accelerated exhumation rates since the middle Miocene (~12 Ma) record a prominent deformation phase in the Talesh Mountains (Figures 9a, 9c, and 9e). The sedimentary, thermochronologic, petrologic, and structural effects of this deformation phase are recorded across the entire collision zone [McQuarrie et al., 2003; Allen et al., 2004; Guest et al., 2007; Ballato et al., 2008, 2011; Mouthereau, 2011, Mouthereau et al., 2012, Madanipour et al., 2013]. Crustal thickening is believed to have occurred during this phase of shortening [Ballato et al., 2013] and resulted in the 45–65 km thick Zagros and Sanandaj-Sirjan Zone [Paul et al., 2010; Motavalli-Anbaran et al., 2011; Jiménez-Munt et al., 2012] and the 35–40 km thick crust of the Alborz Mountains and central Iranian Plateau [Paul et al., 2010; Motavalli-Anbaran et al., 2011; Jiménez-Munt et al., 2012]. Regional uplift of the collision zone, especially in the central Iranian Plateau, was due to the uplift of hot asthenosphere in response to slab rollback and break off [Morley et al., 2009]. Regionally, this deformation phase is attributed to the hard collision, in which the stiff continental crusts of Arabia and Eurasia finally met resulting in crustal shortening and thickening coupled with lateral extrusion of crustal blocks that enhanced uplift and exhumation in the collision zone [McQuarrie et al., 2003; Guest et al., 2007; Morley et al., 2009; Hatzfeld and Molnar, 2010; Ballato et al., 2011]. It is believe that the lag time between the late Oligocene onset of continental collision and the middle Miocene acceleration of deformation is interpreted to reflect continental accretion following the subduction of stretched continental lithosphere in the Arabia- Eurasia collision zone [Ballato et al., 2011].

6 Conclusion

This study presents an integration of bedrock and detrital thermochronology data with structural analyses and restored cross sections to document the exhumation and deformation history of the Talesh Mountains at the northwestern margin of the Arabia-Eurasia collision zone. We suggest that the late Oligocene (~27–23 Ma) cooling and exhumation record the initiation of collision between the stretched continental lithosphere. This initial collision caused shortening that was accommodated along reactivated and/or newly initiated thrust faults including the MD, BD, and Angut faults. Continued convergence in a NE directed stress field resulted in oroclinal bending of the Talesh Mountains due to the progressive interaction with the southwestern corner of the rigid South Caspian Block. The final phase of hard continental collision started at ~12 Ma and resulted in accelerated exhumation. The compressional stress field changed to a northward orientation that resulted in exhumation along the thrust and reverse faults of the northern and southern Talesh and in a strike-slip-dominated regime with limited exhumation in the north trending central Talesh Mountains.

Acknowledgments

We acknowledge the financial support of the Tarbiat Modares University, Iran. We thank Chris Fergusson and an anonymous reviewer for their constructive comments on an earlier version of the manuscript. Claudio Faccenna and Bernard Guest are thanked for their critical comments and editorial handling of the manuscript. Data presented in this study can be obtained from the following link: http://www.rapidshare.com.cn/vzmRX7r.

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Citing Literature

Volume36, Issue8

August 2017

Pages 1538-1561

Accelerated middle Miocene exhumation of the Talesh Mountains constrained by U-Th/He thermochronometry: Evidence for the Arabia-Eurasia collision in the NW Iranian Plateau

Abstract

Key Points

1 Introduction

2 Geological Setting of the Talesh Mountains

3 Materials and Methods