24

Lithospheric folding is an important mode of basin formation in compressional intraplate settings.

25

Basins formed by lithospheric folding are characterized by distinct features in subsidence history. A 26 comparison with extensional basins, foreland basins, intracratonic basins and pull-apart basins 27 provides criteria for the discrimination between these modes of basin formation. These findings are 28 important in deciphering the feed-backs between tectonics and surface processes. In addition, 29 inferences on accommodation space and thermal regime have important consequences for 30 hydrocarbon maturity. Lithospheric folding is coupled to compressional basin and fault reactivation 31 and, therefore, strongly affects reservoir characteristics of sedimentary basins. 32 33 34 models for several basin types have significantly advanced the understanding of extensional 39

Over the last few years much progress has been made in understanding the mechanisms of the 37 formation of sedimentary basins [START_REF] Cloetingh | Tectonic Models for the Evolution of Sedimentary Basins[END_REF][START_REF] Royden | Flexural behaviour of the continental lithosphere in Italy: constraints imposed by gravity[END_REF]. Quantitative

153

The inferred wavelength of these neotectonic lithosphere folds is consistent with the general 154 relationship between the wavelength of lithospheric folds and the thermo-tectonic age of the 155 lithosphere shown by a global inventory of lithospheric folds (Fig. 3 ) [START_REF] Cloetingh | Thermomechanical structure of European continental lithosphere: 1152 constraints from rheological profiles and EET estimates[END_REF] 156 Cloetingh et al., 2005). In some other areas of continental lithosphere folding, smaller wavelength 157 crustal folds have also been detected [START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF][START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF]. Thermal thinning of 158 the mantle lithosphere, often associated with volcanism and doming, enhances lithospheric folding 159 and appears to affect the wavelengths of folds [START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]. and south of the Tien Shan ranges [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]. In this area lithosphere underwent 169 Jurassic reactivation and is characterized by relatively young thermo-mechanical ages (175 Ma) 170 [START_REF] Burg | Shortening of analogue models of the continental lithosphere: 1102 New hypothesis for the formation of the Tibetan plateau[END_REF][START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]. In addition to having undergone thermal weakening, the 171 lithosphere underlying these basins also probably has a weak, quartz-dominated lower crustal 172 rheology. This has resulted in low values for the effective elastic plate thickness (EET) of the order 173 of 15 km [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF].

174

The Central Tien Shan consists of an alternation of ranges and basins separated by reverse faults. as much as 8 km of Cretaceous-Cenozoic sediments (Cobbold et al., 1993). Gravity data suggest 179 that the Ferghana and Tadjik basins are gravitationally overcompensated. The negative Bouguer mantle faulting (seen as Moho offsets in seismic reflection data).

190

Pre-existing thermal structure and variations in crustal thickness have played a major control on the 191 styles and distribution of faulting in this region (Fig. 4), (see also [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]Cobbold et 192 al., 1993). Primary faults appear before the folding develops with a spacing that is proportional to 193 brittle layer thickness. Subsequently the two processes, faulting and folding, co-exist in such a way 194 that faults localize at the inflection points of folds. At this stage, the appearance of faults does not 195 significantly influences the wavelength of folding. Because of the weakness of the lower crust, the 196 upper crust should be completely decoupled from the mantle and interact with it only by flow in the 197 lower crust. Lower crustal flow changes the wavelength and amplitude of the surface folding, which 198 is terminated by the development of a single down-warped megafold [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. Liu et estimated to be 250 Ma, as the main thermal perturbation was associated with a major orogenic 203 phase at 250 Ma ago. These findings are consistent with models in which decoupled lithosphere 204 folds with 30 km and 400 km wavelength correspond to, respectively, crustal and mantle folding.

205

Deep seismic data for basins north of and within the Tibetan plateau (Li et al., 2006;Liu et al., 2006; 206 Zhao et al., 2006) show that folding is not limited to the Tibetan plateau (see also [START_REF] Shin | Three-dimensional fold structure of the Tibetan Moho from GRACE gravity data[END_REF].

208

Folded lithosphere of Iberia

209

As reviewed by [START_REF] Cloetingh | Lithospheric folding in Iberia[END_REF], Tertiary lithospheric folding of Iberia occurs in Variscan 210 lithosphere with wavelengths of 300 km, leading to the development of a system of parallel trending 211 basins and highs (Fig. 5a). This folding generated the distribution of basins and mountain chains,

212

bounded by folds and faults [START_REF] Cloetingh | Lithospheric folding in Iberia[END_REF][START_REF] De | Cenozoic 1206 thick-skinned deformation and topography evolution of the Spanish Central System[END_REF]. The regularity of the

217

A large body of geophysical and geological observations is available on the crustal geometries and 218 stress regimes of Iberia [START_REF] De | Cenozoic 1206 thick-skinned deformation and topography evolution of the Spanish Central System[END_REF], 2008). This is supplemented by thermo-219 geochronology data (De Bruijne & Andriessen, 2002;Ter Voorde et al., 2004)

258

the western part of the plate (De Vicente et al., 2008;[START_REF] De | Large-scale distributed deformation controlled topography along the 1202 western Africa-Eurasia limit: Tectonic constraints[END_REF].

259

In general the dimensions of peneplains linked to extensional basins and passive margins are

267

The European Alpine foreland

268

Geophysical and geomorphological studies (e.g. Lefort & Agarwal, 1996;2000;2002;Marotta et al., 269 2000) indicate large-scale compressional deformation of the European Alpine foreland lithosphere.

270

Although different in structural grain and rheological structure (Tesauro et al., 2007;Cloetingh et (Marotta et al., 2000). The basin underwent a 275 polyphase history [START_REF] Scheck | Evolution of the Northeast German Basin -inferences of a 3D structural 1461 model and subsidence analysis[END_REF]Littke et al., 2008) of extension in the late Triassic 276 followed by lithospheric scale folding in the late Cretaceous-Early Cenozoic (Mazur et al., 2005;277 Marotta et al., 2000). 3D volumetric analysis shows that the compressional event was followed by 278 faster subsidence during the Cenozoic [START_REF] Scheck | Evolution of the Northeast German Basin -inferences of a 3D structural 1461 model and subsidence analysis[END_REF]. A similar observation was made by 279 Van Wees & Cloetingh (1996) (Horvath & Cloetingh, 1996;Cloetingh et al., 2005Cloetingh et al., , 2006;; 314 Horvath et al., 2006;Matenco et al., 2007;Fig. 7a). The river network in the area is affected by 315 differential neotectonic crustal motions (Necea et al., 2005;Dombradi et al., 2010).

316

The Late Neogene evolution of the Pannonian Basin is largely controlled by its interaction with the

325

The Pannonian Basin is underlain by what is probably the hottest lithosphere of continental Europe,

326

characterized by a very low lithospheric strength (Tesauro et al., 2007). Miocene extension resulted 327 in a high level of mantle stretching, leading to a strength distribution with practically zero strength in 328 the subcrustal lithosphere and lower crust (Cloetingh et al., 2005), making it prone to tectonic

340

Quaternary patterns of differential uplift and subsidence with a timing coeval with an increased level 341 of compressional stresses are interpreted as caused by lithospheric folding (Horvath & Cloetingh, 342 1996;Cloetingh et al., 2006;Horvath et al., 2006;Bada et al., 2007). Late-stage compressional

362

The South Caspian Basin with an estimated sedimentary thickness up to 20 km [START_REF] Brunet | The South Caspian Basin: a 1096 review of its evolution from subsidence modelling[END_REF] 363 Egan et al., 2009), is probably one the deepest basins in the world. It is flanked by the Alborz

364

Mountains to the south, with an elevation of 2-4 km (see Fig. 8). The transition between basin and 365 flanking high is abrupt and coincides with a coastline bounded by a major fault system. The 366 formation of the South Caspian Basin is enigmatic. Attempts to explain it as an extensional basin or 367 as a foreland basin have been unsuccessful. There is a lack of evidence for basin extension and 368 both stretching models and topographic loads and slab pull operating on down flexed lithosphere in 369 foreland basin systems cannot explain differential vertical motions of this order of magnitude. As a 370 result, phase changes in continental crust have been proposed as a mechanism for the formation of 371 this superdeep basin (Artyushkov, 2007). A limitation in conclusively resolving the basin formation 372 mechanism is the lack of multi-channel seismic reflection data capable of imaging crustal structure 373 below the thick sedimentary sequences. Over the last few years, however, other constraints on the 374 area have become available (e.g. Guest et al., 2007) (Kaban et al., 2002). Seismic tomography shows slab 382 detachment under the Iranian plateau (Hafkenscheid et al., 2006;Alinaghi et al., 2007) [START_REF] Chu | Current plate motions across the Red Sea[END_REF] all initiated around 5 ± 402 2 Ma (e.g. Axen et al., 2001;[START_REF] Smit | The rift-like structure and asymmetry of 1477 the Dead Sea Fault[END_REF], was interpreted by these authors in terms of a 403 widespread tectonic event. Slab detachment could explain part of the recent uplift of Iran, but does 404 not directly explain the simultaneous dramatic subsidence in the adjacent Caspian Basin. However, 405 slab detachment could have an indirect effect if reducing slab pull forces acting on the downgoing 406 lithosphere promote the development of compressional stresses in both the downgoing and 407 overriding plates [START_REF] Cloetingh | Thermo-mechanical controls on the mode of continental collision in 1173 the SE Carpathians (Romania)[END_REF]Matenco et al., 2007).

408 Guest et al. (2007) gravity data for the basin and followed a similar proposition made for the late-stage deformation of 411 the adjacent Black Sea [START_REF] Nikishin | The Black Sea Basin: tectonic 1415 history and Neogene-Quaternary rapid subsidence modelling[END_REF][START_REF] Cloetingh | Post-rift 1181 compressional reactivation potential of passive margins and extensional basins[END_REF]. Guest et al. (2007) present 412 kinematic models, without addressing the reason for the difference in the order of magnitude of the 413 basin depression and the adjacent high. Analogue experiments [START_REF] Sokoutis | Lithospheric-scale structures 1482 from the perspective of analogue continental collision[END_REF] show that 414 when compression acts on two blocks with contrasting thickness (and rigidity), a major syncline will 415 develop on top of the suture zone separating these blocks, flanked by an anticline of much lower 416 amplitude. These results are consistent with predictions from folding theory that as a result of the 417 acting gravity field, down-warping is mechanically more effective than up-warping against the action 418 of gravity forces. The presented numerical experiments illustrate this (see for example Fig. 9).

420

Basins on folded lithosphere in Eurasia: differences and similarities in lithospheric

429

The Central Asian Basins and the basins of the Iberian plate were initiated by Tertiary folding, 430 whereas the European Alpine foreland underwent a polyphase history, following Carboniferous 431 and Triassic stretching [START_REF] Scheck | Evolution of the Northeast German Basin -inferences of a 3D structural 1461 model and subsidence analysis[END_REF]Le Solleuz et al., 2004), prior to Late Cretaceous-

432

Early Cenozoic folding. Similarly, the Pannonian Basin, formed by back-arc extension in Miocene 433 times, was subsequently compressionally reactivated [START_REF] Cloetingh | Post-rift 1181 compressional reactivation potential of passive margins and extensional basins[END_REF]

469

Folding is commonly associated with periodic deformation of layered structures with contrasting 470 mechanical properties. As a consequence, the harmonic spatial periodicity is often believed to be its 471 primary, almost synonymous, identifying factor. However, from a mechanical point of view, constant 472 wavelength is not a necessary requirement. Folding is a compressional instability developing in stiff 473 layers embedded in a weaker material. The harmonic solution of the equilibrium and conservation 474 equations is a possible solution (Muehlhaus et al., 1998;[START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]Cloetingh et al., 475 1999). The physical mechanism of folding is well understood. In a continuous layered medium the 476 stresses and strains must be continued across the interfaces between the layers. In multi-layer 477 systems with contrasting mechanical properties (e.g. strong and weak layers) this requirement may 478 be difficult to satisfy, because the same amount of shortening in a stiffer layer would require much 479 larger stress than in the neighbouring weaker layer. The system becomes unstable and, in 480 attempting to reduce the stress and strain unconformities at the interfaces between the layers, starts 481 to fold (or buckle) in response to very small perturbations. In non-elastic media with strain-482 dependent properties, locally increased flexural strain at the fold limbs can create weakened zones 483 that significantly facilitate further deformation. These weak or softened plastic or viscous zones are 484 often referred to as inelastic hinges, since the system easily folds at such weakened zones. 489 is likely to be a primary and common response to tectonic compression. Folding probably may 490 continue, in an attenuated form, even after the beginning of subduction [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF], or 491 re-appear when the subduction channel is locked or during the aftermath of collision (Cloetingh et 492 al., 2004;Matenco et al., 2007).

493

The rheological structure of the lithosphere (Watts & Burov, 2003) is a key factor in lithospheric 494 folding. Compressional stresses must build up to a level comparable with the integrated strength of 495 the lithosphere in order to induce folding. Prior to the advent of rheological models based on 496 extrapolation of laboratory experiments [START_REF] Goetze | Stress and temperature in the bending lithosphere as constrained by 1257 experimental rock mechanics[END_REF][START_REF] Carter | Flow properties of continental lithosphere[END_REF]) lithospheric 497 stresses were thought to be incapable of reaching the failure levels required for folding (Turcotte & 498 Schubert, 2002).

499

The simplest expression of the folding process is in oceanic lithosphere, characterized by an 500 absence of a rheological stratification, and with thermo-mechanical ages spanning a limited age 501 window with a maximum of 200 Myr (Geller et al., 1993;Zitellini et al., 2010). Because weak 502 lithosphere is easier to fold than strong lithosphere, one would expect the first observations of 503 lithospheric folding from the rheologically weak continental plate interiors. However, the observation 504 of folding of the relatively strong lithosphere of the central Indian Ocean with a thermo-mechanical 505 age of 80 Ma and with wavelengths of 200-250 km [START_REF] Geller | Heat transfer and intraplate deformation in the 1248 central Indian Ocean[END_REF]McAdoo & Sandwell, 1985; 506 Stein et al., 1989;[START_REF] Bull | Seismic reflection images of intraplate deformation, central Indian 1099 Ocean, and their tectonic significance[END_REF] triggered the interest in this mode of deformation. There 507 may be two reasons for this. First, the recognition of folding, by mapping of basin reflectors and 508 gravity anomalies requires spatial scales of several hundreds of kilometres. This is feasible with 509 marine geophysical surveys, but generally in excess of the spatial scales covered by land surveys.

510

An exception has been the systematic mapping of basins in Central Australia, located on strong 511 lithosphere with a thermo-mechanical age of approximately 700 Ma (Lambeck, 1983;Stephenson & 512 Lambeck, 1985). The Indo-Australian plate is under an exceptionally high level of compression 513 (Stein et al., 1989), as a result of its collision with the Eurasian plate. Under these conditions high stresses might have been more important than the rheology of the plate interior (Beekman et al., 515 1996;[START_REF] Gerbault | Do faults trigger folding in the 1251 lithosphere?[END_REF][START_REF] Gerbault | At what stress level is the central Indian Ocean lithosphere buckling? Earth[END_REF].

516

Secondly, folding in oceanic lithosphere is not affected by surface erosion, making it relatively easy 517 to recognize. The presence of only one dominant folding wavelength characteristic of non-stratified 518 oceanic lithosphere also facilitated a quantitative interpretation. Following the early studies in the 519 Indian Ocean and in Central Australia in the rear of the Indian-Eurasian collision, attention shifted to 520 the other side of the plate contact in Central Asia. Large scale folding patterns were recognized 521 through analysis of geophysical data, including gravity and topography [START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF]Nikishin 522 et al., 1993;[START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF] as well as detailed geological studies (Cobbold et al., 1993).

523

On the basis of an inventory of currently available results of studies of folded continental 524 lithosphere, it appears that in general the observed wavelengths of folding follow theoretical 525 predictions for wavelengths as a function of thermo-mechanical age remarkably well (see Fig. 3). Folding is an unstable deformation that may develop in stratified media with significant competence 532 contrasts, as occur in the continental lithosphere, in response to horizontal loading [START_REF] Biot | Theory of Folding of Stratified viscoelastic media and its implications in tectonics and 1082 orogenesis[END_REF] 533 [START_REF] Ramberg | Contact strain and folding instability of multilayered body under compression[END_REF]. These instabilities may develop under compression or basal shear as a result of 534 strain/stress incompatibility on the interfaces between layers of different competence. In the initial 535 stages of deformation, compressional instabilities cause periodic folds with an exponential growth 536 rate and a dominant wavelength  roughly proportional to 5 -10 times the thickness(es) of the 537 competent layer(s). In the lithosphere, the layer thickness is that of the competent core h. So, h =

541

and where the conditions of the linear theory are more or less satisfied (e.g. thin layer 542 approximation, no strain-softening, plane layers). For linear folding in a Newtonian medium, an 543 asymptotic relationship derived from the thin layer equilibrium equation is [START_REF] Biot | Theory of Folding of Stratified viscoelastic media and its implications in tectonics and 1082 orogenesis[END_REF]Ramberg, 544 1961):

545  l = 2h ( l1 /6 l2 ) 1/3 (1)
Here  l is the "Laplacian" dominant wavelength of folding in the absence of gravity, h is the 546 thickness of the competent layer (crustal or mantle), and  l1 and  l2 are the effective viscosities (or 547 competencies) of the strong layer and weak surrounding material, respectively. The above 548 equation, derived assuming zero gravity, provides estimates of  l /h = 20-40 for typical competence 549 contrasts. This does not hold for most lithospheric-scale cases, where /h = 4 -6 is more common 550 due to the influence of the gravity-dependent terms. A reduction of the normalized characteristic 551 wavelength to values around 4 is also predicted for power-law rheology or plastic layers (Smith, 552 1979). For a single stiff viscous layer on top of an inviscid medium, the dominant gravity-dependent 553 wavelength is [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]:

554  g ~ 2(2h    eff / (g)) 1/2 (2)
where  is the density contrast,   is the strain rate,  eff is the layers' effective ductile viscosity 555 (see Appendix A), n is the power-law exponent, and h is the thickness of the competent layer.

556

At large amounts of shortening this deformation may become a-periodic (Hunt et al., 1996) so that 557 the wavelength and amplitude of folding vary along the deformed transect. Plume-induced periodic 558 deformation [START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]) also includes folding or boudinage caused by shear on the 559 base of the lithosphere. However, the prime cause for folding of continental lithosphere is tectonic 2) predict folding wavelengths that may vary from 30 km (upper crustal folding) to as high as 600-563 800 km. The findings of numerical models [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF] are in excellent agreement with 564 analogue tectonic experiments [START_REF] Sokoutis | Lithospheric-scale structures 1482 from the perspective of analogue continental collision[END_REF]. Numerical experiments show that the 565 wavelength and amplitude of folding is largely controlled by the thermo-mechanical age and less by 566 the rate of shortening [START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]. Lithospheres older than 500 Ma develop nearly 567 the same wavelength of folding as 1000 Ma old lithosphere because of the similar nearly stationary 568 thermal state. Here, therefore, we restrict ourselves to a set of experiments for ages of 150 and 300 569 Ma and convergence rates of 1.5 and 3 cm/yr.

570

Surface topography and sedimentary deposition is described through diffusion-like erosion laws and 571 fluvial transport laws (Gossman, 1976;Kirkby, 1986;Leeder, 1991;Beaumont et al., 1992;Kooi and 572 Beaumont, 1994):

573 dh/dt = div(k * (x,h,h) h ) (slope erosion)
where h is topography, t is time, x is the horizontal coordinate, k * is an experimentally established 574 scale-dependent coefficient of erosion (500 -1000 m 2 /y). For simplicity, we used zero order linear 575 diffusion (k * constant) for the short-range erosion and flat deposition for long-range fluvial transport 576 (Avouac & Burov, 1996). The hill-slope erosion law ensures conservation of matter but holds only 577 for small-scales. At large scales it is no longer applicable, and conservation of matter is not 578 observed, because some sediments are lost due to transport out of the system by long-range 579 processes.

580

Figure 9 shows the results of the experiments for relatively young (150 Ma) lithosphere compressed 581 at rates of 1.5 and 3 cm/yr. Strongly decoupled low-amplitude short wavelength (mantle wavelength 582  = 250 km) folding developed leading to crustal thickening above a synclinal mantle fold. At later 583 stages, harmonic folding was followed by mega-folding [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]Burg & 584 Podladchikov, 1999). For higher rates of shortening these mantle synclines developed in a mode

587

Folding is well developed for medium (300 Ma) age lithosphere (Fig. 10). This case is characterized 588 by long mantle wavelengths (=360 km) and high surface amplitudes (2000 m after 10 Myr). These 589 compare well with observations from Iberia (Fernandez-Lozano et al., 2010), Western Goby and the 590 Ferghana basin in Central Asia [START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF][START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]. At late stages (10 -26

591

Myr since the onset of shortening for 3 cm/y or 20-50% of shortening), folding becomes aperiodic, 592 leading to mega-folding [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF][START_REF] Burg | Lithosphere scale folding: numerical modelling and application to 1105 the Himalayan syntaxes[END_REF]

596

Faleide, 2009). However, it may be relatively rare for folding to continue for periods in excess of 10 597 Myr. More typically is that at a certain stage deformation localizes along single major fault zones 598 [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF][START_REF] Gerbault | Do faults trigger folding in the 1251 lithosphere?[END_REF].

599

As was discussed by [START_REF] Bird | Lateral extrusion of lower crust from under high topography in the isostatic limit[END_REF], Avouac & Burov (1996) and [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF], large-scale 600 undulations of the lithosphere cannot be preserved for a long time (longer than 10 Myr) in the 601 absence of sufficient compression, except for plates with very strong (especially lower crustal) 602 rheology. Otherwise, the folds either will be flattened by gravity-driven crustal flow associated with 603 the large crust-mantle density contrast at the Moho, or deformation will localize along some of the 604 faults created at the inflection points of folds.

605

The fact that folding with wavelengths of 400-500 km and 500-600 km persists long after cessation 606 of the compression in the presumably high-rigidity Australian craton and Russian platform points to 607 a strong rheology yielding a high value for the effective elastic thickness EET (in excess of 60 km),

608

as expected from its thermo-mechanical age [START_REF] Beekman | Mechanical stability of the Redbank Thrust 1076 Zone, Central Australia: dynamic and rheological implications[END_REF]. The long wavelengths of 400-

628

Central Asia [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF].

629

This symmetry is in marked contrast to the asymmetrical shape of foreland basins, formed by 630 flexure in front of an orogenic wedge, which are flanked by a flexural bulge of an amplitude that is 631 only up to 10 percent of the maximum depth of the foreland depression [START_REF] Royden | Flexural behaviour of the continental lithosphere in Italy: constraints imposed by gravity[END_REF]; Zoetemeijer

648

Lithospheric folding results from an instability in the lithosphere due to stress/strain incompatibilities 649 that develop in rheologically stratified layers under compressive strain. This process typically 650 operates with timescales of the order of 1-10 Myr [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. The preservation of 651 lithospheric folds appears to depend strongly on the thermo-mechanical age of the underlying 652 lithosphere [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. Folds in young lithosphere will not be preserved upon relaxation 653 of the stress field, whereas ones in cratonic lithosphere will be well preserved (Cloetingh et al., 654 1999).

677

Thus, folding appears to be more effective in middle aged lithosphere of 300 Ma than in lithosphere 678 of younger ages [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF][START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]. After 2 Myr, a slow uplift phase 679 starts of relatively minor magnitude of the order of a few hundreds of meters to a km. The predicted 680 subsidence in Figure 11 is for the centre of the syncline. As pointed out above, erosion of the 681 uplifted flanks is taken into account, adopting a diffusive equation approach to erosion (see Burov & 682 Cloetingh, 1997;[START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF], for further details).

683

The alignment of parallel highs and basins of similar dimensions has an important consequence for 684 the areal extent of the source areas for sediments available for deposition in the folded depressions.

685

In comparison with foreland basins the distribution of sources is more symmetrical, superseding the 686 volume of sediments that can be eroded from, for example, the flexural foreland bulge. As noted

689

accommodation space in changing the spectrum of wavelengths [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. Erosion 690 acts as a filter, suppressing the shorter wavelengths in folded basin topography. Strong erosion, 691 insufficiently compensated by tectonic deformation, can even wipe out most of the topography.

692

However, if the erosion is tuned to the average elevation rates (Fig. 12), it may dramatically 693 accelerate folding [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF].

694

There has been a rapid development and application of stratigraphic modelling packages to 695 sequence stratigraphy over the last two decades (e.g. Cross, 1989;Alzaga-Ruiz et al., 2008).

696

Among the different types of modelling programs, the forward stratigraphic modelling approach has 697 been applied extensively by basin modellers from academia and petroleum industry (e.g. Lawrence et al., 1990;Granjeon, 2009). Forward stratigraphic modelling commonly starts with an estimate of 699 the initial conditions related to an adopted basin formation mechanism and controlling parameters,

698

726

foreland basins, the initiation of folding is not associated with a thermal instability, unless folding 727 occurs in conjunction with plume activity [START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]. Following the deposition of 728 radiogenic sediments in the folded depression, sediment blanketing will affect the heat flow 729 (Stephenson et al., 1990;Lavier & Steckler, 1997;Van Wees et al., 2009)

736

An example is the Paris basin, where source rocks of Toarcian age were folded in Cenozoic time 737 [START_REF] Bonnet | Large-scale relief 1088 development related to Quaternary tectonic uplift of a Proterozoic-Paleozoic basement: The Armorican 1089 Massif, NW France[END_REF]Le Solleuz et al., 2004).

755

Modes of brittle deformation and faulting

756

The role of pre-existing faults on the development of folding has been investigated through 757 numerical experiments on folding of brittle-elasto-viscous lithosphere (Beekman et al., 1996; 758 [START_REF] Gerbault | Do faults trigger folding in the 1251 lithosphere?[END_REF][START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF] and through analogue experiments on plastic-elastic 759 lithosphere (e.g. [START_REF] Burg | Shortening of analogue models of the continental lithosphere: 1102 New hypothesis for the formation of the Tibetan plateau[END_REF][START_REF] Sokoutis | Lithospheric-scale structures 1482 from the perspective of analogue continental collision[END_REF]. These studies show that lithospheric

772

Following the initiation of folding, and the further build up of compressional stresses, folding will, 773 therefore, usually be accompanied by brittle deformation and faulting in the crust and upper 774 lithosphere. The surface expression of folding in uniform lithosphere is frequently in the form of 775 fault-controlled pop-up and pop-down structures and inverted basins (e.g. Cobbold et al., 1993),

776

accompanied by development of imbrications of bivergent thrust wedges at the areas where 777 shortening is applied [START_REF] Sokoutis | Lithospheric-scale structures 1482 from the perspective of analogue continental collision[END_REF]. These findings (see Fig. 14) are corroborated by 778 observations in the Central System of Iberia, where folding of Variscan lithosphere with a thermo-779 mechanical age of 280 Ma occurs [START_REF] Cloetingh | Lithospheric folding in Iberia[END_REF]. Here the brittle deformation associated 780 with crustal scale folding expressed by a folded Moho is mechanically decoupled from folding of the 781 mantle lithosphere (Fernandez-Lozano et al., 2010; see Fig. 5b).

782

Crustal and mantle faults may develop as a result of folding [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]) (Fig. 15).

783

Folding can continue after these faults develop and folding and faulting can co-exist for times of 784 several Myr (Fig. 5b). For crust-mantle decoupling, a complex distribution of crustal faults and 785 surface undulations can be expected, with crustal faults not necessarily connected to mantle faults.

786

Numerical experiments suggest that mantle faults are unlikely to be longer than a few km in depth 787 [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]. Such faults may also not develop at longer wavelengths of folding 788 connected with stronger lithosphere. However, deep seated faults in mantle lithosphere can develop 789 when cratonic lithosphere is folded [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. The mode of faulting, depth extent and of decoupling between crustal and mantle folding is crucial. Discontinuities within the folded 793 lithosphere will localize the development of deep synclines, flanked by more modest anticlines 794 [START_REF] Sokoutis | Lithospheric-scale structures 1482 from the perspective of analogue continental collision[END_REF] and affect the distribution of faulting (Fig. 14).

795

During the next phase of basin evolution, when a steady state is reached between the creation of 796 accommodation space and sediment supply, only minor faulting will occur. This is in marked [START_REF] Royden | Flexural behaviour of the continental lithosphere in Italy: constraints imposed by gravity[END_REF] is 807 thought to occur in interaction with phase changes in the lithosphere (Artyushkov, 2007). Examples 808 of these might be the Barents Sea (Ritzmann & Faleide, 2009) and possibly the South Caspian 809 Basin, although in the latter case pre-orogenic extension cannot be ruled out (Guest et al., 2007).

810

Basins exposed to changes in tectonic regime will have a polyphase record (Cloetingh & Ziegler, 811 2007), characterized by a superposition of more than two regimes, such as pre-orogenic extension, 812 foreland flexure and late-stage folding. This sequence appears to be characteristic for some very 813 deep foreland basins, such as the Focsani depression of the Romanian Carpathians, with more 814 than 16 km of sediments, which was affected by extension due to the opening of the Black Sea

821

The impingement of plumes on the base of continental lithosphere can induce differential 822 topography similar to the surface deflections characteristic for lithosphere folds (Guillou-Frottier et 823 al., 2007;Burov et al., 2007). Plumes and lithospheric folds frequently interact in space and time in 824 the geological record (Ziegler & Dèzes, 2007;[START_REF] Burov | Controls of mantle plumes and lithospheric folding on modes of intra-1114 plate continental tectonics: differences and similarities[END_REF]. Therefore, the evolution of 825 folded lithosphere basins can be overprinted by the signatures of a plume or vice-versa. This will 826 lead to an amplification of the induced vertical motions, particularly significant for young and European Cenozoic Rift System arrived about 1 Myr ago. This activity was preceded and followed 848 by lithospheric folding that continues to present-day since 17 Myr ago [START_REF] Bourgeois | Separation of rifting and lithospheric folding signatures in the NW-Alpine foreland[END_REF]. [START_REF] Cloetingh | Tectonic Models for the Evolution of Sedimentary Basins[END_REF]Xie & Heller, 2009). Figure 1b andTable Sedimentary basin systems are, by their nature, prone to tectonic reactivation and, therefore, 870 frequently characterised by a poly-phase evolution [START_REF] Cloetingh | Tectonic Models for the Evolution of Sedimentary Basins[END_REF]. Extensional basins 871 are, for example, often formed in areas previously thickened by tectonic compression and 872 subsequently often subject to late-stage compression (Lundin & Dorè, 1997). Examples include the 873 Pannonian basin of Hungary (Horváth & Cloetingh, 1996) and the North Atlantic rifted margins 874 [START_REF] Cloetingh | Post-rift 1181 compressional reactivation potential of passive margins and extensional basins[END_REF]. Similarily, foreland basins are frequently characterised by pre-orogenic 875 extension. Examples include the Carpathian foredeep (e.g. Tarapoanca et al., 2003) and the 876 Aquitaine basin of Southern France, which was the retro-arc foreland basin of the Pyrenees 877 [START_REF] Desegaulx | Consequences of foreland basin development on 1198 thinned continental lithosphere: application to the Aquitaine basin (SW France)[END_REF]. The actual subsidence patterns of these poly-phase systems are often 878 more complex than predicted by the end-member models discussed below that only consider the 879 basin formation mechanism. Retro-arc foreland basins are characterised by a prolonged subsidence 880 history with a distinct thermal signature including the effects of thermal cooling inherited from pre-881 orogenic extension. This is sometimes neglected in comparing pro-arc and retro-arc foreland basins 882 (Naylor & Sinclair, 2007).

883

Basin infill records allow main features, including geometry, vertical motions and faulting 884 characteristics to be unravelled (Zoetemeijer et al., 1993). However, heat flow histories need to be 885 determined through kinematic reconstruction and forward modelling. Results of heat flow history 886 modelling for FB and EB settings (Van Wees et al., 2009) can be subsequently compared to FLB 887 results presented in this paper (Fig. 13). For ICB and PAB we have refrained from a quantitative 888 approach for heat flow predictions as vertical motions driving subsidence cannot be linked in a 889 straightforward way to a single or uniformly distributed lithosphere process. As a consequence, for 890 these settings tentative heat flow patterns have been drawn.

891

Folded lithosphere basins (FLB) that develop as a result of periodic folding are highly symmetrical,

902

Foreland basins (FB) are strongly asymmetrical, with depths usually less than10 km, and have 903 modest flexural bulges of a few hundred meters (Royden, 1993;DeCelles & Gilles, 1996;904 Zoetemeijer et al., 1999). Important differences occur in the nature of the subsidence of pro-and 905 retro-foreland basins (Naylor & Sinclair, 2007). The first are typically associated with short life spans 906 and convex up subsidence patterns with time, whereas the latter show more prolonged subsidence 907 histories and concave up subsidence patterns. Subsidence patterns will be interrupted during 908 thrusting phases, also leading to uplift of the flexural marginal bulge [START_REF] Quinlan | Appalachian thrusting, lithospheric flexure and the Paleozoic 1425 stratigraphy of the eastern interior of North America[END_REF] 909 Zoetemeijer et al., 1993). The predicted heat flow pattern of a pro-foreland setting shown for the 910 Ebro foreland basin (well Jabali from Verges, 1999), marked by rapid basement subsidence and 911 sedimentary infill of ca 4 km in ca 10 My period (other parameters as in Van Wees et al., 2009).

912

Flexure and rapid sediment infill is marked by a reduction of heat flow, followed by a progressive 913 increase related to radiogenic heat production of foreland sediments. Basement fracturing is 914 concentrated at the points of maximum curvature in the flexed lithosphere, at the maximum points of 915 upward and downward deflection, respectively, with intensity higher than inferred for FLB.

916

Intracratonic basins (ICB) share their symmetrical shape and great thickness of their sedimentary

929

Heat flow patterns of ICB will be dominated by sediment infill. ICB are in general characterized by a 930 low intensity of faulting. However, in the absence of a rigorous understanding of the formation 931 mechanisms of ICB, predictions for heat flow are often not available. As a consequence, their heat 932 flow is most often described primarily in terms of the thermal consequences of their burial history.

933

Extensional basins (EB) have received major attention from basin modellers, both because of their 934 relative abundance and importance and because of the availability of a relatively simple quantitative 935 model for their formation (McKenzie, 1978). The latter can be relatively easily be incorporated in 936 forward modelling of their subsidence and heat flow histories (Van Wees et al., 2009). The basin 937 shape is symmetrical or asymmetrical depending on a pure shear or simple shear mode of 938 extension. Rifts can be narrow or wide, depending on the mode of rifting and the rheology of the 939 extending lithosphere (Huismans and Beaumont, 2002). Dependent on the rigidity of the extending 940 lithosphere, rift shoulders with asymmetrical shape can be formed and be prone to erosion (Van 941 Balen et al., 1995, Van der Beek et al., 1995). The development of rift shoulders and their erosion [START_REF] Cloetingh | Tectonic Models for the Evolution of Sedimentary Basins[END_REF]. In the latter case basin migration is a 953 common feature [START_REF] Corti | Transition from continental break-up to seafloor spreading: how fast, symmetric and magmatic[END_REF]. Extensional faulting occurs during the rift phase, with absence 954 of faulting in the post-rift phase. Heat flow reaches a peak during the syn-rift phase, followed by a 955 rapid decrease in heat flow (Van Wees et al., 2009). The syn-rift peak in heat flow is low or almost in the crust and cooling by sediment infill.

964

Pull-apart basins (PAB) are characterized by a narrow fault bounded geometry [START_REF] Smit | Pull-apart basin formation and 1474 development in narrow transform zones with application to the Dead Sea Basin[END_REF], 965 developing over a very short time span. Models have been developed to calculate subsidence 966 histories and heat flow histories for these basins [START_REF] Pitman | The effect of sea-level change on the shelf edge and slope of 1421 passive margin. In: The shelf break: critical interface on continental margins[END_REF], incorporating 967 lateral heat transport in stretching models. These models predict rapid subsidence of several km's 968 over time intervals of the order of 0.1 -1 Myr, followed by a decay in the subsidence with a concave 969 shape upward. Erosion and sediment flux can not keep pace with the subsidence at the basin 970 formation phase, leading initially to a starved basin, followed by a shallowing upward sequence

F o r R e v i e w O n l y
39 scenario is a good first order description, many PAB show a more punctuated fault controlled 974 subsidence history, often terminating by late-stage uplift, as observed for Late Neogene pull-apart 975 basins in the internal zone of the Betics in SE Spain [START_REF] Cloetingh | Flexural interaction and the dynamics of Neogene extensional basin formation in the Alboran-1162 Betic region[END_REF] and for the flanks of 976 pull-apart basins in the Dead Sea Fault System [START_REF] Smit | Pull-apart basin formation and 1474 development in narrow transform zones with application to the Dead Sea Basin[END_REF][START_REF] Smit | The rift-like structure and asymmetry of 1477 the Dead Sea Fault[END_REF].

977

From the above overview, it is apparent that FLB have in common with FB a compressional origin 978 and thermal histories controlled to a large extent by burial histories. At the same time, pronounced 979 differences occur in the duration of subsidence, varying from ultra short (FLB 1-10`s Myr), medium 980 (FB, tens of Myr) and ultra long (ICB, hundreds of Myr). An important difference occurs in the spatial 981 scales of the topography generated by folding and foreland fold-and-thrust belt deformation. In the 982 latter case a much steeper differential topography is generated over typical horizontal differences of 983 the order of 10 km between peak and trough. In contrast, in the case of lithospheric folding 984 horizontal distances between the topographic highs and the axes of the depocentres are typically of 985 the order of hundreds of kilometres. As a result, diffusive transfer of mass will occur at a much 986 slower rate of sediment supply in the case of folded lithosphere than rates of sediment supply 987 locally generated by thrusting. This together with a very rapid initial subsidence of folded lithosphere 988 basins has important consequences for the paleobathymetry of these systems.

989

FLB contrast with extensional and pull-apart basins in the shape of the subsidence curves, being

1629

We use a serial (Maxwell type) body (eq. 1), in which total strain increment in each numeric element 1630 is defined by a sum of elastic, viscous and brittle strain increments. Consequently, in contrast to 1631 fluid dynamic approaches, where non-viscous rheological terms are simulated using pseudo-plastic 1632 and pseudo-elastic viscous terms (e.g. [START_REF] Bercovici | A two-phase model for compaction and damage 3. 1079 Applications to shear localization and plate boundary formation[END_REF][START_REF] Solomatov | Scaling of time-dependent stagnant lid convection; application to 1485 small-scale convection on Earth and other terrestrial planets[END_REF], our 1633 method explicitly treats all rheological terms. The parameters of elastic-ductile-plastic rheology laws 1634 for crust and mantle come from rock mechanics data (Tables A1 andA2

1638

The brittle behavior of rocks is described by Byerlee's law [START_REF] Byerlee | Friction of rocks[END_REF][START_REF] Ranalli | Rheology of the Earth[END_REF] 1660

 eff = e d ij (1-n)/n A -1/n exp (Q (nRT) -1 ) ( A 8 )
For non-uniaxial deformation, the law (A8) is converted to a triaxial form, using invariant of strain 1661 rate and geometrical proportionality factors:

1662

 eff = e d ij (1-n)/n A -1/n exp (Q (nRT) -1 )
where 1663

e d II = (Inv II (e ij )) ½ and A * = ½A•3 (n+1)/2 ( A 9 )
parameters A, n, Q are experimentally determined material constants (Table A1). Using olivine 1664 parameters (Table A1), one can verify that the predicted effective viscosity at the base of the

1682

Central Indian Ocean Basin (CIOB) [START_REF] Geller | Heat transfer and intraplate deformation in the 1248 central Indian Ocean[END_REF]Stein et al., 1989) ; NW European platform 1683 [START_REF] Bourgeois | Separation of rifting and lithospheric folding signatures in the NW-Alpine foreland[END_REF]Lefort & Agarwal, 1996;Marotta et al., 2000); Pannonian Basin (Dombradi 1684 et al., 2010 ); Iberia [START_REF] Cloetingh | Lithospheric folding in Iberia[END_REF][START_REF] De | Cenozoic 1206 thick-skinned deformation and topography evolution of the Spanish Central System[END_REF]Fernandez-Lozano et al., 2010); 1685 Central Asia [START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF][START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF][START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]; Tibet/ Himalayan 1686 syntaxis belt (Burg & Podlachikov, 1999;[START_REF] Shin | Three-dimensional fold structure of the Tibetan Moho from GRACE gravity data[END_REF]; Central Australia (Stephenson & 1687Lambeck, 1985;Lambeck, 1983); Arctic Canada (Stephenson et al., 1990); Transcontinental Arch 1688 of North America [START_REF] Ziegler | Laramide foreland[END_REF]South Caspian Basin (Guest et al., 2007) [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. Black squares 1696 are estimates for wavelengths of differential topography and thermo-tectonic ages inferred from 1697 observational studies. Numbers refer to sites listed in Table 1.

1701

(1999). (b) Neotectonic movements along the profiles AD and CD [START_REF] Burov | Continental lithosphere folding in 1117 central Asia (Part II): Constraints from gravity and tectonic modelling[END_REF]. Modified significantly while the wavelength of folding stayed constant [START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF] A2. Specific rheology and related thermal parameters. Compilation by Burov et al. (2001)

  for Integrated Solid Earth Sciences, Faculty of Earth and Life 8 Sciences, VU University Amsterdam, De Boelelaan 1085, 1081 HV Amsterdam, The Netherlands.

  data, demonstrating prevailing fold wavelengths of 300 to 420 km(Shin et 152 al, 2009).

  folded continental lithosphere have characteristic features. We illustrate this 163 by largely concentrating on the basins of Central Asia, the folded lithosphere of Iberia, the NW 164 European platform, the Pannonian Basin and the South Caspian Basin. and Tadjik basins are compressional basins northwest and north of the Pamir

  168

175

  The Tien Shan terminates west of the right-lateral Talas-Ferghana fault by splaying into two narrow 176 surround the Ferghana Valley. A narrow belt of mountains parallel to the 177 central section of the fault slopes downward to the Ferghana Valley, underlain by a deep basin with 178

  199 al. (2007) reported evidence for lithospheric folding in the Central/ North China Tarim Basin region, 200 initiated during Pliocene times, with good time constraints provided by Ar/Ar dating. Two 201 wavelengths are found of 30 and 400 km. The thermo-mechanical age of the lithosphere is 202

  213 fluvial network pattern of the central-western part of Iberia also indicates large-scale lithospheric 214 folding. Alpine (mainly Cantabrian-Pyrenean related) compressional tectonics was the principal 215 factor for closing internally drained sedimentary basins, including the Ebro, Duero and Tagus basins 216 (Casas-Sainz & De Vicente, 2009).

  , demonstrating 220 accelerated uplift of the highs during Late Miocene-Pliocene times. In spite of the existence of 221 recent localized volcanism in the inner part of the Iberian plate, an important contribution from 222 thermal uplift is questionable due to the distribution of crustal thickness (Casas-Sainz & De Vicente, 223 2009). Fernandez-Lozano et al. (2010) have constructed a three-layer analogue model consisting of 224 brittle upper crust, ductile lower crust and ductile upper mantle for Iberia. These models show that 225 folding is associated with deformation of narrow mountain ranges separated by basins (see Fig.

  226 5b). The narrow mountain ranges representing upper crustal pop-ups form the main topographic 227 reliefs. Shortening is accommodated within the viscous crust beneath the pop-ups leading to lateral 228 the ductile crust. These results are consistent with seismic and gravity data Different patterns in evolution and stratigraphy between the eastern and western 255 Tagus basin may be related to the mechanical decoupling between Iberia and Africa during the 256 Miocene emplacement of the Alboran Domain in the southern part of the plate. This generated 257 extensional stresses in the eastern Iberian plate, coeval with compression and lithospheric folding in

  260 larger, and their evolution involves time scales much longer than those characteristic for high plains 261 of the Iberian plate (Casas-Sainz & De Vicente, 2009). Peneplains associated with compressional 262 or convergent regimes (e.g. Anatolia, Central Iran and Morocco) are characterized by spatial 263 dimensions comparable with Iberia. In addition, their tectonic setting in the hanging wall of the 264 African-Eurasian convergence zone is also similar. Thus the Iberian plate may be an analogue for 265 other present-day intra-mountain basins.

  al., 271 2005) inherited from differences in their Paleozoic-Mesozoic history, different segments of the 272 lithosphere in the northern Alpine European foreland share many similarities in their Cenozoic 273 intraplate deformation. Seismic and geological evidence shows that the North German Basin cannot 274 be explained by classical models of basin formation

  Fig.6b).

317

  Carpathian arc and the Adriatic indentor impinging the Pannonian basin lithosphere (e.g. Bada et al, 318 2001). As a result of the consumption of subductable lithosphere of the European foreland, the 319 Pannonian basin became locked. At the same time, the continuous N-NE directed indentation of the 320 Adriatic microplate built up a high level of compressional stresses in the Pannonian Basin. The 321 structural inversion of the basin is a direct consequence of the temporal changes in the stress field 322 from extension to compression. Recent studies of the stress field together with an extensive array of 323 geological and geophysical data have shown that horizontal stresses are transmitted from the plate 324 boundaries into the interior of the Pannonian Basin (Bada et al., 2007).

  329 reactivation. Therefore, in this area relatively low compressional stresses induced mainly by plate 330 boundary interactions, and further amplified by topography-induced stresses (Bada et al., 2001), 331 could initiate significant intraplate deformation. noted that in the Pannonian basin system, seismic tomography (Wortel & Spakman, 333 2000) supports slab detachment under the adjacent Romanian Carpathians. At the same time, the 334 seismic tomography shows the presence of a hot upper mantle below the folded lithosphere in the 335 Pannonian basin. In this area, therefore, the upper mantle structure reflects more the back-arc 336 extension prior to the folding than the consequences of folding on the density stratification of the 337 mantle. The reverse is true for the crustal configuration. The geometries imaged by the crustal scale 338 deep seismic reflection profiling are clearly dominated by compressional structures resulting from 339 folding (Fig. 7b).

  343 deformation of the Pennonian basin is demonstrated to be of key importance for the geometry of weak back-arc lithosphere of the Pannonian basin and the interplay between lithospheric processes 348 and the intrabasinal topography development. Based on available rheological and seismotectonic 349 constraints, the basin system is represented by a two-layer model with a 10-15 km thick brittle upper 350 crust. The lower crust and the mantle are represented by a low-strength layer of 45 km thickness 351 (Dombradi et al., 2010). The strain rates applied were constrained by GPS velocities (Grenerczy et 352 al., 2005). In the experiments lithospheric folds develop with wavelengths equivalent to 200-250 km 353 (Fig. 7c). The experiments also show that crustal thickness variations largely determine the 354 localisation of stresses. Topographic profiles extracted by laser scans of the model show that 355 surface topography reflects long wavelength buckling of the lithosphere. These results are 356 compatible with Pliocene-Quaternary subsidence, which occurs in the centre of the Pannonian 357 basin, simultaneously with accelerated uplift at the basin margins. The latter has lead to elevated

  deformation in Central Asia, Iberia, the NW European Alpine foreland, 423 the Pannonian Basin and the South Caspian Basin demonstrates common features compatible with 424 an interpretation in terms of basins developed on folded lithosphere. Amongst these are their 425 symmetrical shape, dimensions and the temporal association of their basin formation history with 426 the build up of intraplate stresses. Important differences can be observed between the Central 427 Asian and Iberian basins and the South Caspian Basin on one side and the European Alpine

428

  foreland and the Pannonian Basin on the other side.

460A

  general observation for all these basin systems is the flatness of sedimentary sequences in the 461 depocentres. Faulting plays a minor role in the basin development. Faulting appears to be basin margins, with a highly variable depth extent, varying from upper crustal 463 scale faults bounding pop-up structures in the Central System of Iberia and the Pannonian Basin to 464 minor faulting in the basement of the North German Basin. The thermal regimes of the basins are 465 similar with thermal histories strongly affected by subsidence and sediment deposition on down 466 flexed lithosphere, pre-conditioned by its thermo-mechanical age.

485

  Homogeneous shortening of the lithosphere under horizontal compression requires more work than 486 required by an equivalent amount of shortening by folding. However, shortening by 487 underthrusting/subduction of the lithosphere may be more mechanically efficient than folding but 488 or unable to start immediately after the onset of shortening. For this reason, folding

  538 10-100 km, gives rise to  = 50-1000 km. s", or linear, folding encompasses the cases where compression of the lithosphere takes place 540 from the beginning, leading to the formation of alternating (synclinal) basins and (anticlinal) highs,

  560loading by horizontal far-field forces. For characteristic parameters of various layers within the h = 10 -120 km,  = 10 19 -10 26 Pa s:  = 10 -600 kg m -3 ), equations (1-562

  subduction. For lower convergence rates (Fig.9a) folding does not develop 586 due to the low Peclet number of the system, leading to heat diffusion and weakening of folds.

  and subsequent 593 formation of high-amplitude crustal down-warps. The amplitude of vertical movements may reach 20 594 km ( 10 km) or even more. As pointed out earlier, such high amplitudes of vertical motions are 595 observed in the South CaspianBasin (Guest et al., 2007) and the Barents Sea(Ritzmann &

  609 600 km of crust-mantle down-warping in the eastern Barents Sea, where Mid-Cenozoic a basin system formed during Late Permian-Triassic time(Ritzmann & 611 Faleide, 2009), with a total infill of sediments of the order of 15-20 km is another impressive 612 example of compressional deformation on relatively cold lithosphere.613Most of the areas with present-day long-wavelength continental lithosphere folding formed during or 614 short after the Alpine collision 60 Myr ago. It thus seems that 60 Ma is a characteristic timescale of 615 gravity collapse of the large-scale folds in such intermediate-age lithospheres. In the case of very 616 weak quartz-dominated lower crust, folds may disappear rapidly within 8-15 Myr. Experiments 617 (Cloetingh et al., 1999) for 1000 Ma old lithosphere, with a diabase lower crust, demonstrate that 618 folding-induced deformation persists for 10-50 Myr following the cessation of compression. During 619 this period the amplitude of the folds decreased by less than 10 percent. At this stage, crustal faults 620 may accelerate the gravity collapse of folds, creating inverted basins. lithosphere leads to a symmetrical pattern of down-warped areas (synclines) flanked 627 by highs (anticlines) of similar amplitude and wavelength, like for instance the Ferghana basin in

  632et al., 1999). Both types of compressional basins show linear nature alignment of parallel 633 depocentres and flanking highs. An important difference is the presence of parallel trending 634 depocentres for lithospheric folding, whereas foreland flexural basins have a single depositional 635 system. In addition, basins developed on folded lithosphere have a static location of the axis of their 636 contrast to foreland basins, where the axis of depocentres can migrate with time 637 (e.g.Zoetemeijer et al., 1993).638In both cases, the integrated strength of the lithosphere (Watts and[START_REF] Robin | Vertical movements of the Paris Basin (Triassic -Pleistocene): from 3D stratigraphic database to 1443 numerical models[END_REF] Tesauro et al., 639 2007) defines the characteristic width of the basins. For folding, basin width roughly corresponds to 640 one half of the fold wavelength (), which equals 5 to 10 times the thickness of the strong core of 641 the lithospheric plate, h (h = 10 to 100 km, i.e.  = 50 -1000 km). In the case of crust-mantle 642 decoupling, which occurs in relatively young plates, two dominant folding wavelengths develop643(typically 250 -400 km and 50-100 km). Thus two basin populations may be observed, one 644 imbricated within the other. In the case of flexural foreland deformation, the width of the basin is 645 controlled by the flexural parameter and is thus of the same order as the largest fold wavelength.

655Figure 11a

 11a Figure 11a displays characteristic basin subsidence patterns predicted by folding of 150 and 300

  687earlier, erosion reduces the contribution of gravity-dependent terms and accelerates local 688 , therefore, has an important feed-back with the geometry of the

700

  followed by a forward prediction in time. Most of these software packages are based on sequence 701 stratigraphic concepts(Wilgus et al., 1988). Assumptions for initial subsidence are either ad-hoc or 702 related to stretching (e.g.Lawrence et al., 1990) or foreland flexure (Garcia-Castellanos et al., 703 1997), but to date do not include lithospheric folding as a basin formation mechanism. 704 705 Marine versus continental deposits and closed basins versus capture of drainage 706 Erosion of uplifted areas and sedimentation in the depressions created by folded lithosphere is a 707 self reinforcing process promoting continuing uplift of the highs and subsidence in the depocentres. 708 Thus, such basins will remain closed during much of their evolution. Folded basins will hence be 709 predominantly characterized by continental deposits with only minor deposition of marine sediments 710 during marine incursions or during basin capture. This seems to be the case for folded basins in 711 Iberia (e.g. Duero Basin) and Central Asia (Ferghana, Tarim Basin, Lake Issyk-Kul) (Cobbold et al., 712 1993). As has been pointed out for flexural foreland basins (Garcia-Castellanos et al., 2003), the 713 capture time for opening basins decreases with increasing flexural rigidity. The flexural upwarp to removed by erosion will be almost instantaneous for very weak 715 lithosphere, but requires up to 50-100 Myr for high lithospheric rigidities (Garcia-Castellanos et al., 716 2003). Thus, basins created by folding of cratonic lithosphere probably have a shorter capture time 717 than ones formed in younger lithosphere. controlled by the interplay of lithospheric stresses and inherited strength of 723 the lithosphere. The thermal regime controls the rheological profile and differs between folded 724 basins developed in young lithosphere and basins in cratonic lithosphere. The latter are associated 725 with much lower thermal gradients than basins developed on young continental lithosphere. As for

  , modifying the surface 730 heat flow in the basin centre. As pointed out by Lavier & Steckler (1997) and Ziegler et al. (1998), 731 the effect of sediment fill is to weaken the lithosphere. The low thermal conductivities of the 732 sediments lead to high temperatures in the upper lithosphere and consequently low local yield 733 strength. During basin subsidence sediments deposited in the central parts of the basin might be 734 exposed to temperature windows corresponding to hydrocarbon generation. At the same time, 735 sediments pre-dating the folding may have undergone extra burial due to syn-folding sedimentation.

738

  During basin capture overall cooling takes place. This effect is illustrated by characteristic thermal 739 evolution for different thermo-tectonic ages of 150 and 300 Ma of the folded continental lithosphere, production in the sediments filling in the synclinal depression, shown 741 in Figure 13. Subsidence induced by folding is calculated for the centre of the basin. Basins 742 developing on folded lithosphere are characterized by their relatively low heat flow at the onset of 743 folding, followed by a steady increase with time, doubling the heat flow over a time interval of the 744 order of 5 Myr following the cessation of shortening. This increase is primarily due to the 745 contribution of radioactive heat production in the very substantial pile of sediments accumulating in 746 a relatively short time interval. The patterns of heat flow are similar for different thermo-mechanical 747 ages of the lithosphere, with the oldest lithosphere having the lower heat flow. Older ages enhance 748 the accommodation space and the contribution of the sediments to heat flow but are compensated 749 by a larger decrease in heat flow with aging lithosphere. Another important factor is the mode of 750folding. Crustal scale folding, which is characteristic for Lake Issyk-Kul, will be associated with 751 shorter wavelengths, shallower basin depths and thinner sequences of heat producing sediments 752 than for mantle lithospheric folding. This effect might explain the relatively low heat flow observed in 753 Lake Issyk-Kul(Vermeesch et al., 2004).

760

  folding and faulting can develop simultaneously and that pre-existing crustal scale faults do not 761 prevent, but instead promote the development of folding instabilities. The instabilities, in turn, can 762 initiate the formation of new faults at the inflection points of the folds. Although the usual expectation 763 is that faulting prevents folding, this view neglects the role of the gravity and friction on faults in the gravity and friction. After some sliding (uplift) on a fault, the potential gravitational 767 energy working against friction generated by horizontal shortening forces becomes too high. As a 768 result, the fault locks and the rock mass transmits horizontal stresses as a continuous medium. As 769 the compression continues, one of the folds finally starts to grow faster than others, resulting in a 770 loss of the periodicity and the formation of a mega-fold (see also the discussion in the context of the 771 Ferghana and Tadjik basins of Central Asia).

790

  spacing of faults depend on the thermo-mechanical age of the lithosphere. Once the development 791 of folding begins in a basin, the intensity and depth penetration of faulting will accelerate. The mode 792

797

  contrast with the later phase of basin capture, which is in general characterized by erosion and 798 narrowing of the basin. Due to erosional unloading, stress relaxation will occur, manifested in 799 upward propagating faults. 800 801 Interaction of folding with other tectonic processes 802 Folding, in addition to being a basin forming mechanism, frequently interacts with other tectonic 803 processes. As pointed out by Cloetingh (1988), intraplate compression can modify pre-existing 804 basins. As discussed above, the Pannonian Basin of Central Europe created by Miocene back-arc 805 extension appears to be an example of such a configuration, characterized by a-typical a-periodic 806 folding. Other examples are cratonic sag basins where intraplate compression

815

  Basin, followed by foreland flexure and overprinted by Late Miocene compression(Tarapoanca et 816 al., 2003). As pointed out above(Ziegler & Dèzes, 2007), the Northwestern European foreland 817 indicates folding, overprinting, rifting and foreland flexure. Quantification of the topographies created foreland flexure and correcting for them is essential to reconstruct the shape of the 819 additional accommodation shape created by subsequent lithospheric folding(Bourgeois et al.,

  827 intermediate age lithosphere. Emplacement of hot upper mantle material will raise temperatures in 828 the lithosphere and increase heat flow. Due to the lag time in the propagation of heat, the effect at 829 the surface might become manifest only after several tens of Myr. Plume emplacement might also 830 weaken the lithosphere, making it more prone to compressional fault reactivation after cessation of 831 the folding. 832 Interplay of lithosphere folding and plume impingement on the continental lithosphere occurred 833 probably almost simultaneously in late Neogene times in the Northwestern European foreland. Two 834 examples are the Eifel and Massif Central areas of the Alpine foreland of NW Europe (Cloetingh & 835 van Wees, 2005; Ziegler & Dèzes, 2007). Both areas are sites of main Late Neogene volcanic 836 activity in the European Cenozoic rift system (ECRIS). Seismic tomography (Ritter et al., 2001) 837 shows finger shaped baby plumes with a characteristic spatial diameter of 100 km, extending 838 downward to 400 km. Ziegler & Dèzes (2007) propose that plume activity occurs simultaneously 839 with recent compressional deformation in the Massif Central area (Guillou-Frottier et al., 2007) and 840 the Ardennes/Eifel area. Geomorphological studies constrain the recent uplift of the Ardennes and 841 the Eifel area. An order in magnitude difference occurs between the uplift of the Eifel area, underlain 842 by a plume and the adjacent Ardennes area where evidence for a plume is lacking. The patterns of843uplift appear to be radial, superimposed by a linear NE-SW trend perpendicular to the main axis of 844 plumes primarily develop in the anticlines of lithospheric folds(Burov &

849

  Thermo-mechanical modelling illustrates the relative effectiveness of amplification of lithosphere 850 deformation and topographic effects induced by plumes though folding and vice-versa. Burov & 851 Cloetingh (2009) examined the response times and time-lags involved and whether these baby 852 plumes were more efficient in localizing deformation than large plumes. Plume-affected folding 853 appears to accelerate surface uplift, whereas folding goes into saturation and stagnates when 854 plastic hinges form. For the discrimination of plumes and folding it is critical to access constraints on 855 the presence or absence of radial versus linear symmetry, heat flow anomalies, gravity and geoid 856 data. Plume activity facilitates folding, by dramatically lowering the stress levels required (Burov & 857 Cloetingh, 2009). Plume impact also reduces the fold wavelength and localizes folding above the 858 plume impact area. A general outcome of the modelling and observations is that lithospheric folding 859 as a mechanism for producing thermal perturbations in the lithosphere/upper mantle system is a through lithospheric folding have characteristic features including 865 vertical motions, basin architecture, thermal regime and fault activity, distinguishing them from other 866 basin types

867

  differences and similarities between basins developed on folded lithosphere (FLB), foreland basins 868 (FB), intracratonic basins (ICB), extensional basins (EB), and pull-apart basins (PAB).

892

  with dimensions that can vary from 50 km to 600 km, depending on lithospheric age and shortening 893 rate. These basins can accumulate thick sedimentary sequences up to the order of 20 km. The time 894 scales associated with this process of basin formation are very short, typically a few Myr. The 895 fast that erosion and sediment supply at the basin formation stage typically lags 896 behind, leading to the development of starved basins, followed by shallowing upward sequences 897 after stress relaxation. Subsidence patterns are characteristically convex upward with time. A 898 noticeable feature is the development of significant topography of the order of several km flanking 899 the synclinal depression. No initial heating is predicted, with an increase in heat flow with time 900 following the basin formation. Significant brittle deformation is expected to occur in the folded 901 basement, accelerating during the folding phase.

917

  basin fill with FLB. An important difference is their evolution in time, characteristically associated 918 with very long time spans (typically of the order of 500-800 Myr) with long phases of low subsidence 919 rate punctuated by phases of subsidence acceleration. As pointed out above, this is why ICB are 920 probably affected by short-term phases of enhanced intraplate compression superimposed on other 921 longer time scales. The large dimensions of these basins point to 922 formation in high rigidity lithosphere, that freezes in basin deformation over long time scales. The 923 long time scales involved also imply that the highs induced by compressional upwarping will be 924 commonly eroded. A major difference between ICB and FLB is that current hypotheses for ICB 925 formation are casted only in terms of the tectonic processes operation at their earlier stage such as 926 phase changes in the lithosphere. It appears, therefore, that lithospheric stress fields and the 927 lithosphere rheology (inferred from gravity and flexure) are far more diagnostic than the lithosphere-928 asthenosphere boundary, which is only giving information on the thermal state of the lithosphere.

942

  can lead to feed back with lower crustal flow and faulting in the extending lithosphere(Burov & 943[START_REF] Burov | Erosion and rift dynamics: new thermo-mechanical aspects of post-1111 rift evolution of extensional basins[END_REF]. During the post-rift phase, flexural widening of the basin commonly occurs. A 944 characteristic feature of EB predicted by simple stretching models (McKenzie, 1978) is the 945 succession of rapid initial subsidence in the syn-rift phase and exponentially decaying subsidence in 946 their post-rift phase with a typical thermally controlled time constants of the order of 70 Myr. Feed 947 shoulder erosion and lower crustal flow (Burov & Cloetingh, 1997; Burov & 948 Poliakov, 2001), have pronounced effects on subsidence in the final phases of rifting and in the 949 early post-rift phase. Both can be influenced by changes in the mechanical properties of the 950 lithosphere and its strengthening upon rifting. Multiple rifting phases is another common feature of 951 EB. The duration of the syn-rift phase can vary from quasi-instantaneous (shorter than 20 Myr) to 952 several hundreds of Myr

  Fig. 16 shows such a setting, which is based on the 958 benchmark model presented in Van Wees et al. (2009), As a consequence of the loss of radiogenic 959 heat production in the crust at the end of the post-rift stage values, heat flow evolves to values 960 considerably lower than the initial ones. The EB heat flow pattern is, therefore, the reverse of that in 961 FB where syntectonic heat flow is first lowered due to sediment infilling (Kombrink et al., 2008), but 962 eventually increases due to thermal relaxation and additional radiogenic heat production in the 963 sediments.

  990convex up for FLB and concave for EB and PAB. The heat flow histories differ also. EB and PAB 991 are characterized by a phase of initial heating, whereas FLB and FB are characterized by an 992 absence of initial heating. A common feature of FLB and PAB formation is their capability to cause 993 dramatic differential vertical motions in a short time interval, predicting shallowing upward 994 sequences. FLB are to some extent unique in terms of the magnitude of the uplifted topography 995 surrounding the down warped depocentres. Similar magnitudes are only reached at rift shoulders of 996 basins developing in intracratonic lithosphere(Van der Beek et al., 1995). FLB can be associated 997 with localized brittle deformation and pop-up structures in the upper crust and widespread brittle 998 deformation at deeper crustal and upper mantle levels. by lithospheric folding are characterized by a number of features in 1004 their basin architecture, subsidence history, thermal evolution and faulting patterns making them 1005 distinctly different from other basin systems. Of particular importance are the relatively short 1006 temporal scales involved in their evolution, whereas their spatial scales vary from tens to several 1007 hundreds of km, depending on the rheological stratification of the lithosphere and its thermo-1008 mechanical age. Subsidence patterns are characterized by an acceleration of subsidence in the 1009 depocentres during compressional basin formation, occurring simultaneously with accelerated uplift 1010 of and erosion from flanking basin highs. In contrast with asymmetrical foreland basins, basins on 1011 folded lithosphere are symmetrical with equal dimensions for linearly trending topographic highs and 1012 lows. 1013 Basins on folded lithosphere have a strong tendency to remain closed systems during a major part 1014 of their evolution, with capture times decreasing with thermo-mechanical age from several tens of 1015 Myr to less than 10 Myr. The thermal history of these basins is not characterized by an initial 1016 thermal perturbation, unless folding occurs in interaction with plume activity. In the absence of an 1017 interplay with lithospheric scale thermal perturbations, their thermal evolution will be primarily 1018 controlled by the burial history of the sediments deposited in the accommodation space created by 1019 the folding process. 1020 Intrabasinal faulting and the formation of pop-up structures and inverted basins at upper crustal 1021 levels are important in the structural evolution of these basin systems. The intensity of the brittle 1022 deformation peaks with ongoing shortening of the lithosphere during folding but faulting activity is 1023 dramatically reduced after the cessation of folding. At this stage, with an equilibrium between 1024 sedimentation and erosion, minor faulting occurs as a result of stress relaxation due to erosion of 1025 During capture and overall uplift of basins on folded lithosphere, their accommodation 1026 space is further reduced by unflexing and associated stress release through faulting. initiated during S. Cloetingh's visit to Paris VI as Visiting Professor and finalized 1031 during a subsequent Visiting Professorship at IFP. Comments and stimulating suggestions by Fred 1032 Beekman, Jan-Diederik van Wees, Francois Roure, Seth Stein, Jeroen Smit and Didier Granjeon 1033 are gratefully acknowledged. Thomas Kruijer and Elishevah van Kooten provided valuable technical 1034 assistance. We benefitted from constructive criticism from Dr. Mary Ford and two anonymous ., SOKOUTIS, D., BADA, G., CLOETINGH, S., & HORVATH, F. (2010) Modelling of recent 1221 deformation of the Pannoanian lithosphere: lithospheric folding and tectonic topography. Tectonophysics, 484, .S., MOSAR, J., BRUNET, M.-F. & KANGARLI, T. (2009) Subsidence and uplift mechanisms within 1225 the South Caspian Basin: insights from the onshore and offshore Azerbaijan region. Geol. Soc. London Spec. 1226 Pub., 312, 219-240.

  . (1976) Slope modelling with changing boundary conditions -effects of climate and lithology, Z. 1260 Geomorphol. Suppl., 25, 72-88. 1261 1262 GRANJEON, D. (2009) 3D stratigraphic modelling of sedimentary basins, from source to sink. AAPG Annual 1263 Convention and Exhibition, June 7-10, 2009, Denver Colorado. 1264 1265 GRENERCZY, Y., SELLA, G.F., STEIN, S. & KENYERES, A. (2005) Tectonic implications of the GPS velocity 1266 field in the northern Adriatic region. Geophys. Res. Lett., 32, L16311. 1267 1268 GUEST, B., GUEST, A. & AXEN, G. (2007). Late Tertiary tectonic evolution of northern Iran: a case for simple 1269 crustal folding. Global. Planet. Change, 58, 435-453. 1270 1271 GUILLOCHEAU et al. (2000) Meso-Cenozoic geodynamic evolution of the Paris Basin: 3D stratigraphic 1272 constraints. Geodin. Acta, 13, 189-246. 1273 1274 GUILLOU-FROTTIER, L., BUROV, E., NEHLIG, P. & WYNS, R. (2007) Deciphering plume-lithosphere 1275 interactions beneath Europe from topographic signatures. Global Planet. Change, 58, 119-140. 1276 1277 HAFKENSCHEID, E., WORTEL, M.J.R. & SPAKMAN, W. (2006) Subduction history of the Tethyan region 1278 derived from seismic tomography and tectonic reconstructions. J. Geophys. Res., 111, B08401. 1279 1280 HORVÁTH, F. (1995) Phases of compression during the evolution of the Pannonian Basin and its bearing on 1281 hydrocarbon exploration. Mar. Petrol. Geol., 12, 837-844. 1282 1283 HORVÁTH, F. & CLOETINGH, S. (1996) Stress-induced late-stage subsidence anomalies in the Pannonian 1284 basin. Tectonophysics, 266, 287-300. 1285 1286 HORVÁTH, F., BADA, G., SZAFIÁN, P., TARI, G., ÁDÁM, A. & CLOETINGH, S. (2006) Formation and 1287 deformation of the Pannonian basin: constraints from observational data. Geol. Soc. London Mem., 32, 191-.S. & BEAUMONT, C.B. (2002) Asymmetric lithospheric extension: the role of frictional plastic 1291 strain softening inferred from numerical experiments. Geology, 30, 211-214. 1292 1293 HUNT, G., MUHLHAUS, H., HOBBS, B., & ORD, A. (1996) Localized folding of viscoelastic layers. Geol. 1294 Rundsch., 85, 58-64. . & MAXSON, J. (2001) Lithospheric buckling of the Laramide foreland during Late Cretaceous and .F. & PODLADCHIKOV, Y. (1994) Diapiric ascent of magmas through power law crust and and a functional, respectively. In the Lagrangian method, incremental 1616 displacements are added to the grid coordinates allowing the mesh to move and deform with the solution is obtained in local coordinates, which are then updated in a large strain 1619 mode.1620Solution of (A1) provides velocities at mesh points used for computation of element strains and of 1621 heat advection uT. These strains are used in (A2) to calculate element stresses and equivalent 1622 forces as input for the computation of the velocities for the next time step. Due to the explicit 1623 approach, there are no convergence issues, which is rather a common problem of implicit methods 1624 in case of non-linear rheologies. The algorithm automatically checks and adopts the internal time 1625 step using 0.1 -0.5 of Courant's criterion for propagation of information, which warrants a stable

;

  Kirby & Kronenberg, 1635[START_REF] Quinlan | Models of subsidence mechanisms in intracratonic basins and their applicability to North 1428 American examples[END_REF] Kohlstedt et al., 1995).

  , which 1639 corresponds to Mohr-Coulomb material with a friction angle  = 30° and cohesion C 0  < 20 MPa 1640 (e.g., Gerbault et al., 1998): where  n is normal stress  n =  I +  II dev sin,  I = P is the effective pressure,  II dev is the 1642 second invariant of deviatoric stress, or effective shear stress. The condition of transition to brittle 1643 deformation (function of rupture f) reads as: f =  II dev + P sin -C 0 cos = 0 and f/t= 0. In terms of activation enthalpy, R is Boltzman's gas constant, and T is temperature in K°, 1  and 3  are the 1658 principal stresses. The effective viscosity  eff for this law is defined as:

1665Figure 1 .Figure 2 .

 12 FIGURE CAPTIONS 1675

Figure 3 .

 3 Figure 3. Relationship between wavelengths of lithospheric folding and thermo-tectonic age of

Figure 4 .

 4 Figure 4. (a) Map showing major tectonic features of Tien Shan area of Central Asia , with location

1702after

  [START_REF] Cloetingh | Lithosphere folding: primary response to compression? 1165 (from Central Asia to Paris Basin)[END_REF]. (c) Measured and calculated Bouguer gravity anomalies for lithosphere 1703 under horizontal compression along profile F across the Ferghana basin with corresponding 1704 topography profile (d) (after Burov & Molnar, 1998). Predicted topography and Moho geometry for 1705 Airy and best fitting folding scenario. During compression the amplitude of plate deflection changed 1706

Figure 5 .Figure 6 .Figure 7 .Figure 8 .

 5678 Figure 5. First order features of basin geometry of folded lithosphere of Iberia. (a) Topography

Figure 9 .Figure 10 .Figure 11 .Figure 12 .Figure 13 .

 910111213 Figure 9. Characteristic patterns of continental lithosphere deformation induced by folding of the

Figure 14 .Figure 15 .Figure 16 .

 141516 Figure 14. Results of analogue tectonic experiments for lithospheric folding (after Sokoutis et al.,

. 375 Figure 8 presents a cross section of northern Iran and the South Caspian Basin (Guest et al.,

	376
	377	2007). As shown, wavelengths for compressional folding of the lithosphere are typically in the range
	378	of 400 km. Although the actual basin formation mechanism for the Caspian Basin is not well
	379	resolved (e.g. Artyushkov, 2007), late Neogene folding has been affecting a lithosphere probably
	380	thermally reset by middle Late Jurassic marginal basin formation (Guest et al., 2007) with a thermo-
	381	mechanical age at the onset of collision of 130-150 Ma. Satelite data demonstrate an exceptionally
		large gravity anomaly over the basin

400 of Zagros strike slip and extrusion of central Iran, coarse Zagros molasse deposition, Dead Sea 401 transform reorganisation, Red Sea oceanic spreading

	385	Basin analysis has shown that basement inversion has occurred 20 Ma ago, simultaneously with
		the slowing of Arabian/Eurasian plate convergence and the onset of accumulation of Neogene
		F o
		r
		R e v i e w to a flexural support by the South Caspian basement (Axen et al., 2001). The simultaneous reversal
		O n
		l y
	383	beginning at
	384	about 10-15 Ma.
		16

  argue in favour of compressional deformation of the South Caspian Basin/Alborz

		F o
		r
		R e v i e w
		O n
		l y
	409
	410	mountains. This interpretation was largely based on available industrial seismic reflection data and
		17

2 illustrate

  . Red lines are 1707 Bouguer anomaly predicted by folding model and Moho depression, respectively. Solid grey line is 1708 Airy isostasy. Solid black line is Airy Moho. Triangles indicate observed Bouguer anomaly. Green is 1709 topography above sea level. FB: Ferghana Basin.

  .  33 is density; Q, n, A are material-dependent parameters of ductile flow laws (Kirby & Kronenberg, 34 1987; Kohlstedt et al., 1995). Other parameters from Turcotte & Schubert (2002).

	35
	36
	37
	38	Parameter	Value
	39
		F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r F o r
		R R R R R R R R R R R R R R R R R
		e v e v e v e v e v e v e v e v e v e v e v e v e v e v e v e v e v
		i e i e i e i e i e i e i e i e i e i e i e i e i e i e i e i e i e
		w w w w w w w w w w w w w w w w w
		O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l O n l
			y y y y y y y y y y y y y y y y y

[START_REF] Burov | Gravity anomalies over the Ferghana Valley (Central Asia) and 1120 intracontinental deformation[END_REF]. Three-dimensional fold structures in the Tibetan Plateau have been inferred from

collected in the Cantabrian mountains, the Spanish Central System, and the Toledo Mountains (see

basin fill[START_REF] Sacchi | Role of unconformity bounded units in the stratigraphy of 1457 the continental record: a casestudy from the Late Miocene of the western Pannonian basin, Hungary[END_REF], see Fig.7b) and for its reservoir structures, both crucial in

assessment of its hydrocarbon potential(Horváth, 1995). Analogue tectonic experiments by

Dombradi et al. (2010) further examine accumulation and transmission of stresses in the hot and

Basin Research

clastics in foreland basins (e.g.Fakhari et al., 2008). By 10 Ma, contraction occurred by

underplating of the Arabian crustal units beneath the Iranian plate. Estimated shortening rates from

Pannonian Basin appear to be primarily controlled by the pre-existing geometry of the sediment-

Page 18 of 91 Basin Research

behaviour of the large-scale fault-and-fold systems. As also observed in analogue experiments

[START_REF] Burg | Shortening of analogue models of the continental lithosphere: 1102 New hypothesis for the formation of the Tibetan plateau[END_REF], this continuous behaviour of faulted lithosphere can be explained by fault

Page 30 of 91 Basin Research

[START_REF] Cloetingh | Tectonic Models for the Evolution of Sedimentary Basins[END_REF]. According to presently available data, the plumes in this segment of the

Page 34 of 91 Basin Research

absent in continental extensional settings as a consequence of removal of heat producing elements

[START_REF] Pitman | The effect of sea-level change on the shelf edge and slope of 1421 passive margin. In: The shelf break: critical interface on continental margins[END_REF]. These basins are intensively faulted, also due to their strike-slip

origin, whereas heat flow decays on very short time scales of the order of a few Myr. Although this

BIYO, J.J. & DE URREIZTIETA, M. (1993) Sedimentary basins and crustal thickening. Sediment. Geol., 86,

Shan. C.R. Acad. Sci., Ser. IIa: Sci. Terre Planets, 234,[49][50][51][52][53][54][55][56][57]

and deflection data. J.Geophys. Res., 93,[7747][7748][7749][7750][7751][7752][7753][7754][7755][7756][7757][7758][7759][7760][7761][7762][7763][7764][7765][7766]

lithosphere under the western outer Carpathians.Tectonics, 18,[843][844][845][846][847][848][849][850][851][852][853][854][855][856][857][858][859][860][861]

material. This enables solution of large-strain problems locally using a small-strain formulation. On

Page 58 of 91 Basin Research

Vogelsberg anticline (NVA), Sologne-Franconian Basin Syncline (SFS) and the Burgundy-Swabian

Jura Anticline (BSJA). LBR: London Brabant Ridge with thick crust. BF: Black Forest; FB

location Fig.6a). Thin line: actual topography surface. Bold line: elevation of top basement as

predicted in the absence of rifting. Fold signatures are similar with a wavelength  of 270 km and

amplitude A of 1500 m in the Franconian Basin (Cross section G-G') and in the upper Rhine Graben

area (Cross section H-H'). In the Paris basin, development of folds with assumed same wavelength

and amplitude (stippled bold line in cross section C-C'), implies subsidence of Sologne and uplift of

Normandy.

motions are typically of the order of 250 km. (b) Interpreted seismic section in the central part of the

Pannonian Basin[START_REF] Sacchi | Role of unconformity bounded units in the stratigraphy of 1457 the continental record: a casestudy from the Late Miocene of the western Pannonian basin, Hungary[END_REF] showing evidence for late Neogene crustal scale folding at

scales of 20-40 km. Note syn-rift grabens formed during the Miocene back-arc extension of the

Page 64 of 91 Basin Research

side (b) cross-section located on the top view. Arrows indicate the direction of the moving wall. Note

the similarity to Central Iberia (Fig.5). (Bottom): Geometries predicted by analogue experiment,

shortening a lithosphere with two contrasting blocks, bounded by suture. A deep synclinal

depocentre develops over the suture zone, flanked by an anticline of smaller amplitude.Left-side (c)

final top view of model after 26% of bulk shortening. Right side (d) cross-section located on the top

Page 68 of 91 Basin Research

APPENDIX A: MODEL DESCRIPTION 1600 1601 A1. Equations. The Flamar-Para(o)voz is a "2.5 D" FLAC-like code [START_REF] Cundall | Numerical experiments on localization in frictional materials[END_REF]. It has a mixed 1602 finite-difference/finite element numerical scheme, in which the coordinate frame is Cartesian 2D, but 1610

Here u,    , g, k are the respective terms for the displacement, stress, acceleration due to body 1611 forces and thermal conductivity. The triangular brackets in (A1) specify conditional use of the related 1612 term (in quasi-static mode inertial terms are dumped using inertial mass scaling [START_REF] Cundall | Numerical experiments on localization in frictional materials[END_REF]).

1613

The terms t, , C p, T and H r designate respectively time, density, specific heat, temperature and 1614 internal heat production. The terms D/Dt, D/Dt, F are a time derivative, an objective (Jaumann) principal stresses, the equivalent of the yield criterion (5) reads as:

1647

The elastic behavior is described by the linear Hooke's law:

1648

where  and G are Lame's constants. Repeating indexes mean summation and  is Kronecker's

1652

Mantle convection and a part of lithospheric deformation is controlled by thermally activated creep 1653 [START_REF] Ranalli | Rheology of the Earth[END_REF]Kirby & Kronenberg, 1987). Within deep lithosphere and underlying mantle regions, 1654 the creeping flow is non-Newtonian since the effective viscosity can vary within 10 orders of 1655 magnitude as a function of differential stress:

Where