Lateral Heterogeneity in Compressional Mountain Belt Settings

Convergent orogens are typically linear with laterally continuous, orogen-parallel folds and thrusts. Over the years, geoscience research has revealed evidence for important orthogonal/cross structures as well as lateral heterogeneity in deformation style, igneous activity, metamorphic grade, geomorphology, and seismic activity. To assess the occurrence, causal mechanisms, and implications of these lateral heterogeneities, a selection of convergent orogens, with different tectonic settings and history are reviewed. The Appalachians, the North American Cordillera, the Alps, the Himalayas, the Zagros, the Andes, and several other belts all exhibit a degree of lateral heterogeneity. Major factors driving the lateral heterogeneity and/or cross structures include the pre-existing deformational history of the cratonic blocks involved, lateral change in lithology of crustal rocks, variations in crustal/lithospheric rheologic properties, or changes in plate kinematics. The Appalachian orogenic front mimics the Iapetan rift margin. Pre-existing basement structures have control on preand syn-orogenic sedimentation, which subsequently impacts how an orogenic wedge evolves. A thicker sedimentary column generally evolves into a salient (as opposed to a recess), which is further enhanced by the presence of weak horizons as seen in the Zagros and the Cordillera. Lateral variation in sedimentary facies also creates changes in thrust-ramp geometry. During orogenic contraction, inherited basement structures can be preferentially reactivated based on their orientation. Several cross faults in the Himalayas spatially coincide with orogen-perpendicular, lower plate, basement structures. In a similar way, oceanic subducting plate physiography can also influence deformation in the overriding plate. Along-strike variations in subduction dynamics have been reflected in the Andean deformation. Orogenic extension in the Alps has been accompanied by a system of orogen-parallel strike slip faults and extensional cross faults. It is evident that lateral heterogeneities can form crucial control on the evolution of orogenic belts and can influence seismic rupture patterns, resource occurrence, and landslide-related natural hazards.

• Lateral heterogeneity, expressed as sharp lateral variations, has been observed in most orogens 9 and often coincides with cross structures. 10 • Major controls are basement structures, lateral changes in sedimentary facies, varying crustal 11 rheology, and changing plate dynamics. 12 • Lateral heterogeneities impact orogenic evolution and have geomorphic, seismic, and 13 economic implications. grade of the Barrovian metamorphic series) more strongly than in the recess (Tull & Holm, 2005). 194 Towards the adjoining Virginia recess, a section of the Pine Mountain belt, bounded by the left-195 lateral Jacksboro tear fault and the right lateral Russel Fork tear fault (Figure 3a), demonstrates a 196 double ramp geometry in the middle section, while a single ramp geometry around the edges 197 (Mitra, 1988). Presence of the orogen-parallel, large-scaled (about 35 km wide, (Brewer, 2004)) 198 Birmingham graben beneath the southern Alleghenian thrust belt also had major controls on the 199 foreland sedimentation and subsequent propagation of the thrust sheet (Thomas, 2007). Transverse 200 zones in the southern Appalachians have been interpreted as inherited transform fault related to 201 the Iapeton rifting along the continental margin of Laurentia (e.g.; Figure 4) (Rankin, 1976;202 Thomas, 1977202 Thomas, , 2014. 203

Cross-structures in the Appalachians of Pennsylvania and West Virginia (central 204
Appalachians) such as the Parsons, Petersburg, and Tyrone-Mount Union lineaments ( Figure 2) 205 also have similar effects on the orogen-parallel structures as is seen in the southern Appalachians 206 (Wheeler, 1980;Southworth, 1986). In the northern Appalachians, however, genesis of cross-207 structures presents much more diversity. The Serpentine Lake cross-strike discontinuity (CSD), 208 the Bonne Bay CSD, the Canada Bay CSD, and Belle Isle CSD in Newfoundland mark the 209 thrusts and associated structures are west verging, while towards the south the vergence reverses 220 (Williams & Cawood, 1986). In southern Newfoundland and the Cape Breton islands, the Silurian 221 tectonic wedging of the Avalon craton into the Laurentian craton has resulted in a very narrow 222 Appalachian central mobile belt characterized by stronger syncollisional deformation, high-grade 223 Barrovian metamorphism ('Kyanite and Sillimanite grade, 700˚-750˚C at 6-10 Kbar (Plint & 224 Jamieson, 1989), 5 kbar elsewhere'), and a greater volume of Silurian magmatism than elsewhere 225 (Lin et al., 1994). This intensely deformed and metamorphosed narrow band is bound by the NW- NE-striking, 300 km long, transpressive, dextral fault system (West Jr & Hubbard, 1997;Hubbard, 237 1999) has recently been interpreted to be a result of subduction of an oceanic ridge and related 238 transform fault (Kuiper, 2016;Kuiper & Wakabayashi, 2018). Around the orogenic front in 239 Quebec and northern Vermont, cross structures were formed due to reactivation of earlier-formed 240 transverse, normal faults with an oblique thrust-sense or served as an oblique ramp during the 241 westward thrusting of the Lower to Middle Ordovician shelf-carbonates (Séjourné & Malo, 2007). 242 243

Proposed Factors Controlling Lateral Heterogeneities 244
Lateral heterogeneity along the Appalachians has largely been controlled by the irregularities 245 along pre-collisional continental margins and by inherited basement structures. The eastern 246 continental margin of Laurentia went through a series of rifting events including the Iapetan rifting, 247 all of which acted in concert to develop a series of embayments (concave oceanward) and 248 promontories (convex oceanward) separated by transform faults (Rankin, 1976;Thomas, 1977Thomas, , 249 2014. During subsequent orogenic events (Taconic, Acadian, and Alleghenian orogenies), 250 embayments evolved into salients (convex towards the foreland), promontories evolved into 251 recesses (concave towards the foreland), and the transform fault boundaries evolved into transverse 252 zones ( Figure 4). As Tull and Holm (2005) noted, foreland sedimentary sections along salients are 253 much thicker than along recesses, which has a great control on the geometry and deformation style 254 of the evolving thrust sheets (discussed in detail in the North American Cordillera section of this 255 article). Irregularity along continental margins also has another important control, which results in 256 a sharp lateral heterogeneity along an orogen. A collisional event between two promontories (of 257 two different landmasses) results in a narrower but stronger orogenic deformation and a higher 258 grade of metamorphism that a promontory-embayment collision (Lin et al., 1994). The evolution 259 of rift-related transform faults into transverse zones also had a profound impact on the evolution 260 of the Appalachians mountains by creating barriers or zones that truncated or offset major thrust 261 faults. Along the range front, transverse zones have served as nucleation sites for cross structures 262 in the form of lateral ramps or tear faults. 263 264 3 Cordillera 265

Tectonic Setting and Lateral Heterogeneities 266
The North American Cordilleran orogenic system records the Jurassic (and possibly earlier) to 267 Paleogene history of terrane accretion and the eastward convergence of the Farallon and Kula 268 plates beneath the western margin of Laurentia (North American continent) (Oldow et al., 1989; 269 involved Laramide orogenesis is one of the remarkable changes. Initiation of the Laramide system 291 uplifts has generally been correlated to the subduction of a flat-slab segment (e.g.; Coney & 292 Reynolds, 1977;Dickinson & Snyder, 1978;Saleeby, 2003;Chapin, 2012). North of the Canadian 293 border, the Laramide-style deformation is absent. Orogenic collapse of the Cordilleran system, 294 focused on the Sevier belt, initiated during Middle to Late Eocene due to the change in relative 295 plate kinematics and has resulted in the active Basin and Range province of extensional 296 deformation (e.g.; Constenius, 1996;Yonkee & Weil, 2015). The contractional phase of 297 deformation resulted in spectacular range-parallel fold and thrust belts (Yonkee & Weil, 2015), 298 yet there are important lateral heterogeneities and a number of cross structures along the Cordillera. 299 Basement provinces and their suture zones and major faults along the western margin of 300 the North American continent include the Snowbird Tectonic Zone (Ross et al., 1991;Hope & Eaton, 2002), the Archean Hearne Province (Hope & Eaton, 2002), the Archean Medicine Hat 302 Block (Ross et al., 1991;Lemieux et al., 2000); the Archean Wyoming Province (Wooden & 303 Mueller, 1988;Mogk et al., 1992), and the Paleoproterozoic Yavapai and Mazatzal Provinces 304 (Foster et al., 2006;Whitmeyer & Karlstrom, 2007), from north to the south (Figure 5b). The 305 Trans-Hudson Orogen and the Superior Craton lie towards the east, around the central region of 306 the North America. Towards the west, the basement block include the Archean Grouse Creek 307 Block and the Selway Terrane (Foster et al., 2006). 308 Transverse crustal boundaries between these provinces and the other 309 Archean/Paleoproterozoic transverse basement structures are one of the primary causes of lateral 310 heterogeneity in the Cordillera due to their influence in the subsequent sedimentary and 311 deformational history (e.g.; Paulsen & Marshak, 1999;Sears & Hendrix, 2004;McMechan, 2012). 312 Significant crustal-scale, transverse boundaries around the southeastern Canadian Cordilleran 313 front (Alberta) include the Thorsby Low, a crustal suture within the Snowbird tectonic zone (Ross 314 et al., 1991;Hope & Eaton, 2002), the Red Deer zone, the northern boundary of the Hearne 315 Province (Hope & Eaton, 2002), and the Vulcan Low, a crustal suture between the Medicine Hat 316 Block and Hearne Province ( Figure 5b) (Hoffman, 1988;Ross et al., 1991). These structures and 317 terrane boundaries served as loci for activation of transverse zones during the deposition of the 318 Mesoproterozoic Belt-Purcell sequence (Benvenuto & Price, 1979;Foo, 1979;Root, 1987; 319 Anderson & Davis, 1995;McMechan, 2012). Multiple other transverse structures were also 320 initiated in Mississippian and Triassic as normal faults above these transverse boundaries (Cooley 321 et al., 2011;McMechan, 2012). Segments of some of these transverse zones were reactivated as 322 dextral, oblique reverse faults or as tear faults during Jurassic to Eocene in response to 323 compressional stresses (Bielenstein, 1969;Foo, 1979;Price, 1981 Thompson, 1989;Pilkington et al., 2000;Berger et al., 2008). In northern British Columbia, 327 a geometric inversion of the northern end of a NW-trending "basement uplift" during the Sevier 328 contraction resulted in the formation of NE to E-trending (transverse) contractional structures, 329 which were superimposed on the regional NW-striking structures (McMechan, 2007). Aside from 330 their influence in the sedimentary and deformational history, these basement structures and terrane Range extension in the Eocene, polarity of the fault movement reversed, and this zone served as a 363 dextral, extensional, TZ that facilitated the exhumation of metamorphic core complexes ( Figure  364 5a) (Foster et al., 2007). 365 In the Sevier FTB (USA), the geometry of basement structures had first-order control on 366 the formation of orogenic curvatures and on the evolution of their transverse boundaries (e.g.; 367 Paulsen & Marshak, 1999). A deeper basin generally corresponds to a thicker sedimentary column 368 and more material available to be incorporated into the deforming taper, which results in a wider 369 wedge (salient) (e.g.; Marshak & Wilkerson, 1992;Boyer, 1995). The Helena salient likely formed are strongly converging into the right-lateral, reverse faults within the SWMTZ, due to a gradual 376 clockwise rotation of the shortening direction during their evolution (Whisner et al., 2014). 377 Further south, controls of the basin boundary geometry on the evolution of transverse zones 378 have been exemplified by the Mount Raymond Transverse Zone (MRTZ) and the Charleston TZ. 379 The MRTZ and the Charleston TZ form the northern and the southern boundaries of the 380 Uinta/Cottonwood arch (recess) with the Wyoming Salient and the Provo Salient respectively 381 ( Figure 5a) (Paulsen & Marshak, 1997, 1998. Paulsen and Marshak (1999) noted contrasting 382 structural styles between these two zones and explained this contrast in light of corresponding 383 basement structures. An east-west trending asymmetric basement high, with a gentle northern 384 flank and a steep southern flank, existed just north of the present Uinta/Cottonwood arch, along 385 the boundary between the Archean Wyoming province (north) and Proterozoic terranes (south) 386 (Paulsen & Marshak, 1999). A gentler northern flank meant that the sedimentary thickness 387 gradually increased northwards from the Uinta recess into the Wyoming Salient. The MRTZ 388 initiated above this flank as NNE-trending thrusts along the southern margin of the Wyoming 389 salient, which were later tilted northward creating an E-W strike during the uplift of the 390 Uinta/Cottonwood arch (Paulsen & Marshak, 1997). The steeper southern flank, however, marked 391 an abrupt increase in sedimentary thickness towards the south and thereby formed the boundary 392 between two contrasting taper wedges. The Charleston TZ (Figure 5a) served as zone of lateral ramp between the two contrasting tapers and gradually evolved into a left-lateral strike slip zone, 394 which accommodated the differential motion between the Uinta recess and the Provo salient 395 (Paulsen & Marshak, 1998 and references therein). The southern boundary of the Provo salient 396 with the central Utah segment is the Leamington TZ, which is an ENE trending, >50 m long, 397 complex, cross structure (Lawton et al., 1997;Kwon & Mitra, 2006). In the salient, an initial E-398 directed vergence over the TZ rotated clockwise during subsequent deformational phases, which 399 likely reflects the interaction between a deforming wedge and an oblique ramp (Lawton et al., 400 1997;Paulsen & Marshak, 1999;Kwon & Mitra, 2006). 401 In southern Wyoming, the E-to NE-trending Cheyenne belt (Figure 5b) represents the 402 transverse, crustal suture/ transpressional shear zone between the Wyoming and Yavapai-Mazatzal 403 Provinces, across which the Precambrian geology, metamorphism, and metallogenesis abruptly 404 change between adjacent blocks (Karlstrom & Houston, 1984). During the Laramide orogeny, this 405 weak crustal zone was reactivated as a left-lateral transpressional structure and subsequently as a 406 right-lateral transtensional zone during the Tertiary extension (Bader, 2008). Just to the south, east-  Following the initial collision, the eastern Alps underwent E-W directed orogen parallel 471 extension in the Miocene (Oligocene; Ring, 1994;Steck, 2008). This extension has been referred 472 to as "lateral extrusion" (Ratschbacher et al., 1991). The lateral extrusion has been interpreted as;   (Behrmann, 1988;Selverstone, 1988;517 Fiigenschuh et al., 1997), where blue-schist and eclogitic facies rocks of the Penninic zone have 518 been exposed (Pfiffner, 2014). Exhumation of the Tauern Window has been widely linked to 519 Miocene extension. Rosenberg and Garcia (2011) argue that localized intensive folding 520 deformation due to an irregular geometry of the Dolomite indenter coupled with erosion can also 521 exhume deep-seated rocks without a significant crustal extension. The southern edge of the 522 window has been cut by the NW-striking dextral, normal MV fault. Gently west-dipping mylonitic 523 fabric, with top-to-the-west shear along the Brenner Fault (Behrmann, 1988;Selverstone, 1988), 524 has been overprinted by steeply west-dipping cataclastic zones (Prey, 1989). Much like the 525 Brenner Fault, the Simplon Fault zone is a low-angle, SW-dipping, extensional fault that exhumes 526 the Lepontine Metamorphic Dome in its footwall. Mylonitic shearing that formed the Simplon 527 lateral/oblique ramps, initiated more lateral ramps, and produced complex geometries like en-572 echelon ramp folds and back thrusting (Schönborn, 1992). In the central Southern Alps, inherited 573 transfer zones cut through decoupling surfaces, partition thrust sheets into discrete blocks, and 574 thereby serve to generate a laterally heterogenous deformation style (Laubscher, 1985;Schönborn, 575 1992). Several transverse structures that have been identified in this area include the Lecco Line 576 and Ballabio-Barzio TZ ( Figure 6) (Laubscher, 1985;Schönborn, 1992;Zanchi et al., 2012). The 577 central Southern Alpine thrust belt has been compartmentalized by these transverse zones 578 (Schönborn, 1992). Moreover, many upper Triassic to Jurassic (Bernoulli, 2007), orogen-

Cross Structures in the Helvetics 602
In the Helvetic units, NW to N-striking tear faults were formed due to lateral variation in 603 shortening of the nappe stack, primarily during 35-30 Ma (Hunziker et al., 1986). Nappe 604 imbrication in the Helvetics changes laterally due to the absence or presence of a decoupling layer 605 lateral variations in sediment thickness and geophysical properties along the Himalayan foreland 645 (Burrard, 1915;Oldham, 1917). In the 1970's new data from the oil and gas industry connected 646 these lateral variations to a series of NE-trending basement ridges (Sastri, 1971;Rao, 1973;647 Raiverman, 1983). It was also proposed that these transverse basement structures may influence 648 Himalayan deformation (Valdiya, 1976). In recent years, evidence from field surveys and Geophysical data supports the merging of these thrusts into a master fault known as the Main 659 Himalayan Thrust (Zhao, 1993;Avouac, 2003;Nabelek, 2009). There is an extensional fault 660 system bounding the northern GHS, the South Tibetan Detachment System (STDS; Hodges, 2000). 661 The hanging wall of the STDS is the Tibetan or Tethyan Sedimentary Sequence (Gansser, 1964; including the local presence of dun structures and a series of recesses and salients (Yeats, 1991;680 Mukul, 2010). In some areas the irregular mountain front has been linked to differences in 681 shortening (Dubey, 2001)  Tethys, these faults were activated as right-lateral transform faults in the basement rocks (Talbot 789 & Alavi, 1996). Neogene reactivation of these transverse faults has produced a dragging effect on 790 the earlier-formed, orogen-parallel folds. These faults serve to transmit and distribute the slip along 791 the MRF towards the southeast into the Zagros Folded Belt and the Foredeep (Berberian, 1995;792 Authemayou et al., 2006). Salt diapirs have been intruded along these faults at multiple locations, 793 which suggests that these faults cut into the top of the basement (e.g.; Kent, 1979;McQuillan, 794 1991;Talbot & Alavi, 1996). A certain degree of seismic hazard is associated with these transverse  In the Dezful embayment, however, exposed folds are concentric folds with small wavelengths 815 and probably overlie large concentric folds beneath a detachment surface (Sepehr et al., 2006).

Tectonic Setting and Lateral Heterogeneities 884
The Andean mountain belt is a classic example of an active subduction margin and likely 885 represents processes that were active in the world's collisional mountain belts prior to collision.  highest topographic features in the world within a non-collisional setting (Isacks, 1988).  Cross structures have also been interpreted to have great economic and seismic impacts besides 1064 their influence in tectonic, stratigraphic, and structural evolution (e.g.; Mahoney et al., 2017). In the Papua New Guinea fold-thrust belt, the Jurassic, extensional, transverse, crustal structures 1066 evolved as zones of economically significant copper-gold mineralization during their Late 1067 Miocene-Pliocene inversion (Davies, 1991;Corbett, 1994;Hill et al., 2002). Jurassic salt and Middle-Lower Jurassic shales) in Russia (Sobornov, 1996), the Sulaiman belt 1077 (Paleozoic, Lower Cretaceous, and Eocene strata) in Pakistan (Jadoon et al., 1994), and the Parry 1078 Island belts (Ordovician salt) in the Canadian Arctic (Harrison & Bally, 1988).

Irregular Continental Margins 1097
The geometry of the continental margin(s), prior to the collision/convergence, has a profound 1098 effect on the geometry of the evolving orogenic belt. The orogenic front may mimic the continental 1099 margin geometry (Figure 4) such that any irregularities along the margin are reflected along the 1100 deformational font. This phenomenon has been proposed in the Appalachians. During the Iapetan 1101 rifting (plus the other rifting events) along the Laurentian margin, a series of embayments (concave 1102 oceanward) and promontories (convex oceanward) separated by transform faults were produced 1103 (Thomas, 2014). Along such an irregular margin, the depositional environment is bound to vary 1104 laterally. During subsequent orogenic events, promontories evolved into recesses (concave 1105 towards the foreland), embayments evolved into salients (convex towards the foreland), and the 1106 transform fault boundaries evolved into transverse zones or cross structures (Thomas, 2014). been noted that a failed rift-arm (aulacogen) can also serves as an inherited basement structure and 1122 control structural evolution (Rodgers, 1990;Brown et al., 1997;Perez-Estaun et al., 1997). Crustal 1123 sutures mark boundaries between two continental blocks and form as a result of collisional 1124 tectonics or terrane accretion. Generally, suture zones consist of highly deformed rock units and 1125 form weaker zones within the crust (Hope & Eaton, 2002;Whitmeyer & Karlstrom, 2007). These 1126 zones are known to be intruded by igneous bodies during the subsequent crustal stabilization (Whitmeyer & Karlstrom, 2007), and these intrusion boundaries can also behave as zone of 1128 weaknesses during various tectonic events (Simony & Carr, 1997;Bader, 2009).

Impacts of Basement Structures on Subsequent Sedimentation and Deformation 1138
The presence of basement structures can greatly affect the foreland sedimentation and thereby 1139 have a strong control on the evolution of mountain belts as is seen in the Zagros, the Cordillera, 1140 and the Himalaya. Basement structures can isolate deposition centers in the foreland, which 1141 induces sharp lateral variations in both the thickness and facies of the sedimentary strata. A thicker 1142 foreland sedimentary sequence, when incorporated into the evolving taper wedge, will propagate 1143 much farther (salient) than a wedge consisting of a thinner sequence (recess) (e.g.; Paulsen & 1144Marshak, 1999. Differential tectonic transport between the adjacent segments (salients and 1145 recesses) is often accommodated by cross structures such as tear faults, lateral ramps, and 1146 displacement transfer zones (Paulsen & Marshak, 1999). Orogen-parallel faults and folds are 1147 generally truncated at, or dragged into cross structures as seen in the Appalachians and the 1148 Cordillera (e.g.; Thomas, 2007;Whisner et al., 2014). 1149 Similarly, lateral variations in sedimentary facies will also have a significant impact on the 1150 wedge evolution. An abrupt change in the structural elevation of decoupling layers results in lateral 1151 ramps (Thomas, 1990). Since thrust-ramp geometry is governed by the nature of the sedimentary 1152 column, any lateral variation in sedimentary facies is likely to be reflected in the thrust geometry 1153 (e.g. ;Mitra, 1988). In a broader stratigraphic framework, if the presence of regional decollement 1154 horizons (units of salts, evaporites, and shales) varies laterally, then the deformation style in 1155 adjacent segments of the mountain belts will be very different. If a decollement horizon is present, 1156 then the taper angle will be low and propagate further than adjacent areas. Likewise, in the absence 1157 of such a horizon, the taper will build up to a larger angle in a narrower belt as in the Zagros (e,g,; 1158 Jadoon et al., 1994;Sobornov, 1996;Bahroudi & Koyi, 2003). The depth of these decollement 1159 horizons will also have a first order control on the fold geometry, such that a deeper horizon will 1160 result in folds with larger amplitude (Sepehr et al., 2006). Furthermore, the stratigraphic variation 1161 boundaries will also act as a lateral buttress to locally deflect the transport trajectories (Bahroudi 1162& Koyi, 2003. If basement structures are reactivated during deposition, then syn-sedimentary 1163 faults and/or drape folds can form in the overlying sedimentary column (Thomas, 1990). These 1164 structures will also impact the fold-thrust belt evolution in a similar fashion to other cross-1165 structures (e.g.; Zerlauth et al., 2014). 1166

Reactivation of Basement Structures 1168
Regardless of their origin, inherited basement structures can be broadly categorized into three 1169 groups: orogen-parallel, oblique, and transverse/cross basement structures. Various styles of 1170 preferential reactivation and inversion of basement structures have been discussed in the 1171 Apennines (Tavarnelli et al., 2004;Butler et al., 2006), and are also seen in other mountain belts. 1172 When orogen-parallel basement structures/faults are present, it is possible for them to be 1173 reactivated as reverse faults during orogenic compression (Figure 10a). It has, however, been noted 1174 that basement faults can be deformed under compression without inversion or reactivation (Pantet 1175(Pantet et al., 2020. In a reactivation scenario, the deformation between the basement and the overlying 1176 Variation in crustal rheology is another significant element in inducing lateral heterogeneities 1194 in orogenic belts, which itself is a function of its composition and thermal structure. In general, a 1195 weaker crust is more likely to buckle under contraction resulting in a high crustal relief, whereas 1196 a stronger crust will distribute deformation in more brittle manner (Allmendinger et al., 1997; margin front could result in orogenic curvatures, such that a stronger segment better resists deformation and evolves as a salient (e.g.; Malekzade et al., 2016). As seen in the Zagros, frictional 1200 strength in hinterland faults, which is governed by crustal rheology, can influence the mode of 1201 deformation partitioning in the orogenic belt and thereby have a control on the style and 1202 distribution of deformation (e.g.; Vernant & Chéry, 2006). Thermal structure and rheology of the 1203 crust can vary laterally due to variations in arc magmatism. Therefore, previous tectonic events 1204 will also have an indirect impact on lateral heterogeneity of an orogen. coupling is thought to be governed by crustal composition such that a thick and more felsic crust 1208 has a stronger coupling between the upper and lower units than a thin and mafic crust. Thus, if 1209 there is strong lateral variation in crustal properties due to tectonic events such as terrane accretion 1210 or rifting then such variation will be manifested as laterally heterogenous deformation segments. In orogens associated with oceanic subduction, variations in the relative rates of subduction 1235 versus convergence can create lateral heterogeneity in the over-riding plates as is seen in the 1236 Andes. When subduction outpaces convergence such that the convergent boundary retreats, back-1237 arc extension occurs. These areas may have contrasting evolution when compared to adjacent areas 1238 where there is a better balance of the subduction versus convergence rates. Piquer et al., 2020). Thermo-mechanical tectonic models could potentially be rectified by changing 1275 the lithospheric parametric values along the strike such that they agree with the lateral variation in 1276 crustal deformation, as can be observed on the surface. Finally, cross-structures also have a major 1277 control on the geomorphologic evolution of mountain belts. In the Himalayan front, cross faults 1278 spatially coincide with river channels (Sahoo, 2000;Srivastava, 2018). The anti-Apennine faults 1279 have distinct topographic signatures (Coltorti et al., 1996). Cross-faults can be zones of 1280 weaknesses, and therefore, can also amplify erosional hazards, especially in active mountain belts. 1281

Conclusions 1282
While convergent mountain belts are dominated by spatially and temporally continuous 1283 orogen-parallel structures, geological and geophysical data shows that various forms of lateral 1284 heterogeneity, often marked by cross structures, are ubiquitous in most orogenic settings. In 1285 general, lateral heterogeneities along orogens have been manifested as: 1) along-strike changes in 1286 deformation style; 2) variation in igneous activity or metamorphic grade; 3) variation in seismic 1287 activity; 4) changes in topography and geomorphology; and 5) abrupt lateral stratigraphic changes. 1288 Common drivers behind these lateral heterogeneities include the geometry of the continental 1289 margin, inherited basement structures, lateral variation in stratigraphy of deforming sedimentary 1290 sequences, variation in crustal rheology, along-strike changes in plate tectonic setting, physiography of the lower plate, and obliquity of plate convergence. In most settings, these factors 1292 are interrelated and simultaneously influence the morpho-tectonic evolution of an orogen. Apart 1293 from their influence on foreland sedimentation and orogenic evolution, lateral heterogeneity and 1294 cross structures can have an impact on patterns of seismicity, natural resource occurrence, and 1295 natural hazards. We therefore stress the importance of documenting heterogeneity, mapping cross 1296 structures, and understanding the role these lateral changes play in mountain belt development 1297 along convergent margins.  Cembrano, J., Hervé, F., & Lavenu, A. (1996). The Liquiñe Ofqui fault zone: a long-lived intra-1445 arc fault system in southern Chile. Tectonophysics, 259(1-3), 55-66.