# Pliocene growth of the Dowlatabad syncline in Frontal Fars arc: Folding propagation across the Zagros Fold Belt, Iran

Mahdi Najafi, Elisabet Beamud, Jonas Ruh, Frédéric Mouthereau, Alireza Tahmasbi, Gilen Bernaola, Ali Yassaghi, Hossein Motamedi, Shahram Sherkati, Mohammad Ghasem Hassan Goodarzi, et al.

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Mahdi Najafi, Elisabet Beamud, Jonas Ruh, Frédéric Mouthereau, Alireza Tahmasbi, et al. Pliocene growth of the Dowlatabad syncline in Frontal Fars arc: Folding propagation across the Zagros Fold Belt, Iran. Geological Society of America Bulletin, 2021, 133 (7-8), pp.1381-1403. 10.1130/B35748.1. hal-04903835

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Pliocene Growth of the Dowlatabad Syncline in Frontal Fars: 1

Folding Propagation across the Zagros Fold Belt 2

Mahdi Najafi¹*, Elisabet Beamud², Jonas Ruh³, Frédéric Mouthereau⁴, Alireza Tahmasbi⁵, 3

Gilen Bernaola⁶, Ali Yassaghi⁷, Hossein Motamedi⁵, Shahram Sherkati⁸, Mohammad Ghasem 4

Hassan Goodarzi⁵, Jaume Vergés⁹ 5

¹Department of Earth Sciences, Institute for Advanced Studies in Basic Sciences (IASBS), Zanjan, 6
Iran

²Paleomagnetic Laboratory CCiTUB-Geo3Bcn, CSIC, Lluís Solé i Sabarís, s/n, 08028 Barcelona, 8
Spain

³Structural Geology and Tectonics Group, Geological Institute, Department of Earth Sciences, 10
ETH Zurich, Sonneggstrasse 5, 8092 Zurich, Switzerland

⁴Geosciences Environnement Toulouse (GET), Université de Toulouse, UPS, Univ. Paul Sabatier, 12
CNRS, IRD, 14 av. Edouard Belin, F-31400 Toulouse, France

⁵National Iranian Oil Company, Exploration Directorate, PO Box 19395-6669, Tehran, Iran

⁶Department of Mining-Metallurgy Engineering and Material Science, Faculty of Engineering in 15
Bilbao, University of the Basque Country (UPV/EHU), Plaza Ingeniero Torres Quevedo 1, Bilbao, 16
Spain

⁷Department of Geology, Tarbiat Modares University, Tehran, Iran

⁸PETRONAS, Kuala Lumpur, Malaysia

⁹Group of Dynamics of the Lithosphere (GDL), Geosciences Barcelona, Geo3Bcn - CSIC, Lluís 20
Solé i Sabarís s/n, 08028 Barcelona, Spain

*Correspondence: mahdinajafi@iasbs.ac.ir 22
23 **ABSTRACT**

The integration of biostratigraphy, strontium isotope stratigraphy and magnetostratigraphy allowed to precisely date the >3.0-km-thick marine to non-marine foreland sedimentary succession within the Dowlatabad growth syncline along the Frontal Fars in the Zagros Fold Belt. This area was the missing link to complete the ages of syntectonic deposits in the Fars arc to be able to carry out regional quantified analyses of both the migration of sedimentary belts and of the propagation of folding across the Mesopotamian foreland basin. Both are essential for defining the interplay of basin evolution and sequence of folding.

Deposition of the foreland marine marls in the Mishan Fm. started about 11.5 Ma. The transition to a non-marine basin infill occurred at 4.9 Ma by the progradation of thick fluvial deposits of the Aghajari Fm. with a fast accumulation rate of 63 cm/kyr. The beginning of growth strata deposition and thus the onset of folding in the Dowlatabad syncline is dated at 4.65 Ma. The first appearance of carbonate conglomerates sourced from the Guri limestone at 2.8 Ma marked the progressive dismantling of the nearby growing anticlines. Tectonic deformation in the front of the Fars arc was active at least for 2.85 Myr and ceased at 1.8 Ma beneath the Bakhtyari unconformable conglomerates composed of external clasts derived from the Zagros hinterland.

The compilation of magnetostratigraphic ages reveals that both the migration of Aghajari-Bakhtyari sedimentary belts and propagation of folding front was in-sequence towards the foreland with rates close to 20 mm/yr in the Fars arc and to 15 mm/yr in the Lurestan arc in the last 20 Myr. These high rates of folding propagation are about one order of magnitude larger than age equivalent shortening rates (~4 mm/yr in Fars arc, and ~2 mm/yr in Lurestan arc) and thus implying an efficient detachment level at the base of deformed Arabian sedimentary cover.

Numerical experiments including cover and basement are designed to test the influence of inherited basement structures on the deformation propagation within the cover sequence providing clues on
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the partly coeval in-sequence deformation of the Zagros Simply Folded Belt and the local out-of-sequence Mountain Frontal Fault system as illustrated by regional and local geology.

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**KEY WORDS:** magnetostratigraphy, fold growth, folding propagation, numerical modelling, Zagros Fold Belt, Mesopotamian foreland basin

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## INTRODUCTION

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The Zagros Mountains form a large segment of the Alpine-Himalayan orogenic system between Arabia and Eurasia evolving during a long-lasting and multi-episodic orogeny by a combination of thin-skinned folding and basement-involved thrusting (Fig. 1). The first episode of orogenic deformation was the Late Cretaceous emplacement of the Neo-Tethys ophiolites on top of the Arabian passive margin, interpreted as a volcanic arc collision and resulted in a restricted early foreland flexure in front of the emplaced ophiolite slices (Ricou, 1971; Agard et al., 2005; Saura et al., 2011). The continental collision between Arabia and Eurasia initiated by a soft collision during the Oligocene, characterized by underthrusting of the distal portion of the continental margin (Ballato et al., 2011; Mouthereau et al., 2012; Madanipour et al., 2017). The final closure of the Neo-Tethys occurred as a hard collision phase characterized by increasing crustal thickening since the early Miocene during which deformation propagated towards the craton interior to build the Zagros Mountains and their coupled foreland basin (Hessami et al., 2001; Mouthereau et al., 2007, 2012; Ruh et al., 2014; Pirouz et al., 2015; Vergés et al., 2019; Etemad-Saeed et al., 2020).

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During the Arabia-Eurasia collision, the tectonic evolution of the Zagros Fold Belt was shaped by the interplay of thin-skinned detachment folding of the Phanerozoic sedimentary cover (McQuarrie, 2004; Sherkati et al., 2006; Jahani et al., 2007; Mouthereau et al., 2007; Vergés et al., 2011a), as well as by a significant contribution of reactivated faults in the crystalline Precambrian basement, inherited from Permo-Triassic Neo-Tethys rifting (Figs. 1B and 2) (Jackson et al., 1981;
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Berberian, 1995; Talebian and Jackson, 2004; Mouthereau et al., 2006, 2007; Karasözen et al., 2019).

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The Zagros Mountain Front has a sinusoidal shape, forming salients and recesses including two well-developed arc-shaped folded domains, the Lurestan arc in NW Zagros and the Fars arc in SE Zagros (e.g., Falcon, 1961; Berberian, 1995; McQuarrie, 2004; Sepehr and Cosgrove, 2004; Emami et al., 2010; Tavani et al., 2020). Presently, there is general consensus on the forelandward sequence of folding across both Lurestan and Fars arcs in particular and in the Zagros Fold Belt in general based on the extensive database on timing of deformation (see discussion in Vergés et al., 2019). These data include Strontium isotopic ages (Ehrenberg et al., 2007; Pirouz et al., 2015, 2017) and a significant number of magnetostratigraphic studies (Homke et al., 2004; Khadivi et al., 2010; Ruh et al., 2014; Pirouz et al., 2017; Vergés et al., 2019; Lashgari et al., 2020). They have been particularly useful to determine the stratigraphy of the Mesopotamian foreland basin sedimentary fill and the timing of onset and duration of evolution of growth synclines and coeval adjacent anticlines (Fig. 1B). The folding age of the outermost domain of the Simply Folded Belt is however not well constrained. To date, only the Changuleh syncline along the front of the Lurestan arc, has yielded constraints on the age of folding at the mountain front at 8.1–7.2 Ma (Fig. 1B) (Homke et al., 2004). Because this fold developed above the blind basement thrust of the Mountain Frontal Fault, it is still unclear whether the age determined corresponds to thin-skinned folding or deep-seated basement deformation. Cover folding and basement thrusting were partially coeval as documented in Emami et al. (2010), Saura et al. (2015), and Le Garzic et al. (2019) on a large scale although the Mountain Frontal Fault in Lurestan ends up deforming the folds located in its hangingwall as discussed in Tavani et al. (2020).

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Hence, to obtain age constraints on the outermost folding in the Zagros Fold Belt, we analysed the Dowlatabad syncline in the front of the Fars arc, near the Persian Gulf, which exhibits growth patterns and is situated far away from any step formed by blind basement faults. The growth

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syncline comprises, in its SW flank, a continuous and well-preserved >3.0-km-thick foreland
sedimentary succession that begins with the shallow marine carbonates and marls of the Guri Mb.
and Mishan Fm., sampled for calcareous nannofossils and $^{87}$Sr/$^{86}$Sr isotopic dating, and continues
with the typical fluvial and alluvial continental deposits of the Aghajari and Bakhtyari formations
(2.2-km-thick log), sampled for magnetostratigraphic analysis. Furthermore, the first appearance of
Guri monomictic clasts as well as polymictic external clasts in the syncline and thus local and
regional tectonics and drainage evolution is interpreted and dated. To examine both the timing of
folding and of migration of sedimentary belts across the Zagros foreland, these ages were combined
with a regional structural cross-section extending from the Main Zagros Thrust to the SW of the
Persian Gulf, and with a compilation analysis of published magnetostratigraphic studies. In
addition, numerical experiments of basement-involved fold-and-thrust belts were conducted to test
the effect of inherited basement faults on the propagation of folding in the sedimentary cover.
Numerical results are compared to regional patterns of deformation front propagation based on
magnetostratigraphic results presented here.

109 **Figure 1**

110 STRUCTURE OF THE SE ZAGROS FOLD BELT

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The Zagros foreland Fold Belt extends along 2000 km from eastern Turkey to southern Iran,
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and is divided into different tectonic domains based on its structural and stratigraphic variations
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(Berberian, 1995; Sepehr and Cosgrove, 2004) (Fig. 1B). The arc-shaped Fars province in the SE
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Zagros is the largest tectonic domain of the Zagros Fold Belt, with more than 280 km of width (Fig.
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1B). Several studies presented geological cross-sections from the Fars province (McQuarrie, 2004;
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Sherkati et al., 2006; Alavi, 2007; Mouthereau et al., 2007; Oveisi et al., 2009; Yamato et al., 2011;
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Najafi et al., 2014), although they do not present a consensual interpretation of structural
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geometries at depth or of mechanisms of folding. In order to better define the structure of SE
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Zagros, we integrated field data with new seismic lines and well data to construct a revised structural cross-section. The section extends from the Main Zagros Thrust (MZT) to the south of the Persian Gulf in Saudi Arabia. It is oriented NE-SW, with a length of ~540 km and extending downwards to Moho depths (Fig. 2). The Zagros structural features cropping out along the cross section are described from hinterland to foreland.

The Main Zagros Thrust is the SW boundary of the Sanandaj-Sirjan metamorphic-magmatic belt and is characterized by the exposure of ophiolite and radiolarite oceanic remnants, partially over-thrusted as tectonic nappes on top of the High Zagros Imbricated Zone (Fig. 2). The High Zagros Imbricated Zone constitutes highly deformed imbricated and duplex thrust systems (Tavakoli-Shirazi et al., 2013; Sadeghi et al., 2016).

The Fars arc is characterized by remarkably regular and long NW-trending concentric folds with homogenous wavelengths of about 15 km (McQuarrie, 2004; Sherkati et al., 2006; Mouthereau et al., 2007; Fig. 2). This Simply Folded domain is limited by the Imbricated Zone to the north, and by the Persian Gulf to the south. Folds in the Fars arc involve a 10-12-km-thick pile of Phanerozoic sedimentary cover. At the base of the succession, ~2 km salt-bearing evaporites of the Ediacaran-Early Cambrian Hormuz series decouple the deformation from the Precambrian crystalline basement and detachment folding is the main deformation mechanism of the sedimentary cover in the Fars arc (Fig. 2).

In addition to the thin-skinned detachment folding, the contribution of inherited basement reverse faults in the folding of Fars arc is well-constrained for the High Zagros Fault, Surmeh Fault, and Mountain Front Fault (Figs. 1B and 2), based on their seismicity (Jackson and Fitch, 1981; Berberian, 1995; Talebian and Jackson, 2004; Karasözen et al., 2019), structural relief (Homke et al., 2010; Najafi et al., 2018) and magnetic anomalies (Teknik and Ghods, 2017). The Mountain Front Fault is also inferred from the trace of the Mountain front flexure, which is strikingly evident
in the Kirkuk, Lurestan, and Dezful domains (Fig. 1B). In the Fars arc, however, the frontal flexure
is less evident since structural steps are smaller, and therefore there is not a full agreement in the
exact location of such faults at depth (e.g., Jackson et al., 1981; Berberian, 1995; Mouthereau et al.,
2007).

The total shortening across the Fars arc is about 65 km (21%), calculated by comparing the
present length of layers (~241 km) between the pin and loose lines in the A–B cross-section, and
their pre-folding original length once restored to flat (~306 km). The pin line was located at the
deformation front near the Persian Gulf coastal line and the loose line was located in the footwall of
the High Zagros Fault (Fig. 2). The near top Bangestan Group (Upper Cretaceous carbonate unit)
was selected for unfolding since it is well-preserved at surface and well-imaged on seismic lines.
Our calculation represents a minimum value, since we do not account for internal deformation as
layer parallel shortening.

The Zagros deformation front can be traced subparallel to the Iranian coast in the Persian Gulf.
To the SW of this deformation front structures are not directly related to the frontal deformation of
the Zagros Fold Belt (Fig. 2). The Persian Gulf along the section is characterized by N-S-trending
low amplitude and long wavelength anticlines formed by dextral reactivation of inherited transverse
faults (Burberry et al., 2011; Soleimany et al., 2011), and by the Hormuz and Fars salt diapirs
(Jahani et al., 2007; Stewart, 2018; Ezati-Asl et al., 2019).

Figure 2

162 THE DOWLATABAD GROWTH SYNCLINE

The Dowlatabad syncline, is located close to the front of the Fars arc, 40 km to the northeast of
the Persian Gulf coastal line, and 250 km to the southwest of the Zagros suture zone (Figs. 1B and
2 for location). Characteristic pre-, syn- and late-growth strata patterns allow to define the relative
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timing of the deformation in this key area of the frontal portion of the Fars arc (Figs. 3 and 4).

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Furthermore, deposition of about 3.5-km-thick syntectonic sediments make the syncline an ideal

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structure to date the foreland succession. The foreland stratigraphy in the Dowlatabad syncline is

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composed of a marine to continental mixed evaporite-carbonate-siliciclastic sequence including

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Gachsaran Fm. evaporites, Mishan Fm. marls and carbonates (including the Guri Mb. limestone),

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and the upward coarsening Aghajari Fm. mudrocks, sandstones and conglomerates, ending with the

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Bakhtyari Fm. conglomerates at top (Fig. 5).

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## Structure of the Dowlatabad Syncline: Pre-growth, Growth and Post-growth strata

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The 12-km-wide Dowlatabad syncline is limited by the Sefid anticline to the north and the NW

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segment of the Pazan anticline to the south (Fig. 3). The Sefid anticline is an approximately 45-km-

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long and 11-km-wide, NW-SE trending, doubly plunging anticline, with an amplitude of 2.9 km.

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The anticline plunges 10° toward 305° at the NW nose, and 5° toward 112° at the SE nose. The stiff

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limestone of the Guri Mb. exposed in the core of the Sefid anticline forms the fold shape at surface

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that shows a simply rounded geometry with an opening angle of ~115°. The SW flank of the Sefid

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anticline is partially covered by unconformable Bakhtyari conglomeratic units (Fig. 3).

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The Pazan anticline is located to the south of the Dowlatabad syncline. The anticline is formed

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by two originally separate anticlines labelled as NW Pazan and SE Pazan that have an en-echelon

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oblique linkage together (Fig. 3A), as described in several examples in Zagros (Bretis et al., 2011;

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Fernandez and Kaus, 2014; Collignon et al., 2015). The NW Pazan is a 30-km-long and 14-km-

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wide, SW-verging fold, with an amplitude of 3.85 km. At surface, the anticline shows an

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asymmetric and rounded geometry with an opening angle of ~125°. It plunges 15° toward 308° at

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the NW nose and 7° toward 100° at the SE nose. The SE Pazan anticline shows a symmetric

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geometry and plunges 10° toward 290° at its NW nose. The formerly NW and SE Pazan anticlines
could grow with slightly different ages, as deduced by their change in direction to the south of
Dowlatabad and by the distribution of growth strata units as described below.

The integration of our field and remote-sensing mapping, together with seismic interpretations
by Motamedi et al. (2012) and Najafi et al. (2014) permits to construct a geological cross-section
across the Dowlatabad syncline (section CD; Fig. 3B). The syncline shows a closed asymmetric
chevron-type geometry at surface whereas its deeper geometry is open and displaying a box-fold
shape as imaged in the published seismic lines across the Frontal Fars (Fig. 3B).

The Dowlatabad syncline, especially at the Aghajari level, shows a smooth topography related
to recent river erosion that does not favour the identification of pre-growth, growth and post-growth
strata in a cross section. The change between pre-growth and growth units, and therefore the local
folding onset, is identified by comparing the thicknesses in the sampled section CD with the EF
section located ~10 km to the NW along the SW flank of the syncline, clearly showing a reduction
in bedding thickness towards NW (location in Figs. 3A and 4).

Along the SW flank of the Dowlatabad syncline the Aghajari Formation shows two units, a
lower unit with constant thickness, and therefore pre-growth, and an upper unit that displays a
significant thinning towards the NW termination of the NW Pazan anticline (Fig. 4). The
unconformity at the base of growth strata is labelled U1 (Figs. 3A and 4). The correlation of
bedding along the flank of the syncline has been done by remote-sensing mapping and field
measurements indicating a 1700 m thick middle and upper parts of the Aghajari Fm. growth strata
in section CD thinning to ~950 in section EF (SW of Dowlatabad village) and eventually to about
500 m towards the NW towards the fold ending (Fig. 3A and 4). The middle-upper part of the
growth strata unit is characterized by an internal ~5° angle unconformity in the centre of the
syncline (labelled U2 in Figs. 3A and 4). This unconformity does affect only few meters of
Aghajari deposits and evolves to a parallel contact with underlying Aghajari beds along the study
transect CD (Fig. 4). Therefore, unconformity U2 should not mean any problem in interpreting the magnetostratigraphy results but is important for the characterization of the growth unit internal configuration (location in Figs. 3A and 4). The Aghajari succession just above unconformity U2 contains the first appearance of carbonate conglomerates sourced from the Guri limestone providing important information about the progressive dismantling of the growing Pazan anticline.

The thick Aghajari growth unit is in turn eroded by the base of the Bakhtyari Formation, unconformity U3, and therefore showing the characteristic depositional wedge pattern of growth strata (Fig. 3A). In the NE flank of the syncline, Aghajari beds dipping between 35° and 60° crop out beneath the unconformity at Bakhtyari base. Above U3 Bakhtyari conglomerate beds gently dip (<15°) and downlap towards the NE (Figs. 3 and 4). This indicates that the conglomeratic strata did not accommodate a considerable amount of deformation and were therefore late-growth in this locality as reported elsewhere in the NW Zagros (Homke et al., 2004), northern Dezful embayment (Najafi et al., 2018), and in the Fars (Mouthereau et al., 2007).

Figure 3

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## Stratigraphy of the Dowlatabad Pre-growth and Growth Successions

The Miocene to Pleistocene siliciclastic foreland succession deposited coevally with the progressive growth of the Zagros Fold Belt constitutes the upper-most part of the Zagros sedimentary cover (James and Wynd, 1965; Kashfi, 1980; Koop and Stoneley, 1982; Khadivi et al., 2010; Ruh et al., 2014; Vergés et al., 2019; Etemad-Saeed et al., 2020). In this study, the Zagros foreland succession was examined across the Dowlatabad growth syncline in the Frontal Fars arc for the first time (Fig. 5). The Miocene to Pleistocene stratigraphy constituted by the Gachsaran, 213 214 215 216 217
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Mishan (including the Guri Mb.), Aghajari and Bakhtyari formations that is deposited on top of the late Oligocene-early Miocene Asmari carbonate platform.

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The Gachsaran Fm. is composed of about 500 m mudrocks, halite, gypsum, anhydrite, and thin-bedded carbonates in the SW flank of the Dowlatabad syncline. However, it grades to the detrital Razak Fm. towards the Interior Fars (James and Wynd, 1965; Motiei, 1993). In the Dowlatabad syncline, the Mishan Fm. overlays the Gachsaran evaporites and consists of 1060 m green and grey mudrocks intercalated with well-bedded yellowish and grey fossiliferous carbonates. The Mishan Fm. is composed of ~300 m of very thick-bedded to massive reefal limestones of the Guri Mb. at its base (Figs. 5A and B). The Mishan Fm. is deposited in a shallow open marine environment in the Frontal Fars arc (Rahmani et al., 2010; Pirouz et al., 2015).

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Above the marine Mishan Fm., the Aghajari Fm. was deposited into a siliciclastic shoreline and meandering fluvial system across the Zagros foreland basin (James and Wynd, 1965; Najafi et al., 2018; Pirouz, 2018). The Aghajari Fm. shows a ~2000-m-thick coarsening upward detrital succession and consists of light-grey to reddish medium-thick sheet-like sandstone beds and red mudrocks. The sandstone beds present abundant sedimentary paleocurrent indicators including flute casts, cross-stratifications and ripple marks (Fig. 5C). The first appearance of conglomerate and pebbly sandstone layers in the Aghajari Fm. is observed around point 2900 m on the stratigraphic log, above the U2 unconformity (Fig. 5A). The conglomerates are monomictic and composed of medium-well sorted and poorly-medium rounded carbonate pebbles in the Dowlatabad syncline (Fig. 5D).

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The Bakhtyari Fm. is deposited into a braided and high-energy fluvial system (James and Wynd, 1965; Najafi et al., 2018; Pirouz, 2018). The Bakhtyari Formation represents the uppermost late- to post-growth unit in the Dowlatabad syncline and is composed of more than 300 m of

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258 coarsening-upward, clast-supported and imbricated polymictic conglomerates, which are
259 intercalated with mudstones and cross-stratified red sandstone channels (Fig. 5E).

260 **Figure 5**

261 **Provenance of the Aghajari-Bakhtyari Deposits Filling the Dowlatabad Syncline**

262 Paleocurrent indicators, including cross-stratifications, flute casts and clast imbrications, were
263 measured in 13 sites across the Dowlatabad syncline (Fig. 5). Measured paleocurrents within the
264 Aghajari show great dispersion but large numbers are directed towards the ESE and SE with few
265 NNW and NNE directions. A major change is seen in the few measures collected in the Bakhtyari
266 conglomerates displaying S and SSE directions. The ESE to SE directions seem to indicate well-
267 developed Aghajari longitudinal fluvial systems flowing subparallel to foreland and to growing
268 anticlines. This fluvial system is pushed towards the SW by the arrival of S to SSE Bakhtyari
269 transverse alluvial system as discussed in Vergés (2007) and Pirouz et al. (2011), among others.
270 The conglomerates of the Aghajari Fm. are dominated by sub-angular limestone clasts, mainly
271 sourced from the Guri Mb., pointing towards a local source area from adjacent and nearby growing
272 anticlines (Fig. 5D). However, on top, the conglomerates of the Bakhtyari Fm. document the input
273 of well-rounded metamorphic, igneous and radiolarian chert clasts derived from distal sources in
274 the Zagros orogenic core (Fig. 5E). These observations point to an up section change in sediment
275 provenance, from proximal carbonate sources supplied by outstanding anticlines in the Frontal Fars
276 during the deposition of the Aghajari Fm. to distal source areas from the Zagros hinterland during
277 the deposition of the Bakhtyari Fm. Provenance results from the Azanak growth syncline in
278 northern Dezful embayment show substantial differences in both compositions and ages with
279 respect to the study area (Etemad-Saeed et al., 2020), indicating the complexity of the drainage
280 system in the Mesopotamian foreland basin.
281 DATING OF THE DOWLATABAD SECTION

In order to define the absolute ages of foreland succession in the frontal Fars, a new magnetostratigraphic section was sampled across the Dowlatabad growth syncline. The magnetostratigraphic ages constrained from 208 drilled sites were combined with ages provided by 20 samples of calcareous nannoplankton and 7 ⁸⁷Sr/⁸⁶Sr isotopes ages (Figs. 3A, 4 and 5A for locations of samples). Integration of all these results on the Gachsaran, Mishan, Aghajari and Bakhtyari formations provided a precise chronostratigraphic framework. The analysed succession along CD section is characterized by maximum thickness recording continuous deposition and is not apparently affected by any of the unconformities described in the syncline.

290 Magnetostratigraphy

The Dowlatabad (DW) magnetostratigraphic section is ~2-km-thick, its base is located at 28.076883° N, 52.327367° E and its top at 28.118167° N, 52.338550° E (Figs. 3A, 4 and 5A). The DW section begins at the top of the marls of the Mishan Fm., includes the entire shallow marine and fluvial Aghajari Fm. and ends up within the coarse fluvial deposits of the Bakhtyari Fm. 208 sites were drilled along the ~2200 m stratigraphic section with an average sampling resolution of 10 m/site. Two samples per site were obtained with a water-cooled electrical power drill with diamond core bits cooled by water. The samples were oriented in situ by a magnetic compass fixed to an orienting device with a clinometer. Sampling focused on fine-grained materials such as, mudstones, siltstones and fine-grained sandstones, as they are more efficient in recording the Earth's magnetic field than coarse-grained materials.

Paleomagnetic analyses were carried out in the Paleomagnetic Laboratory of Barcelona (CCiTUB-Geo3Bcn, CSIC). They consist of stepwise thermal demagnetization and subsequent measurement of the natural remanent magnetization (NRM) of 1 sample per site. Thermal
demagnetizers MMTD80 (Magnetic Measurements Ltd.) and TSD-1 (Schönstedt) and a
superconducting rock magnetometer SRM755R (2G Enterprises) were used for this purpose.
Additionally, undesired generation of new magnetic minerals during heating was controlled by
measuring the bulk magnetic susceptibility after each demagnetization step with a KLY-2
susceptibility bridge (Geofyzika Brno). Up to 16 demagnetization steps were applied to the samples
progressively tightening the temperature increments from 100 °C to 50, 40 and 30 °C up to 670 °C.
Demagnetization data are represented in vector-end point orthogonal diagrams (Zijderveld, 1967)
(Fig. 6A). In most of the samples, a viscous component parallel to the recent magnetic field is
observed, which is usually removed below 200 °C and which was not considered for further
analyses. Above this temperature, the characteristic components usually kept stable up to 670 °C,
pointing to hematite as the main remanence carrier. Characteristic components (ChComp) for each
sample were obtained by means of linear regression (Fig. 6A). ChComp's were classified as class 1
and 2 depending on yielding errors (class 1: below 10; class 2: below 20, Table S1). Finally, class 3
was assigned to those components in which polarity could be deduced but the components were
either obtained with only two points or yielded higher errors. Class 3 directions were not
considered as reliable when building the local magnetostratigraphy, but this was not a big issue as
only 8 class 3 sites were obtained, representing less than 4% of the total directions. Mean directions
for both normal and reversed polarity sets (Fig. 6B) suggest no vertical axis rotations in Frontal
Fars, in agreement with previous magnetotectonic studies in this area (Aubourg et al., 2008). Best
grouping of the directions after tectonic correction point to a primary origin of the magnetization.
Moreover, these directions passed the reversal test with a B classification (γc: 8.0 and γ0: 4.1). The
obtained directions were then computed to calculate the virtual geomagnetic pole paleolatitude
(VGP) at the site level (Fig. 7, Table S1). Positive VGPs are interpreted as normal polarity and
negative VGPs latitudes as reverse polarities and they are represented accordingly when building
the local magnetostratigraphic section (Fig. 7, black for normal magnetozones and white for the
reversed ones). A total of 14 magnetozones were defined by at least two consecutive sites of the same polarity (Fig. 7). The obtained local magnetostratigraphy has been correlated to the Global Polarity Time Scale (Gradstein and Ogg, 2012) (Fig. 7) in order to get the absolute depositional ages. It can be observed that there is a very good match between the local magnetostratigraphy and the GPTS, which favours an independent correlation. Nevertheless, it is also well supported by the chronological constraints yielded by the Sr and calcareous nannofossil samples from the Mishan Fm. and Guri Mb (Fig. 8), as it will be discussed in the following section.

The magnetostratigraphic succession sampled along CD section, in the SW flank of the Dowlatabad syncline, is the thickest in the syncline and thus records the most continuous deposition. In addition, the section is away from the unconformities described in the syncline (Fig. 4). However, the beginning of growth above U1 is recorded by a shorter reverse magnetozone in the local magnetostratigraphy (R2) compared with the associated chron in the GPTS. This implies either a strong decrease in the sedimentation rate (from 74 to 24 cm/ky, Fig. 8) or even a short stop in sedimentation not longer than few tens of thousands of years since reversed polarities are maintained after the growth (Figs. 7). The intra-growth strata U2 does affect only few meters of Aghajari deposits and evolves to a parallel contact with underlying Aghajari beds along the study transect CD (Fig. 4). The U3 at the base of late-growth Bakhtyari conglomerate occurs around the boundary between N7 and overlying R7 of local polarity corresponding to C1 and C2 chrons in the GPTS, respectively. However, no effect of this unconformity is observed in the reversal pattern, since the relative length of the magnetozones is preserved, implying that no significant time is missed in the sampled section (Fig. 7). In general, the longer chron duration and lower sedimentation rate reduces the ability of magnetostratigraphic data in detecting short time hiatuses, as observed in the Southern Pyrenees (Garcés et al., 2020) and in the Fars arc of Zagros (Ruh et al., 2014). Despite unconformities, the good match of this magnetostratigraphic log with the GPTS is a strong indication of a correct correlation as discussed in Garcés and Beamud (2020).
356 **<sup>87</sup>Sr/<sup>86</sup>Sr Isotopic Dating**

357 Strontium Isotope Stratigraphy is a dating method based on the <sup>87</sup>Sr/<sup>86</sup>Sr ratio in the
358 composition of marine authigenic sediments that record the chemical composition of the sea water
359 at the time of mineral precipitation (Elderfield, 1986). The ratio of <sup>87</sup>Sr/<sup>86</sup>Sr in sea water has
360 changed significantly through geological times in response to the input of varying proportions of Sr
361 from continental crust and upper mantle sources. These variations constitute a global standard that
362 has been widely used for dating of carbonate sediments. Rapid monotonous increase in the <sup>87</sup>Sr/<sup>86</sup>Sr
363 ratio from 0.7082 to 0.7092 during the Neogene, makes this period very promising for Sr dating as
364 age uncertainties are less than 1 Ma (McKenzie et al., 1988; Hodell et al., 1991; McArthur et al.,
365 2001). In this study, 7 samples were taken from limestones and marly limestones of the topmost
366 part of the Gachsaran Fm., Guri Mb. and Mishan Fm. (Fig. 8). Following the quality criteria for Sr
367 dating (Elderfield, 1986), samples were processed at the “Laboratorio de Geocronología y
368 Geoquímica Isotópica of the Universidad Complutense de Madrid” in Spain. The obtained results
369 are summarized in Table S2. The numerical ages were determined by plotting the obtained <sup>87</sup>Sr/<sup>86</sup>Sr
370 values on the <sup>87</sup>Sr/<sup>86</sup>Sr LOWESS (Locally Weighted Scatterplot Smoother) calibrated curve
371 (McArthur et al., 2001) (Fig. 8 and Table S2). Five samples of the Guri Mb. at the base of Mishan
372 Fm. (SR81, SR82, SR83, SR85, SR76) cluster between 12.2 and 11 Ma (Fig. 8), suggesting an age
373 around 11.5 Ma for the deposition of this unit in the front of Fars arc. This age is in agreement with
374 reported ages between 11.5 ± 0.5 and 11.1 ± 0.1 Ma for the Mishan Fm. samples from Coastal Fars
375 next to the modern coast of the Persian Gulf (Pirouz et al., 2015). However, two samples from
376 topmost part of the Gachsaran Fm. (SR80 and SR93) show an age around 15 Ma.
377 **Calcareous Nannofossils**

378 In this study, 20 samples were collected to obtain the calcareous nannofossils content in the
379 Mishan Fm. marls and first meters of the Aghajari Fm. (Figs. 3A, 4, 5A and 8, for sample
380 locations). All the samples were prepared from unprocessed material as smear slide using standard
381 procedure (Bown, 1998). The smear slides were analysed using a Zeiss Axioplan petrographic
382 microscope at different magnification (x 750-x 1250). 15 of the 20 samples provided positive
383 results (Table S3). The studied samples were, in general, of poor quality due to the poor
384 preservation and low abundance of the autochthonous assemblages and the presence of abundant
385 reworked species (more than the 50% of the assemblage, in most of the samples). The lowermost
386 samples (NF77-NF84) contained almost no nannofossils, and the few observed ones were usually
387 poor preserved, most of them being reworked dissolution resistant species. The calcareous
388 nannofossil abundance and preservation improves up-section. Samples NF87, NF88 and NF92
389 were rich in moderately to well-preserved nannofossils, but most of them were reworked species.
390 Finally, from sample NF94 onwards both abundance and preservation got worse again. Table S3
391 summarizes the obtained results for Neogene species. The standard schemes of Martini (1971) and
392 Okada and Bukry (1980) have been adopted for this study. Cretaceous and Paleogene reworked
393 species are not included in order to simplify the presentation of results. It is worth noticing the
394 reduced amount in both number of species and number of specimens of the *Discoaster* genera in
395 the studied assemblages, pointing to a deposition in a shallow restricted sea.

396 Regarding biostratigraphy, we have obtained disparate results along section. Due to scarcity and
397 low preservation of calcareous nannofossils in the lower part of the section, it was difficult to
398 assign them to a biozone. In sample NF77 autochthonous calcareous nannofossil assemblages are
399 absent whereas samples NF81-NF84 showed only some Miocene species with a wide stratigraphic
400 range such as *Calcidiscus leptoporus*, *Reticulophenestra haqii*, *Unbilicospahaera jafari* and
Umbilicosphaera rotula. In Table S3 the lowermost zone that can contain the observed assemblage is indicated. Samples NF81–NF84 are, thus, related to zone NN1-NN2, although they could belong to a higher zone. Samples NF85 and NF86 belonged to Zone NN7 as they recorded the presence of *Calcidiscus macintyrei* and *Reticulofenestra pseudoumbilicus* (Table S3). The preservation of the assemblages, however, is still poor and thus they could pertain to a younger zone. Samples from the middle part of the section were the best quality ones. The presence of *Discoaster quinqueramus*, *Discoaster berggrenii* and *Discoaster bergenii* in samples from NF87 to NF92 indicated that this stratigraphic interval belonged to the lower part of Zone NN11a (Zone CN9a). There were no evidences suggesting a younger age, as species from genera *Amaurolithus* (CN9b) or *Ceratolithus* (CN10b, NN12) were not identified. However, its absence may obey to the general poor preservation state of the nannofossils assemblages. Actually, species from genera *Catinaster* (NN8-NN10) or *Triquetrorhabdulus* (Paleogene-NN12) were neither identified. The uppermost samples (NF94–NF97) were, again, of bad quality hindering any Zone assignation.

## 414 Integrated Dating Results

In the Dowlatabad section, the oldest sampled part of the Gachsaran Fm. shows Sr isotopic ages of ~15.5 Ma. In the Guri Mb. at the base of the Mishan Fm., despite some uncertainties, most of the samples yield Sr isotopic ages around 11.5 Ma, between 12.5 and 11.1 Ma.

On top of the Mishan Fm., samples NF87, NF89, NF90 and NF92 show the only calcareous nannofossil assemblages that provide an accurate zone determination. This assemblage is assigned to zone NN11a, with an age ranging from ~8 to 5.5 Ma (Fig. 8). The moderate to poorly preserved calcareous nannofossil assemblages identified in samples NF85 and NF86, in the middle part of Mishan Fm. have been assigned to zone NN7, ranging from 11.8 to 10.8 Ma (Fig. 8). This would mean that between sample NF86 and NF87, less than 100 m apart, there would be a time gap of about 2.8-3.8 Myr. In the Dowlatabad section there is no evidences of a sudden change in the
425

sedimentation rate or the presence of a hiatus between both samples that could explain this long-
time gap in less than 100 m of sedimentary succession. Therefore, samples NF85 and NF86 are
more likely to belong to younger zones, NN9 and NN10 respectively, and the absence of NN9 and
NN10 biomarkers is therefore most likely due to the poor preservation of calcareous nannofossil
assemblages of these samples.

426

427

428

429

430

The contact between Mishan and Aghajari formations is located close to the base of the magnetostratigraphic section and has an age slightly younger than 4.9 Ma within the C3 chron in the GPTS. Further up, the contact between Aghajari and Bakhtyari formations coincides with the boundary between C1 and C2 chronos, with an age of 1.8 Ma (Fig. 8).

431

432

433

434

The timing of the onset of deformation in the Dowlatabad syncline is marked by the beginning of
growth strata after the deposition of ~230 m of Aghajari Fm. at paleomagnetic site DW226. The
correlation with the GPTS shows an age of 4.65 Ma for the contact between pre-growth and growth
strata (Fig. 8). The growth succession has a maximum thickness of 2000 m in the syncline and was
deposited during 2.85 Myr. The base of the gently folded late-growth Bakhtyari Fm., from
paleomagnetic site DW131 upward to DW150, has an age of 1.8 to 1.5 Ma (Fig. 8). Coevally, the
reorganization in the drainage network took place from longitudinal to transverse arrangements,
evidenced by changes both in clast composition with the input of external clasts, and paleocurrent
directions in the Bakhtyari conglomerates (Fig. 8). Therefore, the amplification of the NW Pazan
anticline in the Frontal Fars lasted for 2.85 Myr, from 4.65 Ma to 1.8 Ma, which yields a 1.4 mm/yr
rate of fold amplification.

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436

437

438

439

440

441

442

443

444

445

The sedimentation rates for the Aghajari Fm. were deduced from the proposed correlation of the
local magnetostratigraphic section to the GPTS (Fig. 8). The base of the section is characterized by
a change in sedimentation rates from 46 to 74 cm/kyr. The base of growth strata is marked by a
significant short period during which there was either a decrease in sedimentation rates to only 26

446

447

448
449 cm/kyr or a short break in sedimentation above the U1 unconformity. This was followed by a
450 subsequent increase to 79 cm/kyr. Sedimentation rate keeps high, ranging between 62 and 75
451 cm/kyr, during most of the growth development. The end of growth occurred with sedimentation
452 rates around 62 cm/kyr and they slightly increased to 68 cm/kyr close to the contact with the
453 Bakhtyari Fm. The estimation of the sedimentation rate for the Bakhtyari Fm. is problematic since
454 there are few tie points within the formation. About 200 m of the Bakhtyari sediments were
455 deposited from 1.8 to 1.5 Ma in the Dowlatabad syncline, which yields a sediment accumulation
456 rate of 68 cm/kyr. These comparatively high sedimentary accumulation rates resemble those
457 determined in the northern Dezful embayment where there is a significant accumulation of
458 syntectonic deposits (Najafi et al., 2018; Lashgari et al., 2020).

459 Although much less constrained, from top Guri Mb. at ~11.5 Ma to base of the Aghajari Fm. at
460 ~4.9 Ma, the marls of the Mishan Fm. would have been deposited at a rate of 12 cm/kyr. Using
461 these 12 cm/kyr deposition rates of the Mishan Fm., the age of the top of Gachsaran Fm. can be
462 assumed as ~14 Ma. This age would be in agreement with the results of the SR80 and SR93
463 isotopic samples (Fig. 8).

464 **Figure 8**

465 NUMERICAL MODELLING

466 Integrated dating of the Miocene sequences and related progressive and angular
467 unconformities of the Dowlatabad syncline give insight into the timing of local folding and the
468 switch from marine to non-marine sedimentary systems. A comparison of the obtained data to
469 previous studies of folding ages across the Zagros Fold Belt (Homke et al., 2004; Khadivi et al.,
470 2010; Ruh et al., 2014; Pirouz et al., 2017; Vergés et al., 2019; Lashgari et al., 2020) provides a
471 better understanding of the temporal evolution of deformation propagation at a large-scale.
However, the effects of basement faulting on the cover folding propagation is difficult to discriminate from dating growth folding in the cover sequence alone. Therefore, a couple of numerical experiments were designed to better understand the large-scale structural evolution of a basement-involved fold-and-thrust belt and the deformation front propagation of its cover sequence in particular.

## Methods and Model Setup

The numerical model is based on a thermo-mechanical fully-staggered finite difference code with a visco-elasto-plastic/brittle rheology. For a more detailed description of the governing equations and their numerical implementation, we refer to Ruh (2017) and Ruh et al. (2018). The applied Eulerian box size measures 600 x 60 km in width and height, respectively, with a resolution of 1201 x 121 nodal points and 16 Lagrangian markers per node cell (Fig. 9). From bottom to top it consists of 25 km of diabase basement, 2 km of low-viscous detachment (equivalent to the Ediacaran-Early Cambrian Hormuz salt) between 100 km < x < 550 km, and 10 km of upper sedimentary crust with two weak interlayers (equivalent to the pre-folding Phanerozoic sedimentary sequence of the Zagros Fold Belt). The top 23 km exhibit a so-called sticky-air layer of low density and viscosity to ensure zero stresses along the evolving rock/air interface, mimicking a free surface (Crameri et al., 2012). Rheological parameters and density of all marker types are given in Table S4. The initial temperature distribution describes a linear geotherm within the rock sequence with 0 °C at the rock/air interface and 600 °C at the base of the model. Velocity boundary conditions resemble those from sandbox experiments. At the base and the left side, a velocity of 1 cm/yr is applied, whereas the right side represents a rigid backstop with zero velocities in either direction (Fig. 9). The rock/air interface undergoes linear diffusion with a diffusivity coefficient of 10⁻⁶ m²/s to mimic erosion and sedimentation during folding (Ruh et al., 2018). All models in this study ran for 15 Myr to accommodate shortening equal to 25% of their original length.
In the Fars arc of the Zagros, basement reverse faulting is well-documented for the High Zagros Fault, Surmeh Fault, and Mountain Front Fault (Figs. 1B and 2) dominantly based on seismicity (Jackson, 1980; Berberian, 1995; Talebian and Jackson, 2004; Karasözen et al., 2019). These studies showed that fault planes of basement faults are relatively steep (40–50°), pointing towards reactivated inherited normal faults. To investigate deformation propagation within the upper crustal sequence and to furthermore test the effects of inherited basement faults (see Table S4 for rheological parameters) on deformation propagation and overall structural evolution, two experiments are presented here (Fig. 9): i) Model 1 lacks of predefined inherited basement faults and ii) Model 2 exhibits two inherited listric faults (initially at x = 400 km and x = 550 km) with an angle of 45° at the top of the basement shallowing downwards, geometrically comparable to the interpreted basement thrusts in the Fars (Fig. 2). Although the listric basement thrusts in the Zagros are relatively shallow (30–45°; Fig. 2), they were interpreted as reactivated listric normal faults (Jackson, 1980).

Figure 9

Modelling Results

Figure 10 illustrates the temporal evolution of numerical experiments without (Fig. 10A–C) and with pre-existing listric faults within the basement (Fig. 10D–F). In the case of no inherited basement faults, marker composition (upper panels) indicates significant basement thrusting close to the backstop wall (<120 km) after 5 Myr of shortening (Fig. 10A). Strain-rate plots infer diffuse rather than localized brittle failure of the basement up to 500 km away from the backstop wall. At this stage, the cover sequence is folded at x ≈ 150–250 km (~100 km in front of the most active basement thrust). Developing surface basins (synclines) between anticlines are filled by synorogenic sediments (Fig. 10; yellow colour). After 10 Myr of shortening, a sequence of faults, each with ~1–2 km offset, developed within the basement up to ~270 km away from the backstop wall (Fig. 10B). The uppermost cover sequence shows typical fault-propagation folding with thrusts in
the anticlinal hinges. Folding is more intense above thrusted basement (x = 100–300 km) and at the salt pinch-out at x ≈ 460 km. Although basement faults may exhibit offsets of several kilometres, they do not crosscut the salt detachment, which allows the cover sequence to migrate forward along it. Strain rates indicate that after 15 Myr of shortening, active deformation takes place in a thin-skinned manner in front of the salt pinch-out and its accompanying basement thrust (Fig. 10C).

Inherited lower-strength basement faults strongly affect the structural and mechanical evolution of numerical fold-and-thrust belts presented here. After 5 Myr of shortening, a significant amount of shortening is accommodated by the basement fault closer to the backstop at x ≈ 350 km (Fig. 10D). A large-wavelength anticline forms above the reactivated basement fault. The second invariant of the strain-rate tensor furthermore indicates that also the frontal basement fault beneath the salt pinch-out slips early during the shortening process. The downward dispersion of strain rate localization of reactivated basement faults at y ≈ 45 km indicates the brittle ductile transition. After 10 Myr, both basement faults exhibit offsets of ~3–5 km and the upper crustal sequence is shortened above the salt detachment by multi-layer fault-propagation folding (Fig. 10E). After 15 Myr of shortening, major out-of-sequence basement faulting occurs close to the backstop (x ≈ 50–150 km), affecting the style of deformation in the upper crustal sequence (Fig. 10F). Activity along the frontal basement fault leads to thin-skinned shortening of sequences above intermediate weak layers pushed by the hangingwall (x ≈ 450–500 km).

Summarizing, if there are no pre-existing weak faults within the basement, the upper crustal sequence shows a roughly constant wavelength of folding/thrusting, forming a gently dipping slope (Fig. 10C). A major forward-verging thrust fault roots into where the salt pinches out at the wedge toe. The basement is faulted regularly beneath the weak salt layer by mainly forward-verging thrusts. Reactivation of two inherited listric faults in the basement leads to more intense strain localization along them during shortening (Fig. 10F). Areas around the inherited faults experience less deformation than in the experiment without inherited faults. Reactivation of those basement
547 faults strongly affects deformation of the upper crustal sequence, where long-wavelength folds
548 develop above the uplifted basement structures. Basement deformation is also restricted along the
549 backstop and the inherited faults in case of two pre-existing faults (Fig. 10F). The overlying
550 sedimentary sequence is folded with long-wavelength anticlines above the reactivated faults.

551 **Figure 10**

552 DISCUSSION

553 This section is organized around three major topics: a) Early Miocene–Pleistocene propagation
554 of folding and migration of sedimentary belts across the Fars arc; b) Deformation propagation and
555 effects of basement-involvement inferred from numerical modelling; and c) Timing of folding in
556 the Zagros and its implications for the Arabia-Eurasia collision.

557 Early Miocene–Pleistocene Propagation of Folding and Migration of Sedimentary Belts

558 Across the Fars Arc

559 The results of our study provide high-resolution age constraints on a well-preserved and
560 continuous foreland succession as well as the timing of onset and cessation of folding in the Frontal
561 Fars arc (Fig. 8). These results are then integrated with published magnetostratigraphic sections
562 across the Fars arc in SE Zagros Mountains to establish a new Miocene–Pleistocene
563 chronostratigraphic panel along a NE-SW trend, from the footwall of the Main Zagros Thrust
564 (MZT) to the coast of the Persian Gulf (Fig. 11A). The most striking fact from the revised
565 Miocene–Pleistocene chronostratigraphy for the Fars region is that the sedimentary units grouped
566 in formations are diachronous. The transition between marine marls of the Mishan Fm. and the
567 Aghajari continental sediments is dated to 18.8 Ma in the Dehmoord syncline ahead of the Main
568 Zagros Thrust (Pirouz et al., 2017), whereas its age is 16.6 Ma in the Qalat syncline (Khadivi et al.,
569 2010) and 6 Ma in the Ghol Ghol syncline in Interior Fars (Ruh et al., 2014). Finally, it gets as
570 young as 4.9 Ma in the Dowlatabad growth syncline in the frontal part of the onshore SE Zagros
571 (this study). The contact between the Aghajari Fm. and the conglomerates of the Bakhtyari Fm.
572 occurs at 16.5 Ma in the footwall of the Main Zagros Thrust (Pirouz et al., 2017), whereas it is 14.8
573 Ma in the Qalat syncline (Khadivi et al., 2010), and 3.2 Ma in the Ghol Ghol syncline in Interior
574 Fars (Ruh et al., 2014), and eventually 1.8 Ma in the Frontal Fars (this study). The lateral facies
575 changes between non-marine Aghajari and shallow marine Mishan formations was already
576 interpreted in the pioneering work of James and Wynd (1965) in the Fars.

577 We consider the age variation of both Mishan–Aghajari and Aghajari–Bakhtyari contacts across
578 the Zagros to calculate the rate of migration of the sedimentary belts (progradation) toward the
579 foreland (Fig. 11). The Aghajari and Bakhtyari sedimentary belts have been migrating with rates of
580 18.9 mm/yr and 17.9 mm/yr, respectively, across the SE Zagros Mountains. The same calculation
581 for the NW Zagros shows that the rate of the Aghajari sedimentary belt migration of 19 mm/yr is
582 faster than the rate of the Bakhtyari sedimentary belt migration (11.4 mm/yr). The progressive
583 migration (progradation) of Mishan, Aghajari and Bakhtyari formations towards the foreland is
584 explained by the tectonic propagation in the same direction as in the Lurestan (Pirouz, 2018;
585 Vergés et al., 2019).

586 Using magnetostratigraphic data, Vergés et al. (2019) calculated Aghajari and Bakhtyari
587 sedimentary belts migration rates in Lurestan of NW Zagros to be 36 mm/yr and 18 mm/yr,
588 respectively. They also show that the folding front propagation toward the foreland occurred at a
589 mean rate of 30 mm/yr. In another study in SE Zagros, Pirouz et al. (2015) proposed between 17.8
590 and 50 mm/yr for the progradation of the Mishan facies based on Sr dating of carbonates.

591 The combination of our determined age for the growth strata in the Frontal Fars with published
592 ages across SE Zagros reveals that folding front propagation was non-steady with highly variable
rates (between 8 and 34 mm/yr) during the evolution of the Zagros Fold Belt (Fig. 11).
Deformation started around 18 Ma in the Dehmoord syncline ahead of the Main Zagros Thrust
(Pirouz et al., 2017). The deformation front propagated very fast, 34 mm/yr, reaching the Qalat
syncline at 14 Ma (Khadivi et al., 2010) (Fig. 11A). It was then migrating slowly, 8 mm/yr, across
the Interior Fars and reached the Ghol Ghol syncline at 3.8 Ma (Ruh et al., 2014). Almost at the
same time (4.6 Ma), folding started in the Dowlatabad syncline in the Frontal Fars arc (this study).
Despite these changes, the average rate of folding front propagation from the suture zone to the
current deformation front has been around 20 mm/yr from Burdigalian to recent. A similar
calculation made in the NW Zagros shows between 7 and 55 mm/yr, with an average rate of 15
mm/yr for the propagation of deformation front, from Aquitanian to Present (Fig. 11B). Although
our compilation suggests mean rates of folding front propagation in both SE and NW Zagros (20
and 15 mm/yr, respectively), it also shows that the folding front propagation rates have been
significantly variable throughout the lifetime of the Zagros Fold Belt.

The migration of sedimentary belts in a foreland basin is related among other factors to the ratio
between the available basin space, because lithospheric bending, and the amount of basin filling
sediments: a high ratio produces sedimentary aggradation whereas a low ratio generates
progradation of sedimentary sequences. The progradation of sedimentary belts and the propagation
of deformation should not follow a similar path as the latter is highly conditioned by factors such as
the existence of detachment levels, amount of syntectonic sedimentation, load distribution in the
hinterland, and others (e.g., Graveleau et al., 2012). In the Fars arc, however, sedimentary
progradation and folding propagation towards the SW occurred with great synchrony as the
Aghajari and Bakhtyari calculated migration rates are 18.9 and 17.9 mm/yr and the propagation of
folding is 20 mm/yr (Fig. 11).

The rates of syntectonic sediment accumulation vary across the Zagros foreland (Fig. 11A). The
sedimentation rate of the Aghajari growth deposits varies between 22 and 28 cm/kyr from the Main
618
619

Zagros Thrust towards the Interior Fars (Khadivi et al., 2010; Ruh et al., 2014; Pirouz et al., 2017), reaching the highest rate of 63 cm/kyr in the frontal portion of the Fars arc (Fig. 11A).

620
621
622
623
624
625
626
627
628
629
630
631
632
633

As described in section 2, shortening in the Fars arc accounts for about 65 km and 21% during the last ~16 Myr, giving rates of ~4 mm/yr. This almost two-times shortening amount is, however, comparable to the 20% shortening obtained for the Lurestan arc during 22 Myr (Casciello et al., 2009; Vergés et al., 2011b). The higher shortening rate in Fars arc could be achieved by a greater number of thrusts reactivating inherited faults. In both areas, shortening rates are significantly lower than folding front propagation rates of ~15 mm/yr in the Lurestan arc and ~20 mm/yr in the Fars arc. There are two likely causes for this fast propagation of Zagros folding towards the foreland. The reactivation of pre-existing faults in the Arabian basement due to the transmission of far field stresses as was suggested as a main reason (Mouthereau et al., 2007; Farahpour and Hessami, 2012; Pirouz et al., 2015). An alternative or combined explanation would be the presence of Ediacaran–Early Cambrian Hormuz salt-bearing evaporites at the basement-cover boundary (Fig. 2). Numerical and analogue experiments support that the strength of the basal detachment affects the relation between propagation and shortening rates (Cotton and Koyi, 2000; Ruh et al., 2012, 2013).

634

**Figure 11**

635
636

**Effects of Inherited Basement Faults and Basal Detachment Level on Deformation**
**Propagation Inferred from Numerical Modeling**

637
638
639
640

In this study, numerical models are used to explore the propagation of folding across fold belts in an attempt to draw conclusions about the potential roles of a basal detachment level and inherited basement faults, applicable to the vast amount of data from the Zagros Fold Belt compiled in this work. Figure 12 illustrates rock composition after 15 Myr of shortening (upper panels) and
641 averaged horizontal shortening rate of the cover sequence (lower panels) for the entire time span of
642 the respective experiments (0–15 Myr). In the following, we discuss similarities and discrepancies
643 of the two presented experiments with and without deep-seated inherited basement faults and
644 discuss numerical results with respect to observations from the Zagros.

645 Both experiments show a very low surface taper angle (~1°), in agreement with calculations in the
646 Lurestan and Fars arcs that are typical of fold belts detached above an efficient evaporitic horizon
647 at depth as proposed by McQuarrie (2004), Ford, (2004), and discussed in Vergés et al. (2011a).
648 Furthermore, both experiments show in-sequence and fast forelandward propagation of folding
649 (lateral migration of shortening) within the cover sequence above the efficient basal detachment
650 level equivalent to the Hormuz salt (Fig. 12). This fast propagation forms a wave of folding
651 deforming strips of previously undeformed foreland that can account for the almost synchronous
652 growth onset of the Ghol Ghol (Ruh et al., 2014) and Dowlatabad (this study) synclines in the Fars
653 arc at 3.8 and 4.6 Ma, respectively (Fig. 11A).

654 Besides the aforementioned common points, there are particular discrepancies between the
655 experiment with and without inherited basement faults. The temporospatial evolution of shortening
656 rates in the experiment without inherited basement faults shows a generally constant rate of folding
657 propagation that slows down during the last ~5 Myr (Fig. 12A). A feature of this model is that the
658 basement-cover boundary, i.e. the weak basal detachment, maintains a smooth topography despite
659 having a basement affected by the same amount of shortening as the cover sequence. Another
660 outcome is that folds have a limited duration of growth restricted to a few million years. This is in
661 agreement with field data where the entire folding process was determined by dating the late- or
662 post-growth strata units, like in the Changuleh, Jarik and Dowlatabad synclines showing active
663 growth thru 5, 5.8 and 2.85 Myr, respectively (Homke et al., 2004; Lashgari et al., 2020; this
664 study).
In contrast, the experiment with inherited basement faults results in an interrupted pattern of shortening propagation over time (Fig. 12B). Despite the general trend of forward propagation of deformation, zones of very low horizontal shortening (<1 mm/yr) occur above the hangingwall of the two inherited basement faults (Fig. 12B) that underwent relatively early inversion (Figs. 10D–F). These large basement-involved monoclines act as buttressing zones separating cover folding domains where folds remain active throughout the duration of the experiment. The early activity of basement thrusts at depth can explain the relatively early (~7.6 Ma) growth of the Changuleh anticline actually located above the flexure of the Mountain Frontal Fault in the Lurestan arc (Homke et al., 2004; Emami et al., 2010). Such early growth with respect to the general deformation propagation is also inferred in the southern Fars region by Mouthereau et al. (2006). The longer duration of these basement blind thrusts, specifically their late activity, can be interpreted as partially out-of-sequence since cover folds above the thrust hangingwalls are passively transported. These sequential relationships are observed in the Lurestan front where the forelimbs of the large Anaran and Siah Kuh anticlines gravitationally collapsed along the hangingwall of the Mountain Front Fault (e.g. Emami et al., 2010; Casini et al., 2018). Therefore, although the regional forelandward sequence of deformation is clearly represented in the model, the pre-existence of basement faults results in a more intricate deformation sequence including local out-of-sequence folding.

As a summary, the general sequence in both experiments occurs in-sequence, advancing toward the foreland and thus comparable with magnetostratigraphic ages of folding onset younging in the same direction (Figs. 11 and 12). In the experiment without inherited faults, the propagation of the onset of folding is continuous and anticlines grow for a limited period of time of few million years. In the second experiment, propagation of folding onset shows a more discontinuous pattern where inherited faults concentrate shortening in the basement and thus a more complex scenario, in which local out-of-sequence folding may occur as a result of basement inheritance. However, the present
data set on ages of onset and duration of folding for the Zagros Simply Folded Belt, although
abundant and strategically distributed, is probably not sufficient to determine potential local out-of-
sequence folding activity that should further be analysed in more detail with multidisciplinary
studies. It is worth noting that the in-sequence folding propagation established in this study does
not indicate a significant disruption in the propagation wave of folding at a large-scale. This
undisrupted pattern indicates that the inverted basement thrusts were not significantly affecting the
forward migration of folding. This could be related to a more complex thrusting geometry at depth
favored by the distribution of the effective basement-cover interface detachment level that could
transfer shortening from deeper thrusting in the hinterland to Hormuz salt or equivalent far away in
the foreland. In this scenario, the basement thrusting observed in the Zagros Simply Fold Belt
would uplift already folded cover sequences stepwise as illustrated in Lurestan (e.g., Saura et al.,
2015) and Dezful (e.g., Etemad-Saeed et al., 2020).

702 **Figure 12**

703 Timing of Folding Across the Zagros and its Implications for the Arabia-Eurasia Collision

The combination of our results in the Dowlatabad growth syncline with eight published
magnetostratigraphic sections across the Zagros Fold Belt, indicates that the onset of cover folding
is everywhere younger than 22 Ma in both Lurestan and Fars arcs (Fig. 13). The oldest growth ages
were determined to be Aquitanian–Burdigalian in the Shalamzar and Dehmoord synclines, ahead of
the Main Zagros Thrust (Pirouz et al., 2017) (Fig. 13). However, the onset of Arabia-Eurasia
suturing is slightly older and is marked by the first arrival of clastics from the Sanandaj-Sirjan to
the Zagros foreland basin as discussed in Mouthereau et al. (2012) and Gholami Zadeh et al.
(2017). These clastics are found in red beds of the Razak Fm. in Iran and Suwaise Unit in Iraqi
Kurdistan, with a late Oligocene age (Chattian, ~26–25 Ma) based on recently published detrital
zircon and magnetostratigraphic data (Pirouz et al., 2017; Barber et al., 2018; Koshnaw et al.,
714 2018). The slow exhumation of the Sanandaj-Sirjan zone along the hangingwall of the Main
715 Zagros Thrust started at ~24 Ma (Chattian), and accelerated between 20 and 14 Ma (Burdigalian-
716 Langhian), as determined by thermochronological data in Kermanshah and Lurestan in NW Zagros
717 (Barber et al., 2018). The first foreland sediments with Zagros provenance arrived to the Dezful
718 embayment during late Oligocene–early Miocene times (~25–18 Ma) (Etemad-Saeed et al., 2020).
719 However, these risings of tectono-topographic reliefs in the Zagros can be linked to either the onset
720 of continental collision (soft-collision) or to a more advanced stage of orogenic evolution (hard-
721 collision) (Ballato et al., 2011; Mouthereau et al., 2012).

722 The Arabia-Eurasia collision, therefore, incepted in Chattian (late Oligocene), and the Zagros
723 folding initiated in Aquitanian–Burdigalian (early Miocene). The folding then propagated
724 forelandward through the Interior Fars, Lurestan, Izeh Zone, and north Dezful embayment during
725 the Langhian and Serravallian (middle Miocene), as defined in Afrineh, Chaman-Goli, Jarik and
726 Qalat growth synclines (Khadivi et al., 2010; Vergés et al., 2019; Lashgari et al., 2020). According
727 to these ages, the propagation of folding was occurring similarly along the strike of the Zagros fold
728 belt from Lurestan to Fars arcs during middle and late Miocene times. The good correlation along
729 the strike of the Zagros Fold Belt of onset of folding ages indicates that the present geometry of
730 salients and reentrants bounding the Lurestan arc, the Dezful embayment and the Fars arc does not
731 seem to influence the propagation of the folding wave (Fig. 13). The width of the folded domain
732 was about 160 km at the end of Miocene whereas the maximum present-day width is 230 km in the
733 Lurestan arc and 280 km in the Fars arc. This difference in the width of the Simply Folded Belt,
734 apart from other factors such as the spatial variation of Hormuz salt at depth (Bahroudi and Koyi,
735 2003), could be influenced by the inherited segmentation of the Arabian margin as proposed in
736 Navabpour et al. (2014). The partly coeval growth of the Ghol Ghol and Dowlatabad synclines in
737 the Frontal Fars arc indicates that a number of close anticlines grew in a period of few million years
738 from early Pliocene (Zanclean) to early Pleistocene (Gelasian) (Fig. 13). The numerical
739

simulations of the Fars folding by Yamato et al. (2011) and Fernandez and Kaus (2014) also show
the coeval growth of several anticlines within a limited range on top of Hormuz salt-bearing basal
detachment level as highlighted by our numerical modelling results.

740

741

742

This Neogene folding postdates two earlier pre-collision tectonic events in SE Zagros, i.e., a) the Hormuz salt diapirism from Palaeozoic to Recent (Jahani et al., 2007; Perotti et al., 2016; Stewart, 2018) and, b) the Late Cretaceous reactivation of N-S trending inherited basement faults from the Late Cretaceous (Burberry et al., 2011; Saura et al., 2011; Soleimany et al., 2011; Farahzadi et al., 2019). Both events developed broad and low amplitude antiforms in the sedimentary cover, which are still preserved in the Persian Gulf to the SW of the Zagros deformation front (Fig. 13). To the NE of this front, both diapiric domes and N-S trending folds are overprinted by the NW-SE trending Zagros Fold Belt structures as those presented in this study.

743

744

745

746

747

748

749

750

Figure 13

751

CONCLUSIONS

752

The ages of the syntectonic foreland deposits of the Frontal Fars domain have been determined in the Dowlatabad growth syncline by the combination of calcareous nannofossil analysis, <sup>87</sup>Sr/<sup>86</sup>Sr isotopic dating and magnetostratigraphy for the marine to non-marine mixed evaporite-carbonate-siliciclastic successions of Mishan, Aghajari and Bakhtyari formations. These ages have been integrated with provenance and paleocurrent analysis, as well as within the framework of a detailed structural study of the growth syncline and limiting anticlines. Following we summarize the main results of this study.

753

754

755

756

757

758

759

• The deposition of marine marls of the Mishan Fm. started around 11.5 Ma in the Dowlatabad growth syncline. The transition to a non-marine basin condition occurred at 4.9 Ma by the progradation of thick fluvial deposits of the Aghajari Fm. with a fast accumulation rate of 63

760

761
762 cm/kyr. Monomictic conglomerate beds, above the intra-Aghajari U2 unconformity, composed of
763 Guri Mb. carbonate clasts were sourced from nearby growing anticlines and deposited between 2.8
764 and 1.95 Ma. At the uppermost part of the foreland succession, the Bakhtyari polymictic
765 conglomerates derived from Zagros hinterland, unconformably cover the Frontal Fars synclines
766 from 1.8 to <1.5 Ma, with an accumulation rate of 68 cm/kyr.

767 • Folding of the Frontal Fars arc started at 4.65 Ma, as marked by the beginning of growth
768 strata above unconformity U1. Several anticlines grew almost coevally during the early Pliocene.
769 The major pulse of folding in the Dowlatabad growth syncline lasted for 2.85 Myr, ceasing at
770 around 1.8 Ma, in early late Pleistocene. The NW Pazan anticline amplified with a rate of 1.4
771 mm/yr.

772 • Age variations of both Mishan–Aghajari and Aghajari–Bakhtyari contacts from the Zagros
773 suture zone towards the deformation front mark the migration rate of foreland sedimentary belts
774 (progradation) that is consistent with the propagation of folding obtained from dating of growth
775 strata across the Lurestan and Fars arcs. The propagation of folding and the progradation of the
776 depositional systems occurred by a mean rate of ~20 mm/yr in SE Zagros (Fars arc), and ~15
777 mm/yr in NW Zagros (Lurestan arc), during the early Miocene–Pleistocene. Propagation rates vary
778 through time and space between 34 mm/yr in an early stage close to the High Zagros Fault and 8
779 mm/yr close to the Zagros front.

780 • High rates of folding propagation are about one order of magnitude larger than age
781 equivalent shortening rates, ~4 mm/yr in Fars and 2 mm/yr in Lurestan, indicating the existence of
782 an efficient detachment level at the base of the folded sedimentary cover enhancing the stress
783 propagation towards the foreland.

784 • Numerical experiments of basement-involved fold-and-thrust belts, based on the Fars arc
785 parameters, show in-sequence folding comparable to results from magnetostratigraphic ages of
786 folding across the Zagros Simply Folded Belt. In experiment without inherited faults, the onset of
787 folding is roughly continuous and each anticline grows during few million years whereas the
788 experiment with inherited basement faults displays a more complex scenario with discontinuous
789 onset of folding, longer duration of anticlines growth and local out-of-sequence propagation of
790 deformation.

791 ACKNOWLEDGEMENTS

792 This research was funded by the DARIUS Programme as well as the Group of Dynamics of the
793 Lithosphere (GDL), Geosciences Barcelona, Geo3Bcn, CSIC, Spain. We thank the Paleomagnetic
794 Laboratory CCiTUB-Geo3Bcn, CSIC where the paleomagnetic analyses were conducted, and also
795 the Laboratorio de Geocronología y Geoquímica Isotópica of the Universidad Complutense de
796 Madrid, Spain, where the Sr samples were processed. In addition, the study of calcareous
797 nannofossils was supported by the Research Group of the Basque University System IT-930-16.
798 We are also grateful to the National Iranian Oil Company (NIOC) for their support during
799 fieldwork. We thank DARIUS Programme and its sponsors, and projects Alpimed (PIE-CSIC-
800 201530E082) and Subtetis (PIE-CSIC-201830E039) for additional support. We are grateful to
801 Stefano Tavani and an anonymous reviewer as well as to the Associate Editor for their constructive
802 reviews, which greatly helped to improve the quality of this paper. We dedicate this work to the
803 memory of Mohammad Sepehr, a lover of life and geology, who died unexpectedly during our first
804 reconnaissance trip to the study region.

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FIGURE CAPTIONS

Figure 1. A) Location of the Zagros foreland Fold Belt on the NE margin of the Arabian plate, B) Map of the Zagros tectonic domains and major basement faults. The current study area is located in the Dowlatabad growth syncline in the front of the Fars arc. The age of folding initiation of previously studied growth synclines are indicated on the map, including Shalamzar, Dehmoord (Pirouz et al., 2017), Chaman-Goli, Afrineh (Emami, 2008; Vergés et al., 2019), Changuleh (Homke et al., 2004), Jarik (Lashgari et al., 2020), Qalat (Khadivi et al., 2010), and Ghol Ghol (Ruh et al., 2014) synclines. These ages were determined based on the magnetostratigraphy of growth strata.

Figure 2. Regional structural cross-section extending from the Main Zagros Thrust to Saudi Arabia, across the High Zagros, Fars arc and Persian Gulf. The onshore segment closely follows the cross-sections by McQuarrie (2004) and Mouthereau et al. (2007). The depth of Moho is based on geophysical studies and numerical modeling (Paul et al., 2010; Jiménez-Munt et al., 2012; Tunini et al., 2015; Dashti et al., 2019).

Figure 3. A) Geological map of the Dowlatabad growth syncline neighbors by the Sefid Anticline to the north and the NW Pazan Anticline to the south (modified from 1:100,000 scale geological map series provided by NIOC). Location of paleomagnetic, strontium and calcareous nannofossil sampling sites are indicated on the map. B) CD geological cross-section constructed by projecting pre-, syn- and late-growth strata patterns across the Dowlatabad syncline and the NW Pazan Anticline.
Figure 4. A) Oblique satellite view and B) schematic evolution model of the Dowlatabad syncline with indication of the unconformities, growth strata units and sediment provenance. Unconformity U1 characterizes the base of growth strata, which define a gradual thinning of middle and upper Aghajari units from 1700 m to 500 m towards the NW, and the onset of folding in the Pazan anticline. The U2 is a local intra-growth angular unconformity at the base of first appearance of Guri Mb carbonate clasts derived from the dismantled crest of the Pazan anticline. The U3 represents a syncline-scale unconformity at the base of the late-growth gently dipping Bakhtyari conglomerate layers mostly composed of volcanic and metamorphic clasts from hinterland.

Figure 5. A) Stratigraphy of the Miocene-Pleistocene foreland deposits in the Dowlatabad growth syncline. B) Very thick-bedded to massive carbonates of the Guri Mb. C) Flute cast paleocurrent indicators in medium to coarse grained sandstones and mudrocks of the Aghajari Fm. D) Monomictic carbonate conglomerates and pebbly sandstone of the topmost part of the Aghajari Fm., with sub-angular clasts from the Guri Mb. E) Polymictic well-rounded clast-supported conglomerates of the Bakhtyari Fm., showing imbrication and containing sandstone channels.

Figure 6. A) Demagnetization diagrams of representative samples for the Mishan, Aghajari and Bakhtyari formations showing normal (DW203-1A, DW130-1A) and reversed (DW211-1A, DW136-1A) polarities. In red, the characteristic component (ChComp) with indication of the selected range of temperatures used to calculate it. BTC: before tectonic correction; ATC: after tectonic correction. B) Equal-area stereographic projection of the characteristic components (quality 1 and 2) and the obtained mean directions before (BTC) and after (ATC) tectonic correction with the 95% confidence angle. N: number of directions; dec: declination; inc: inclination; k: precision parameter; α95: confidence angle; rsum: module of the sum vector for the N unitary vector directions. It can be observed that directions are better group ATC and that they do not reveal any vertical axis rotation.
Figure 7. Stratigraphic log of the studied section with the location of magnetostratigraphic samples. Measured paleocurrent directions and changes in clast composition are also indicated. In addition, declination, inclination and latitude of the virtual geomagnetic pole, local magnetostratigraphy and the proposed correlation to the Global Polarity Time Scale (GPTS) (Gradstein and Ogg, 2012) are presented respectively from left to right. Circles represent class 1 and 2 samples and white squares are used for class 3 samples. Half bars show single site magnetozones, not considered as reliable.

Figure 8. Integration of magnetostratigraphy, Sr isotopic and calcareous nannofossil ages for about 3.5 km of foreland sediments in the Dowlatabad section. Local magnetostratigraphy for the topmost Mishan, Aghajari and Bakhtyari formations is correlated to the GPTS2012 (Gradstein and Ogg, 2012). Sr isotopic ages for the topmost Gachsaran and Mishan formations are plotted on the LOWESS curve of McArthur et al. (2001). The mean values, illustrated by small circles and error ranges, are shown by bars. The vertical bars show uncertainty of Sr isotopic ratios and horizontal bars indicate the age uncertainty for each sample. Calcareous nannofossils are assigned to their specific biozones. Using the age ties, rate of sediment accumulation is calculated for each depositional interval.

Figure 9. Setup of the initial stage of numerical models. From bottom to top, each model consists of 25 km of diabase basement, 2 km of low-viscous detachment (equivalent to Hormuz salt), and 10 km of upper sedimentary crust with two weak interlayers. All model parameters are given in Table S1. (A) shows a thin-skinned model without inherited basement faults, while (B) represents a thick-skinned model with listric inherited basement faults.

Figure 10. Temporal evolution of numerical experiments without (A–C) and with listric inherited basement faults (D–F). Upper panels: Composition of rock markers. See Fig. 9 for color code. Lower panels: Second invariant of the strain-rate tensor. Blue colors indicate fast deformation, white denotes slow deformation.
Figure 11. A) Revised Miocene-Pleistocene foreland chronostratigraphy for the SE Zagros (Fars arc) constructed by synthetizing the results of this study with previously published magnetostratigraphic sections (Khadivi et al., 2010; Ruh et al., 2014; Pirouz et al., 2017). Sedimentation rates (cm/kyr) for the different sections in SE Zagros with indication of time of growth initiation are also represented below the chronostratigraphic chart. B) Foreland chronostratigraphy for the NW Zagros (Lurestan arc) based on the previously published magnetostratigraphic ages (Homke et al., 2004; Pirouz et al., 2017; Vergés et al., 2019). Beginning time of deformation according to the base of growth strata and propagation rates of deformation front (pr) are represented on the chronostratigraphic charts (A, B).

Figure 12. Deformation propagation patterns obtained from numerical experiments. A) Experiment without pre-existing basement faults. B) Experiment with two listric pre-existing basement faults. Upper panels: Composition of rock markers after 15 Myr of shortening. See Fig. 9 for color code. Lower panels: Temporal evolution of horizontal shortening rate within the cover sequence.

Figure 13. Map depicting the foreland propagation of deformation across the Zagros Fold Belt during early Miocene to Pleistocene (modified from Vergés et al., 2019). In addition, onset of folding in Persian Gulf is indicated for the Hormuz salt diapirs and the N-S Arabian folds. The map is defined based on well-dated growth synclines including Shalamzar, Dehmoord (Pirouz et al., 2017), Chaman-Goli, Afrineh (Emami, 2008; Vergés et al., 2019), Changuleh (Homke et al., 2004), Jarik (Lashgari et al., 2020) Qalat (Khadivi et al., 2010), Ghol Ghol (Ruh et al., 2014), and Dowlatabad (this study) synclines. The age of the folding onset is marked for each growth syncline. The out-of-sequence reactivation of the Mountain Front Fault is highlighted. Locations of two foreland chronostratigraphic panels of Fig. 11 are illustrated on the map.
<figure><img src="image_2.png" /><figcaption>Fig. 1</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>Fig. 2</figcaption></figure>
<figure><img src="image_4.png" /><figcaption>Fig. 3</figcaption></figure>

B) Cross-section of the Dowlatabad syncline

<figure><img src="image_5.png" /></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 4</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>Fig. 5</figcaption></figure>

<figure><img src="image_8.png" /></figure>

<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>
<figure><img src="image_11.png" /><figcaption>Fig. 6</figcaption></figure>

<figure><img src="image_12.png" /></figure>
<figure><img src="image_13.png" /><figcaption>Fig. 7</figcaption></figure>

Base: 28°076883 N, 52°327367 E
<figure><img src="image_14.png" /><figcaption>Fig. 8</figcaption></figure>
<figure><img src="image_15.png" /><figcaption>Fig. 9</figcaption></figure>
<figure><img src="image_16.png" /><figcaption>Fig. 10</figcaption></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>
<figure><img src="image_22.png" /><figcaption>Fig. 11</figcaption></figure>

<figure><img src="image_23.png" /></figure>
A) No inherited basement faults

<figure><img src="image_24.png" /></figure>

B) Deep-seated inherited listric faults

<figure><img src="image_25.png" /></figure>
<figure><img src="image_26.png" /><figcaption>Fig. 13</figcaption></figure>