significant role in the development of the magma-poor southern North Atlantic and 50

Alpine Tethys rift systems and that the role of inheritance may change as the physical 51 conditions vary during the evolving rifting and as rift-induced processes (serpentinization; 52 magma) become more important. Thus, it is not only important to determine the "genetic 53 code" of a rift system, but also to understand how it interacts and evolves during rifting. 54 Understand how far these new ideas and concepts derived from the well-studied 55 hyperextended rift systems of the southern North Atlantic and Alpine Tethys can be 56 translated to other less explored hyperextended rift systems will be one of the challenges 57 of the future research in rifted margins. 58 59 60 61 1. Introduction 62 Although there is general agreement in the community that inheritance has an important 63 influence on the architecture and tectonic evolution of rifted margins, it is not yet clear 64 how inheritance controls the initiation and evolution of rifting and lithospheric breakup. 65 Most studies advocating for a strong role of inheritance refer to the observed along strike 66 changes of the rift architecture and segmentation of rifted margins that seem to correlate 67 with changes in the onshore geology (Dunbar and Sawyer, 1989; Pique and Laville 1996; 68 Ring 1994; Tommasi and Vauchez, 2001). While Taylor et al. (2009) and Gerya (2012) 69 showed that major oceanic transform faults do not extend into the continents, Behn and 70 Lin (2000) discussed the importance of the segmentation observed along the U.S. East 71 coast passive margin for the incipient structure of the oceanic lithosphere. Thus, at present 72 it is not clear what is the importance of crustal inheritance in controlling the segmentation 73 in oceanic domains. Other studies discussed the reactivation of pre-existing faults, or 74 localization of deformation along major lithological boundaries (Butler et al. 2006; 75 (Bellahsen et al., 2012). These observations are most of the time local and are unable to 76 explain why early stages of rifting are often diffuse and do in most cases not coincide 77 with the location of final breakup (e.g. Dunbar and Sawyer 1989; Bertotti et al. 1993; 78 Manatschal and Bernoulli 1998). In the North Atlantic (Lundin and Doré, 2011) and the 79 NW Shelf of Australia (Direen et al. 2008), the final breakup cuts not only through 80 ancient orogenic belts, but also through earlier rift systems, suggesting that inheritance 81 does neither control the location of breakup, nor the onset of seafloor spreading. 82 Unfortunately, thick sedimentary and/or magmatic sequences prevent detailed 83 investigation of the pre-rift crustal and lithospheric structures and composition. Thus, 84 although the common belief is that inheritance has an important control on the evolution 85 of rift systems, at present it is difficult to assess under what conditions and how it controls 86 rifting and lithospheric breakup. 87

The aim of this paper is to define types of inheritance and to show, using 88 observations and models, how these different types of inheritance may control subsequent 89 rifting events. We use examples from the ancient and present-day Alpine Tethys, 90

Pyrenean-Bay of Biscay and Iberia-Newfoundland rifted margins to illustrate and discuss 91 ideas and concepts on how to integrate the role of inheritance in the study of rifted 92 margins. Since all rift systems discussed in this paper affect Variscan lithosphere and are 93 magma-poor, the ideas developed in this paper may not be applied globally. Thus, the aim 94 is not to propose models that can explain in a global way the influence of inheritance on 95 rift systems, but to discuss how "inheritance" can be described and integrated in 96 descriptions of hyperextended rift models leading to seafloor spreading.

97

Due to the very limited data from hyperextended rifted margins available in the 98 open domain and the difficulty to link their yet little understood evolution to inherited 99

The role of inheritance in structuring hyperextended rift A long-standing question in Earth Sciences is related to the importance of 8 inheritance in controlling tectonic processes. In contrast to physical processes that are 9 generally applicable, assessing the role of inheritance suffers from two major problems: 10 firstly, it is difficult to appraise without having insights into the history of a geological 11 system; and secondly all inherited features are not reactivated during subsequent 12 deformation phases. Therefore, the aim of this paper is to give some conceptual 13 framework about how inheritance may control the architecture and evolution of 14 hyperextended rift systems. 15

In this paper, we use the term inheritance to refer to the difference between an 16

"ideal" layer-cake type lithosphere and a "real" lithosphere containing heterogeneities. 17

The underlying philosophy of this work is that the evolution of hyperextended rift 18 systems reflects the interplay between their inheritance (innate/"genetic code") and the 19 physical processes at play (acquired/external factors). Thus, by observing the architecture 20 and evolution of hyperextended rift systems and integrating the physical processes, one 21 my get hints on what may have been the original inheritance of a system.

22

We first define 3 types of inheritance, namely structural, compositional and 23 thermal inheritance and develop a simple and robust terminology able to describe and link 24 observations made at different scales using geological, geophysical and modelling 25 approaches. To this, we add definition of rift-induced processes, i.e. processes leading to 26 compositional changes during rifting (e.g. seprentinization or decompression melting).

27

Using this approach, we focus on 3 well-studied rift systems that are the Alpine Tethys, 28

Pyrenean-Bay of Biscay and Iberia-Newfoundland rift systems. However, as all these 29 examples are magma-poor, hyperextended rift systems that evolved over Variscan 30 lithosphere the concepts developed in this paper cannot be applied universally. For the 31 studied examples we can show that: 32 1) the inherited structures did not significantly control the location of breakup and the 33 structures in the southern North Atlantic example 34

2) the inherited thermal state may control the mode and architecture of rift systems, in 35 particular the architecture of the necking zone 36

3) the architecture of the necking zone may be influenced by the distribution and 37 importance of ductile layers during decoupled deformation and is consequently 38 controlled by the thermal structure and/or the inherited composition of the curst 39 4) conversely, the deformation in hyperextended domains is strongly controlled by weak 40 hydrated minerals (e.g. clay, serpentinite) that result form the breakdown of feldspar 41 and olivine due to fluid and reaction assisted deformation 42 5) inherited structures, in particular weaknesses, are important in controlling strain 43 localization on a local scale and during early stages of rifting 44 6) inherited mantle composition and rift-related mantle processes may control the 45 rheology of the mantle, the magmatic budget, the thermal structure and the localization 46 of final rifting 47 48

These key observations show that both inheritance and rift-induced processes played a 49 orogenic structures, we will focus our paper on the well explored Variscan system in 100

Western Europe and its relation to the late rift evolution of the Alpine Tethys and 101 southern North Atlantic. Due to the novelty of the subject, this paper does not review but 102 does rather present some considerations on how inheritance may control hyperextended 103 rift systems based on observations and simple numerical models. 104 105 2. Defining inheritance in rift systems and rifted margins 106

Although the terms "inheritance" and "hyperextended rifted margins" are often 107 used, clear and universal definitions are necessary in order to discuss the impact of 108 inheritance on hyperextended rift systems. The aim of this chapter is to develop a 109 terminology that is simple, robust and that can describe and link observations made at 110 different scales using geological, geophysical and modelling approaches. The underlying 111 philosophy is that each rift system inherited a set of particular properties and boundary 112 conditions similar to a "genetic code", and that the evolution of rift systems reflects the 113 interplay between their inheritance (innate) and the physical processes at play (acquired) 114 during rifting. 115 116

Inheritance 117

In contrast to oceanic domains that are newly formed and do, in theory, not bear 118 inheritance, continental domains are the result of polyphase tectonic, magmatic and 119 metamorphic events, often resulting from the superposition of several Wilson cycles. 120

Continental domains preserve a long history with a complex lithological and structural 121 record that cannot be described by few physical parameters only. Therefore, defining the 122 original state of the lithosphere at the onset of a new tectonic phase, in this case rifting, is 123 difficult. In this study, we use the term "inheritance" to refer to the difference between a 124 "real" and an "idealized" lithosphere (Fig. 1). The "idealized" lithosphere (Fig. 1a) is 125 defined as a thermally equilibrated horizontally homogeneous layer-cake lithosphere 126 made of a quartzo-feldspatic upper and middle crust and a mafic lower crust and a 127 peridotitic mantle (for definition of single layers see also [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF]. In contrast 128 to this "idealized lithosphere", a "real" lithosphere (Fig. 1b) may differ by its thermal 129 state, compositional heterogeneities and structural complexity. 130

We distinguish between three types of inheritance, namely "thermal", "structural" 131 and "compositional" inheritance. Although the three can, practically, not be disassociated, 132

for simplicity we will discuss and exemplify the three independently. Figure 1b shows, in 133 a conceptualized manner, the 3 types of inheritance in a section across a lithosphere at the 134 end of an orogenic collapse but before onset of rifting. In this paper we only show 135 examples of rift systems that developed over Variscan post-orogenic, lithosphere. Other 136 types of lithospheres, not discussed in this paper, are cratonic or complex accreted 137 lithosphere, the latter formed by different terrains and separated by multiple suture zones 138 (e.g. Western North America). 139 140

Thermal inheritance 141

Two components have to be considered: 1) inherited heat and 2) potential of heat 142 production. Inherited heat decreases with time and is consequently a function of the age of 143 the lithosphere. This implies that the older the lithosphere, the colder the thermal structure 144 and the stronger the lithosphere. Thus, the time lag between the last thermal perturbation 145 of the lithosphere and the onset of rifting is of major importance in assessing the thermal 146 state of the lithosphere at the start of a rifting event. In contrast, the potential of heat 147 production depends on the presence of radiogenic elements, which is mainly controlled by The existence of structures inherited from previous orogenic events in the crust 176 and the underlying lithospheric mantle result in mechanical weak zones. These "weak" 177 structures may correspond to ancient thrust belts or remnants of subduction zones. When 178 these structures formed in the brittle layers, they are often linked to hydration processes.

179

When they occur in the ductile layers, mainly as mylonitic shear zones characterized by 180 grain size reduction and crystallographic preferred orientations, they can result in a strong 181 anisotropy [START_REF] Tommasi | Continental rifting parallel to ancient collisional belts: 1663 an effect of the mechanical anisotropy of the lithospheric mantle[END_REF][START_REF] Ring | The influence of preexisting structure on the evolution of the Cenozoic 1625 Malawi rift (East African rift system)[END_REF]. Of particular importance are intra-182 continental strike-slip systems, because these structures can affect the whole lithosphere 183 [START_REF] Tommasi | Continental rifting parallel to ancient collisional belts: 1663 an effect of the mechanical anisotropy of the lithospheric mantle[END_REF], conversely to extensional or compressional structures, 184 which are typically limited to the brittle lithosphere and become more diffuse in the 185 ductile layers. It is also important to note that not every inherited structure has the 186 potential to be reactivated. Reactivation depends on its orientation relative to the stress 187 field, on it's lateral continuity and on the rheological contrast of the encompassing rocks 188 relative to the other lithospheric layers (e.g. [START_REF] Chenin | Influence of offset weak zones on the development of 1445 rift basins: activation and abandonment during continental extension and breakup[END_REF] 196 Pinvidic et al., 2013a). These studies resulted in new observations and interpretations of 197 how rift systems evolve in time and space. This process enables us to describe the 198 evolution of rifted margins and to classify them according to the following characteristics: 199 the mid-ocean ridge from breakup and onset of steady-state magmatic activity. However, 250 more and more evidence show that many rifts initially fail, migrate or step outside the 251 initial rift system before breakup occur resulting in so called "offset basins" (Chenin and 252 Beaumont, 2013). Here we call this phenomenon "polyphase rifting". Well documented 253 examples of polyphase rift include the NW Shelf offshore Australia (Direen et al. 2008), 254 the North Atlantic [START_REF] Lundin | Hyperextension, serpentinization, and weakening: A new 1534 paradigm for rifted margin compressional deformation[END_REF], the northern segment of the South Atlantic 255 [START_REF] Blaich | Northeastern Brazilian margin: Regional 1416 tectonic evolution based on integrated analysis of seismic reflection and potential 1417 field data and modelling[END_REF] Akin to modern orogens, the Varsican orogen overprinted a continental 291 crust/lithosphere inherited from older orogenic events. The limit between parts of the 292 crust/lithosphere that were affected relative to those that remained unaffected by a severe 293

Variscan overprint corresponds to the transition between the fold and thrust belt and the 294 flexural foreland basin (Fig 3a andb). While this limit is well preserved across Northern 295

Germany and Southeastern England and Ireland, its southern limit is strongly overprinted 296 by the Mesozoic rifting and the Alpine orogeny (see Fig. 3). The main difference between 297 the distinct orogenic domains shown in Fig. 3a of a complex evolution that ended with major late-to post-Variscan extensional and 333 magmatic events. Although the exact timing and processes of this evolution are not yet 334 well constrained on the scale of the whole orogen, it looks as if at about 310 to 300 Ma 335 the magmatic system changed from an arc-dominated to an intracontinental magmatic 336 setting. This change is associated with the emplacement of prominent underplated mafic 337 bodies and extrusive magmatic systems simultaneous with crustal thinning (Fig. 3c).

338

Evidence for magmatic underplating is widespread and can be observed in all exposed 339 sections preserving pre-rift lower crust (e.g. Ivrea zone in the Southern Alps (Peressini et 340 al. 2007), the Malenco-Campo-Grossina units in the Eastern Alps [START_REF] Hansmann | U-Pb zircon geochronology of a 1489 tholeiitic intrusion and associated migmatites at a continental crust-mantle 1490 transition, Val Malenco, Italy[END_REF]341 Tribuzio et al. 1999), the Santa Lucia complex [START_REF] Rossi | Variscan to eo-Alpine events 1630 recorded in European lower-crust zircons sampled from the French Massif Central 1631 and Corsica, France[END_REF] from the French massif [START_REF] Féménias | A Permian underplating event in late-to post-orogenic tectonic setting. 1467 Evidence from the mafic-ultramafic layered xenoliths from Beaunit (French Massif 1468 Central)[END_REF]).

344

This event, also referred to as the "Permian event", was concomitant with a 345 change in the morpho-tectonic evolution of the former orogenic system. It is expressed as 346 a shift from a classical mountain belt topography to one dominated by plateaus dissected 347 by intra-continental sedimentary basins [START_REF] Burg | Syn-to post-thickening extension : 1437 mode and consequences[END_REF][START_REF] Froitzheim | Alpine tectonics of the Alps and 1475 Western Carpathians[END_REF]). This 348 evolution can be compared with the observed orogenic collapse in the Basin and Range 349 Province of the western US. The Permian Variscan collapse is not only recorded by the 350 important magmatic additions to the crust (intrusions and underplating), but also by 351 granulite facies metamorphism [START_REF] Galli | 1479 Granulites and charnockites of the Gruf Complex: Evidence for Permian ultra-high 1480 temperature metamorphism in the Central Alps[END_REF], see also [START_REF] Mohn | Unravelling the interaction between tectonic and sedimentary processes during 1564 lithospheric thinning in the Alpine Tethys Margins[END_REF] for a review), 352 which indicates a very high thermal gradient (Müntener et al. 2000, Schuster andStüwe 353 2008). These authors described the thermal evolution during the late-to post-orogenic 354 event and determined the thermal state of the lithosphere at the end of Permian time, i.e. 355 before onset of rifting. These studies showed that the magmatic and thermal evolution 356 recorded in the continental crust is intimately related to processes acting in the underlying 357 mantle lithosphere (Fig. 3c). As discussed in the following, these processes strongly 358 controlled the final structure of the lithosphere at the end of the Variscan cycle and pre-359

determined the conditions at the onset of the Mesozoic rifting. Mantle rocks drilled along the conjugate Iberia-Newfoundland margins (e.g. ODP 373

Sites 637 (Galicia Margin), 987, 899, 1068, 1070 (Iberia Abyssal Plain) and 1277 374 (Newfoundland) (see Fig. 4b) show distinct compositional heterogeneities (Fig. 4a; 375 [START_REF] Müntener | High degrees of melt extraction recorded by spinel 1583 harzburgite of the Newfoundland margin: The role of inheritance and consequences 1584 for the evolution of the southern North Atlantic[END_REF] and references therein). While the mantle rocks from 376 the Newfoundland margin (ODP Site 1277) are made of depleted harzburgites (high 377 rations of Cr# in clinopyroxene vs. spinel), those from the Iberia margin are less depleted 378 and consist of infiltrated plagioclase peridotites [START_REF] Müntener | High degrees of melt extraction recorded by spinel 1583 harzburgite of the Newfoundland margin: The role of inheritance and consequences 1584 for the evolution of the southern North Atlantic[END_REF]. The 379 authors interpreted the strongly depleted harzburgites drilled at ODP Site 1277 at the 380 Newfoundland margin as derived from the hanging wall of a Caledonian subduction (Fig. 381 4b). Thus, these mantle rocks sampled the strongly depleted supra-subduction mantle that 382 may have produced the magma of the Bay of Island arc complex now exposed in 383

Newfoundland. This interpretation assumes that mantle rocks from the Ocean Continent 384

Transition (OCT) are derived from the pre-rift sub-continental mantle proving the 385 existence of a strong mantle inheritance. The observation that the mantle rocks from the 386

Iberia-Newfoundland margins show different degrees of depletion is a key observation, 387 because mantle composition is a key parameter controlling the magmatic volume that can 388 be produced during rifting. 389 390 3.1.2.2 Post-Variscan to early rift inheritance 391

In the Alps and Pyrenees, remnants of mantle rocks are common and associated 392 with lower to middle crustal granulites and deep marine sediments (Jammes, 2009;393 Lagabrielle, 2009;[START_REF] Manatschal | A type sequence across an ancient magma-poor 1544 ocean continent transition: the example of the western Alpine Tethys ophiolites[END_REF]) (Fig. 4d). Studies of these mantle 394 rocks in the Pyrenees (Lherz, LeRoux et al. 2007), Liguria and Corsica [START_REF] Piccardo | Mantle processes during ocean formation: Petrologic records in 1615 peridotites from the Alpine-Apennine ophiolites[END_REF]395 Rampone and Hofmann, 2012) (Montanini et al. 2012, and the Alps (Lanzo, Malenco, 396 Totalp; [START_REF] Müntener | Plagioclase Peridotites in 1587 Ocean-Continent Transitions: Refertilized Mantle Domains Generated by Melt 1588 Stagnation in the Shallow Mantle Lithosphere[END_REF][START_REF] Piccardo | Geodynamic evolution of the Jurassic Ligurian Tethys viewed from 1617 the Mantle perspective[END_REF] showed that the mantle rocks 397 exhibited different degrees of modifications prior and during rifting. The uppermost 398 subcontinental mantle that locally still preserves primary contacts to the overlying crust 399 (e.g. Manlenco, Hermann et al. 1997) was too cold to either melt or react with magma 400 during the Permian orogenic collapse and subsequent rifting. This type of mantle, referred 401 to as inherited mantle, is exposed in the most proximal parts of the OCT where it is often 402 associated with remnants of the thinned continental crust (e.g. Malenco or Tasna; 403 [START_REF] Müntener | Refertilization 1579 of mantle peridotite in embryonic ocean basins: trace element and Nd-isotopic 1580 evidence and implications for crust-mantle relationships[END_REF][START_REF] Müntener | Plagioclase Peridotites in 1587 Ocean-Continent Transitions: Refertilized Mantle Domains Generated by Melt 1588 Stagnation in the Shallow Mantle Lithosphere[END_REF] (Fig. 4d). In contrast, the mantle rocks occurring in the more 404 oceanward part of the OCT show a more complex evolution (Fig. 4d). Müntener et al. 405 (2004) showed that these mantle rocks record two major events: 1) a strong depletion 406 during the late-to post-orogenic event; and 2) an enrichment of the mantle during early 407 stages of rifting. The first event is dated as Permian by Sm/Nd model ages, interpreted to 408 date the melting of this mantle under asthenospheric conditions (Rampone and Hofmann,409 2013; [START_REF] Müntener | Refertilization 1579 of mantle peridotite in embryonic ocean basins: trace element and Nd-isotopic 1580 evidence and implications for crust-mantle relationships[END_REF], 2010[START_REF] Piccardo | Geodynamic evolution of the Jurassic Ligurian Tethys viewed from 1617 the Mantle perspective[END_REF]. This event may be linked 410 with the emplacement of the gabbros shown in Fig. 3c at the base of the crust during 411

Permian time. The second event is associated with the infiltration and reaction of magma 412 with the previously depleted mantle. Thus, this second event has to be post-Permian but 413 pre-exhumation. These observations show that the subcontinental mantle records a 414 complex, polyphase evolution including depletion and fertilisation cycles. Thus, the 415 inherited subcontinental mantle underlying the domains that underwent rifting 416 experienced a strong late-to post-Variscan imprint. How this may have controlled the 417 subsequent rifting is discussed in the following section. 418 419

Control of Variscan inheritance on the large-scale evolution of the Tethyan and 420

Iberian rift systems 421 422

The extent of late Variscan overprint in the lower crust and mantle compared to pre-423 and syn-orogenic compressional structures begs to the question of their relative influence 424 on the large-scale evolution of the subsequent rift system. In the following, we discuss the 425 role of the Variscan inheritance, in particular of lower crustal and mantle compositional 426 inheritance in controlling the large-scale strain distribution and magmatic evolution 427 during the subsequent rift event.

428

An interesting observation is that the location of hyperextension and later breakup 429 of the Alpine Tethys and Iberian Atlantic localized along the former boundaries of the 430

Variscan orogeny, leaving the central parts -thus all the previous suture zones -of the 431 orogenic lithosphere preserved (e.g. Bohemian massif, Massif Central and Armorican belt 432 and in the Iberian peninsula; Fig. 3a). This observation questions if the location of 433 continental necking and lithospheric breakup is pre-determined by orogenic or post-434 orogenic processes.

435

How far gravitational processes, i.e slab breakoff and/or delamination of the 436 orogenic roots of an over-thickened orogenic crust are responsible for the subsequent rift 437 evolution compared to late-to post-Variscan inheritance is yet unclear. The observation 438 that in Western Europe crustal necking and breakup affected the boundaries between 439 internal and external parts of the former orogeny may suggest an "orogenic" control. One 440 factor that may control the location of necking is the late-to post-Variscan mafic 441 underplating, associated with the depletion of the underlying lithospheric mantle (Fig. 4c; 442 further discussion in section 5). Indeed, this process seems to be widespread but localized 443 underneath internal parts of the Variscan orogen (Fig. 3b). Thus, as shown by the 444 observations reported above, the Permian magmatic event strongly modified the crustal 445 and mantle composition underneath the realm affected by the Variscan orogen. This event 446

is well documented in all the geological sections that expose remnants of the Late 447

Variscan lower crust and mantle in the former distal portions of the Tethyan margin. It is 448 also drilled along the most distal Iberia margin (ODP Site 1967 in Fig. 4b), found in 449 xenolithes all over Western Europe [START_REF] Féménias | A Permian underplating event in late-to post-orogenic tectonic setting. 1467 Evidence from the mafic-ultramafic layered xenoliths from Beaunit (French Massif 1468 Central)[END_REF] and in remnants exhuming 450 lower crustal rocks in the Alps (e.g. Malenco;Herrmann et al. 1997). This suggests that 451 the last thermal event that may have modified in a sever way the lithosphere and modified 452 the composition of the lithospheric mantle and the lower crust is not related to the 453

Variscan collision, but the late-to post-orogenic extensional collapse of this orogen. In 454 the case of the Variscan orogeny, the post-orogenic collapse resulted in a stronger 455 lithosphere underneath previously overthickened crust comparing to the more external 456 and thinner crust in the former fold and thrust belt that was not modified (Fig. 3b).

457

A determinant observation for the magmatic evolution of the rifting event is that the 458 lithospheric mantle underneath Western Europe exhibits different degrees of 459 modifications prior and during rifting (Fig. 4a). The existence of depleted and enriched 460 mantle domains underneath crustal domains undergoing extension is important, since it 461 may explain differences in the magma production during rifting. Models assuming a 462 homogeneous subcontinental mantle explain the lack of syn-rift magma as either related 463

to very slow extension rates or a very cold mantle lithosphere or a combination of the two.

464

The depleted mantle domains that formed during the late to post-orogenic Permian event 465

may have lacked the necessary mineral components to produce melts during the 466 subsequent extensional event. It may explain why the rift system remained magma-poor. The architecture of the necking zone along the Western Iberia margin is shown in 500 three sections in Fig. 5a). The northern section corresponds to the ISE1 line (Iberia 501 Seismic Experiment; [START_REF] Zelt | Assessment of crustal velocity 1701 models using seismic refraction and reflection tomography[END_REF], the central section to the composed LG12 -TGS 502 line [START_REF] Sutra | How does the continental crust thin in a hyperextended 1648 rifted margin? Insights from the Iberia margin[END_REF] and the southern section to the IAM 9 line (Dean et al., 503 2000) (for location along the Iberia margin see Fig 3b). All sections are dip lines, i.e. 504 parallel to the kinematic transport direction. The necking domain has been defined as the 505 domain limiting original thick crust (30km) from crust thinner than 10km. The necking 506 zone is defined by an oceanward increase in the total ß value, space accommodation and 507 the coupling of the deformation between crust and mantle (for a more detailed definition 508 of the "necking domain" see [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF]. In the northern section (ISE 1) the 509 necking profile is wide and defined by a double neck. In the central section (LG12-TGS) 510

the necking profile shows a single neck of intermediate width. In the southern section 511 (IAM 9) the necking is sharp. In contrast to the change of the necking profile along strike 512 of the margin, it is interesting to note that the width as well as the angle of aperture (angle 513 between top and base of the basement, the later corresponding to a "S" type reflection; e.g. 514 [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF]) of the coupled domain remains similar along the margin (Fig. 5a). 515 516 Implications 517

Thinning of the crust/lithosphere is one of the most fundamental processes related 518 to rifting. Extreme crustal thinning is typically localized within the so-called necking 519 domain, which can display strong lateral changes, as shown for the Iberian margin (Sutra 520 and Manatschal, 2012). The necking profile at rifted margins can be controlled by the 521 inherited bulk rheology of the crust, in particular by the relative distribution of ductile vs. 522 brittle layers [START_REF] Mohn | Necking of 1567 continental crust in magma-poor rifted margins: Evidence from the fossil Alpine 1568 Tethys margins[END_REF]; see also discussion of modelling results in Section 4).

523

The observations show that the more ductile the crust, the more difficult it is to neck. 524

Conversely, strong and more rigid crust tends to produce sharper necks. Other controlling 525 parameters are the obliquity of extensional systems and the strain rate (Jammes et al., 526 2010). For the example, the Iberian margin formed by orthogonal extension more than 527 100 myr after the last thermal imprint, suggesting that the lithosphere was thermally 528 equilibrated, and under similar strain rates along its whole length (see [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF].

529

As the necking profiles are quite different along the margin, one can assume that their 530 architecture is primarily controlled by the composition of the crust and mantle. If we 531 assume a homogeneous mantle composition, the change in the necking profile observed 532 along the Iberia margin may provide some information about the inherited crustal 533 composition and indirectly about the bulk rheology of the crust at the onset of rifting. As 534 shown in Fig. 5b, the 3 sections cross different parts of the Variscan belt, which are likely 535 to show different crustal compositions (i.e. different percentages of ductile material and/or 536 inherited structures able to localize deformation). In the North (Fig. 5b), where the 537 necking profile is a wide double neck, the crustal structure is likely to be complex, 538 probably bearing several brittle-ductile transitions and/or strong lateral variations of the 539 crustal structure. In the South, where the necking profile is narrower, the crustal structure 540 is likely to be simpler, consisting of more brittle material and/or controlled by an 541 important inherited crustal scale structure [START_REF] Alves | Contribuition from potential field methods to regional tectonic studies 1391 of the Western Iberian Continental Margin. B. Geocience Petrobras[END_REF]. 542

Another important observation is that all sections shown in Fig. 5a display similar 543 width and angle of aperture of the coupled domain along strike (Fig. 5a). The map in Fig. 544 5c shows that the width of the coupled domain, which lies between the point of 545 exhumation and the coupling point (see [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF] for details), does not change 546 along the margin, conversely to the width of the necking domain. Deformation in the 547 coupled domain only occurs via brittle processes, while it is primarily controlled by 548 ductile layers in the decoupled domain. 549

Based on the drilling results, the refraction velocities and in analogy with material 550 that can be found in comparable positions in the Alps, we assume that the crustal rocks 551 are strongly hydrated and that the bulk rheology of the hyperextended crust is mainly 552 controlled by the newly created phylosilicates such as illite, chlorite and serpentinite. This 553 would explain that, when the crust is thinned to less than 10 km and becomes entirely 554 brittle, the initial compositional or structural differences do not impact deformation any 555 more and all crusts, independent of the initial composition, start to deform in a similar 556 way. Thus, while thinning and formation of the necking domain may be controlled, on a 557 crustal scale, by inheritance, deformation in the hyperextended and exhumation domains 558 is most likely not controlled by inheritance but by hydration processes, i.e. by a rift-559 induced process. This applies, however, only to the uppermost parts of the hyperextended 560 crust and exhumed mantle (< 6km deep). On a lithospheric scale, the thinning of the 561 lithosphere and onset of magma may eventually control the localization of deformation 562 and magma-production controlling final breakup (see discussion section for more 563 discussion). 564 565

Control of inheritance on the segmentation of rifted margins 566

To what extend oceanic structures in general and transform faults in particular are 567 inherited from the previous rift system is highly debated [START_REF] Bonatti | 1425 Transform migration and vertical tectonics at the Romanche Fracture Zone, 1426 equatorial Atlantic[END_REF]. Taylor et al. 568 (2009) showed that oceanic transform faults do not penetrate into the margin, and (Gerya, 569 2012(Gerya, 569 , 2013) ) showed that prominent structures observed on the continent do not necessary 570 continue into oceanic crust. However, at most present-day rifted margins the initial stages 571 associated with the development of highly segmented rift-transform margins are often 572 masked by thick sedimentary sequences and the relation between the rift structures, syn-573 tectonic sediments and magmatic additions remain poorly constrained. A good place to 574 understand the role of inheritance on the segmentation of margins and hyperextended 575 domains, as well as the location of breakup and structures within the oceanic crust is the 576 Iberia Atlantic and Bay of Biscay -Pyrenean domain (Fig. 6). Alves (2011) showed that major lineaments observed and mapped onshore Iberia 584 can be observed and followed on potential field maps into the Zone of Exhumed 585

Continental Mantle (ZECM). This observation supports not only the idea that the mantle 586 is of subcontinental origin in the ZECM (see previous chapter), but also that these 587 structures are rooted into the lithospheric mantle. However, these lineaments disappear 588 further oceanwards as indicated by the observation that the magnetic anomaly 34 is 589 continuous and not affected by transfer faults along the Iberia-Newfoundland margins 590 (Fig. 6a). Thus, in this case, the structures observed in the continent do not control the 591 segmentation in the adjacent oceanic crust. Furthermore, none of the major transfer 592 structures along the southern North Atlantic, with the exception of the Torre Madeira -593

Nova Scotia and Charlie Gibbs fracture zones seem to penetrate into the hyperextended 594 domains. This also suggests that inherited continental structures are unlikely to control the 595 structure of the adjacent oceanic crust.

596

Mapping of rift domains in the Bay of Biscay -Pyrenean system enabled the 597 description of inherited pre-and syn-rift structures and their relationships to the first 598 oceanic crust. Jammes, (2009) and Jammes et al. (2010b) showed that a network of 599 inherited structures such as the Toulouse and Pamplona faults localized and segmented 600 hyperextended rift basins. According to potential field maps, both the Toulouse and 601

Pamplona faults root deeply in the crust/lithosphere (Jammes et al. 2010b). A detailed 602 map of the rift domains (Tugend et al. 14) (Fig. 6b) shows the complex relationships 603 between the inherited Variscan and Late-to post-Variscan structures and the evolution of 604 the rift domain in space and time. On the large-scale maps (Fig. 6a) as well as on a 605 smaller-scale one inherited structures in the continent seem not to control the structures in 606 the adjacent ocean. This is well shown by the observation that the limit between the 607 hyperextended and oceanic domains remains unaffected by inherited structures mapped in 608 the continent. However, The observation that major structures observed in the continent can be prolonged 616 into the zone of exhumed mantle (e.g. Iberia Abyssal Plain; Alves, 2011) supports the 617 idea that the mantle underneath the OCT has a subcontinental origin. However, in the case 618 of the Iberian margin, these structures do not control the segmentation in the adjacent 619 oceanic crust. Furthermore, the major transfer structures in the oceanic crust along the 620 southern North Atlantic do not penetrate into the continent, suggesting that inherited 621 continental structures do not control the structure of the adjacent oceanic crust. The 622 observation that the first oceanic crust truncates earlier rift structures in the southern 623

North Atlantic (Fig. 6) suggests that breakup and seafloor spreading are not inflenced by 624 inheritance but are the result of deep-seated asthenospheric processes. The excellent exposure and preservation of the Err detachment system (Fig. 7a) 656 permit the determination of the relationships between inherited late Variscan (Permian) 657 structures and younger Alpine thrust faults (green and red structures in Fig. 7). Over wide 658 areas the Err detachment system corresponds to two well-defined, top-to-the-west 659 surfaces, referred to as the Err and the Jenatsch detachment faults (dark and light blue 660 structures in Fig. 7). These detachment surfaces are either overlain by extensional 661 allochthons, or, where the faults were directly exhumed at the seafloor, by syn-rift 662 sediments (Fig. 7c; see also [START_REF] Masini | Anatomy and tectono-1552 sedimentary evolution of a rift-related detachment system: The example of the Err 1553 detachment (central Alps, SE Switzerland)[END_REF] for details).

663

A first observation establishes the relation between the detachment system and a 664 late Variscan batholith made mainly of granite. Throughout most of the observed area, the 665 granites form the footwall of the detachment system, while they are nearly absent in the 666 hanging wall. The primary intrusive contact between the host rocks and the granites, as 667 well as the occurrence of host rocks within the granites, can be observed in many places 668 in the footwall, suggesting that the detachment localized along, or near, the roof of the 669 batholith. A second key observation is the relationship between the detachment system 670 and a fault bounding a former Permian basin filled with a >500m thick volcano-671 sedimentary sequence of Permian age and sealed by lower Triassic shallow marine 672 dolomites (Fig. 7b). As shown in Fig. 7c, the Jurassic detachment is overlain by the thick 673

Permo-Triassic succession in the area of Piz Surgonda. Directly to the North, the lower 674

Triassic dolomites are directly overlain, along a primary contact, by micaschists and 675 orthogneisses that are in the hanging wall of the Jurassic detachment (Fig. 7c). This 676 observation suggests that the Permian basin terminated abruptly north of the Surgonda 677 area along an East-West striking steep, south-dipping fault forming the northern 678 termination of the Permian basin (see Fig. 7d). Detailed mapping of the detachment 679 surfaces shows that the northern termination of the Permian basin corresponds to the 680 location where the Jurassic detachment fault system forms an East-West trending lateral 681 ramp. Thus, the lateral ramp of the detachment fault reactivates an inherited surface 682 corresponding to the northern termination of the Permian basin (see Fig. 7d). Mapping of 683 the major top-to-the-west Alpine thrust fault shows that it reactivates the Jurassic 684 detachment system in the south. In the north, where the Jurassic detachment fault forms a 685 lateral ramp, the Alpine thrust truncates this structure. This observation suggests that the 686

Alpine thrust was unable to reactivate the lateral ramp and incised into the footwall of the 687 Jurassic detachment. This can explain the observation that the Jurassic detachment system 688 was not reactivated during Alpine convergence and preserved its former relationship 689 between footwall and hanging wall in the vicinity of the lateral ramp. 690 691 Implications 692

The mapping of the well preserved Permian, Jurassic and Alpine structures in the 693 area of Piz Err -Piz Surgonda shows that the detachment faults localized either at 694 existing lithological contacts such as the limit between intrusive granites and the 695 surrounding host rocks, or that the detachment reactivated pre-existing structures. In the 696 example of Err detachment system, the complex 3D structure of the detachment system is 697 strongly controlled by the existence of a Permian basin as shown by the section that is 698 perpendicular to transport direction in Fig. 7c. Cretaceous sediments show that salt had to be in direct contact with the exhumed mantle 735 already before Pyrenean convergence. An additional and independent argument is that a 736 hydrothermal metamorphic event, associated with mantle exhumation, and dated at 107+/-737 3Ma, affected Triassic evaporites [START_REF] Thiébaut | Métamorphisme et halocinèse crétacés dans les 1656 évaporites de Betchat le long du Chevauchement Frontal Nord-Pyrénéen (Haute 1657 Garonne et Ariège, France)[END_REF][START_REF] Thiébaut | Métamorphisme des évaporites triasiques du versant nord des Pyrénées centrales et 1661 occidentales[END_REF]. All these observations 738

show that there is a close relationship between salt, detachment faults and mantle 739 exhumation. 740 741

In the Parentis basin, diapiric structures occur mainly in the southern part (Mathieu, 742 1986;[START_REF] Biteau | The Aquitaine Basin[END_REF]. Salt migration initiated in mid Cretaceous time and lasted until 743 the end of Eocene time. The relation between the salt and the rift structures can be 744 observed within a section running parallel to the coastline, referred to as the "Aquitaine 745 coast section" (Fig. 8c). In this section, Jammes et al. (2010a) described along the southern 746 part of the section an extensional detachment structure that is related to salt migration.

747

Apart from the unmistakable evidence for salt diapirs (supported by drilling), there are 748 some hints for asymmetric blocks bounded by growth structures that are underlain by a 749 continuous, strong sub-salt reflector (Fig. 8c). The overall structure resembles that of 750 extensional allochthons. These structures look also similar to those observed in the 751

Arbailles, where tilted, fault-bounded blocks with Urgonian growth structures are 752 associated with salt diapirs [START_REF] Canérot | Rifting eocrétacé et halocinèse sur la marge ibérique des Pyrénées 1442 Occidentale (France). Conséquences structurales[END_REF] the supra-salt units from deformation in the sub-salt units. While the supra-salt 769 deformation was mainly accommodated by block tilting, the sub-salt deformation was 770 accommodated along a top basement detachment fault drilled at wells SGM1 and CTS1 771 (Fig. 8c). We propose therefore that the tilting of the Upper Triassic to Lower Cretaceous 772 carbonates, associated with the growth-structures (probably of Aptian age) occurred 773 simultaneously with, but decoupled from major crustal thinning in the underlying 774 continental crust (see model in Implications: 781

The study of Triassic evaporite layers and their relationships to younger 782 extensional detachment faults exhuming deeper crustal and mantle rocks in the Parentis 783

and Mauléon basins enables us to suggest that the presence of a pre-rift salt layer in an 784 area undergoing extreme crustal thinning can control the geometry and evolution of rift 785 systems and obscure the rift related structures in the underlying basement. During an 786 initial stage of rifting, pre-rift salt layers act as a decoupling horizon between sub-and 787 supra-salt units suppressing the extension of the crustal detachment faults into layers 788 above the salt and, by doing so, eliminating breakaways at the seafloor (Fig. 8d). As a 789 consequence, the detachments sole out along the ductile salt layer and no sub-salt material 790 can be exhumed to the seafloor. In his pioneering work, [START_REF] Buck | Modes of continental lithospheric extension[END_REF] showed that the thermal structure of the 804 lithosphere is controlling, together with the composition and crustal thickness, the 805 evolution of rift systems. Depending on the thermal structure and crustal thickness, Buck 806 (1991) distinguished between "narrow", "wide" or "core complex" extensional systems. 807

Analogue models [START_REF] Brun | Mantle exhumation at passive margins[END_REF] and numerical models (Huismans and 808 Beaumont, 2007;Jammes et al., 2010) showed that, apart from the thermal structure and 809 composition of the lithosphere, also the strain rate and obliquity may control the 810 architecture of rift systems. The aim of this modeling study is to investigate the effect of 811 the thermal structure on the dynamic behavior of the lithosphere during hyperextension. 812

We consciously limit our models to very simple/basic to better investigate the very impact 813 of different thermal structures. In contrast to compositional and structural inheritance, 814 thermal inheritance is transient, therefore it would be difficult to investigate the effects of 815 the thermal inheritance on hyperextension using fossil examples.

816

Following the work of Lavier and Manataschal (2006) we test the effects, on 817 extensional styles and evolution, of variations in initial thermal structure in a 818 rheologically layered lithosphere including a mafic (gabbroic) lower crust (Fig. 9) 819 assumed to be inherited from a phase of post orogenic collapse (see paragraph 3.1.1.2). 820

The mantle is modelled as dry olivine and is assumed to be strong and depleted. We also 821 include weakening of the crust by hydration where fault zones or ductile shear zone are 822 active (for details see Fig. 9c) [START_REF] Lavier | A mechanism to thin the continental lithosphere at 1524 magma-poor margins[END_REF]. Finally, to limit the size of 823 the parameter space we assume that the initial crustal thickness is dependent on the initial 824 thermal structure. The numerical domain is initially 450 km wide and 100 km thick. (Fig. 825 9a). The crustal thickness is 35 km including a 7 km thick mafic layer at its base. Both 826 numerical techniques and boundary conditions are identical to those used in Lavier and 827 Manatschal (2006). To simulate rifting an extensional half rate of 2.5 mm/yr is applied on 828 both side of the model (Fig. 9a). We vary the thermal structure assuming that the 829 lithosphere cools rapidly after the phase of post-orogenic collapse. The lithospheric 830 thickness is fixed to 100 km and we calculate an initial geotherm based on a fixed value 831 of temperature at the Moho given by geological observation (i.e. [START_REF] Hansmann | U-Pb zircon geochronology of a 1489 tholeiitic intrusion and associated migmatites at a continental crust-mantle 1490 transition, Val Malenco, Italy[END_REF]).

832

We define three cases: (1) the Moho temperature is set at 450°C for an old lithosphere 833 (500 Myr) (Fig. 10), (2) 100 Myr later at 600°C (Fig. 11) and (3) immediately after post 834 orogenic collapse at 750°C (Fig. 12). The first case could correspond to an ancient mobile 835 belt and the intermediate case to the Tethyan rifting similar to what is observed in the 836 Alps (for details see [START_REF] Manatschal | New models for evolution of magma-poor rifted margins based on 1541 a review of data and concepts from West Iberia and the Alps[END_REF]). Even though these models are relatively simple 837 when compared with the complex inheritances that could be ascribed (see previous 838 sections), we obtain a wide range of tectonic behaviours and evolutions. 10, 11 & 12). We distinguish 3 different styles of 844 deformation: (see captions for details): (1) localized extension over a symmetric graben 845 that roots in the mantle lithosphere that leads to rapid thinning and necking of the 846 lithosphere until the mantle is exhumed at the surface (Fig. 10), (2) distributed extension 847 followed by a phase of thinning during which necking of the lithosphere occurs for an 848 intermediate lithosphere (Fig. 11), (3) highly distributed extension and near pure shear 849 deformation in case 1 with a hot/young lithosphere (Fig. 12). At first order the main 850 control on the different styles of extension is whether brittle deformation in the upper 851 crust is coupled to brittle deformation in the mafic lower crust and uppermost lithospheric 852 mantle. In case 3 the hot ductile lower crust allows for decoupling to be maintained over 853 very long time scales (30 myrs). In case 2 after being decoupled for ~ 13 Myr the middle 854 to lower crust eventually are thin enough that brittle upper crustal deformation can couple 855 to the deformation in the uppermost embrittled lithospheric mantle. At this moment in 856 time the lithosphere thins over a major lithospheric scale fault zone extending in the 857 mantle. In case 1 the lower crust and mantle lithosphere are mostly brittle and the 858 deformation is immediately coupled, leading to very localized rifts rooted in the mantle 859 lithosphere. The initial rapid subsidence of this keystone block, defined as block H in 860 [START_REF] Lavier | A mechanism to thin the continental lithosphere at 1524 magma-poor margins[END_REF], is followed by rapid necking of the lithosphere and 861 extension and thinning of the internal parts of block H. Case 3 is an analogue of the 862 stretching phase and case 1 of the thinning phase defined by Lavier and Manatschal 863 (2006). Weakening of the crust by hydration [START_REF] Lavier | A mechanism to thin the continental lithosphere at 1524 magma-poor margins[END_REF]) develops 864 listric detachment faults that distribute the deformation in the upper crust. It also allows 865

for the formation of a weak middle crustal layer that decouples brittle deformation 866 between the upper and lower gabbroic crust. Finally the strong gabbroic lower crust 867 formed during the previous phase of extensional collapse act as a guide that distribute the 868 stress elastically or viscoelastically across the whole extensional domain. It becomes a 869 stress concentrator when brittle deformation can localize and propagate through it by 870 plastic (Mohr Coulomb yield) weakening as a function of plastic strain. Figure 13 shows 871 the state of the lithosphere and its topography after 150 km of extension for each case. At 872 first order case 1 develops a narrower conjugate margins with a large zone of exhumed 873 continental mantle. Case 3 develops upper-lower plate asymmetric conjugate margins in 874 a wider rift but a narrow zone of final thinning. This model does not reach exhumation 875 after 150km of extension. Case 2 is similar to that described in Lavier and Manatschal 876 (2006). The extension has stretched the lithosphere down to 20 km thickness (stretching 877 phase) before thinning it down to 10 km during the thinning phase preceding the 878 exhumation phase that brings some serpentinized mantle to the surface along a rolling-879 hinge exhumation fault. 880 881 882

Implications 883

These models confirm that the thermal structure is a key parameter in controlling 884 the rift evolution [START_REF] Buck | Modes of continental lithospheric extension[END_REF]. The results of the models shown in Figures 10, 11 and 12 885 also emphasize the importance of the thermal structure controlling the necking 886 architecture of the rifted margin. For higher temperature, a larger fraction of the upper 887 lithospheric mantle and the continental crust are ductile. This results in the delocalization 888 of the deformation and makes it hard to develop an area of necking where thinning of the 889 crust and mantle are coupled and Moho-topography develops (Fig 12). In contrast, 890 thinning of the crust and mantle are easily coupled and form sharper necks in colder 891 settings (Fig 10). Thus, the evolution of a rift system depends on the thermal structure of 892 the lithosphere in addition to its composition, rates and obliquity. For the same crustal 893 thickness or tectonic province, young and hot lithosphere may be difficult to thin resulting 894 in wide necking profiles, while old and cold lithosphere (Archean or Paleoproterozoic) 895 may be strong, resulting in narrow necking profiles. The Variscan lithosphere underlying 896 the Alpine Tethys and Iberia Atlantic margins follows a phase of Permian orogenic 897 collapse that introduces a large amount of heat by magmatic processes and lithospheric 898 thinning about 70 myr before onset of rifting. It may still be quite hot and may represent 899 an intermediate case between very young and very old lithosphere. Of course, these ideas 900 do neither consider compositional variations, nor strain rate, which may strongly modify 901 the bulk rheology of the extending lithosphere (see paragraph 3.2.1 and Fig. 5). At the 902 moment, we do not see how we can overcome the problem of constraining complex 903 structural and compositional inheritance, since we are not yet able to place realistic 904 "structural" or "compositional" inheritances in model and allow them to compete with the 905 thermal one. However, already taking into account variations in the thermal structure may 906 explain why most rift systems, except for those that are related to subductions (e.g. back-907 arc situations; Tyrrhenian sea; Woodlark, Gulf of California), can not develop directly 908 from an orogenic collapse into a rift system (e.g. Basin and Range). It is likely that rift 909 systems cannot start within a crust/lithosphere that is too hot, probably because thinning 910 of the crust and the mantle cannot be coupled. In contrast, it may also be difficult to 911 initiate a rift system in a very cold lithosphere. This may explain why many rift systems 912 try, in an early stage, to avoid these domains and may prefer weak orogenic lithosphere or 913 other types of weaknesses. 914 915 5. Discussion 916

Relative role of inheritance vs. rift-induced processes 917

The integration of inheritance in rift models is a key to eventually be able to 918 predict and understand the evolution of rift systems. However, as compositional, 919 structural and thermal inheritances cannot be disassociated one from each other, it 920 remains at present difficult to link particular rift structures or their evolution directly to 921 one specific type of inheritance. Furthermore, existing structures or weaknesses are not 922 always entirely reactivated as exemplified by the relation between Jurassic extensional 923 detachment fault and Alpine thrust in the vicinity of the a lateral ramp in paragraph 3.3.1, 924 (Fig. 7). Reasons may be that the geometry of the pre-existing structures/weaknesses are 925 complex (e.g. listric faults, ramps etc.), limited in their lateral extend (e.g. confined 926 evaporite layers), their orientation was not favourable to be reactivated, or the inherited 927 weakness is located in a layer not stiff enough to concentrate the stress (e.g. Chenin and 928 Beaumont, 2013). Another important point is that inherited weaknesses can either interact 929

or not depending on their scale and location within the lithosphere. • conversely, the deformation in the uppermost parts of the crust and exhumed mantle in 945 hyperextended, coupled domains seems not to be controlled by inheritance, but by 946 newly formed weak materials (clay, serpentinite, ...) 947

• mantle composition and mantle processes occurring before and during rifting may 948 control the magmatic budget, the thermal structure and the localization of final rifting 949 (e.g. Alps, Pyrenees and Iberia examples; Fig. 4) in addition to the rheology of the 950 mantle 951

• inherited structures within the crust do not significantly control the location of breakup 952 and the structures in the oceanic crust (e.g. Bay of Biscay and Iberian Atlantic; Fig. 6) 953

• the inherited thermal structure controls the mode and architecture of rift systems, in 954 particular the necking architecture (e.g. modelling approach; Figs. 10, 11, 12) 955 956

These key observations enable us to introduce inheritance in the rift models proposed 957 to describe the evolution of rifted margins. However, it is important to distinguish 958 between inheritance, and rift-induced processes, both of which can control rifting.

959

Examples of rift-induced processes include formation of clay, serpentinization or 960 decompression melting. It is also important to note that the role of inheritance may change 961 as the physical conditions vary during the evolving rifting and as rift-induced processes 962 become more important. Inherited mantle composition may become particularly important 963 when the thermal structure of rifting evolves and magma production becomes an 964 important factor. 965 966

The role of inheritance within an evolving rift system 967 968

Modes of rifting 969 [START_REF] Lavier | A mechanism to thin the continental lithosphere at 1524 magma-poor margins[END_REF] proposed a dynamic model for magma-poor rift 970 systems, for which they defined different modes of rifting depending on the deformation 971 process(es) at play. These modes are referred to as stretching, thinning and exhumation 972 (Fig. 14). As discussed by [START_REF] Bronner | Magmatic breakup as an explanation for magnetic anomalies at magma-poor rifted 1429 margins[END_REF], breakup at magma-poor rift systems may 973 only occur if an excess magmatic event is able to break the strength of the lithosphere and 974 to trigger steady state seafloor spreading. In order to capture the breakup process, we 975 define a "magmatic mode" pre-dating the "seafloor-spreading mode" (onset of steady-976 state seafloor spreading).

977

Independently of their inheritance, margins may go through a sequence of 978 stretching, thinning, exhumation (not mandatory) and magmatic breakup. Although these 979 modes have been defined for the Iberia-Newfoundland and Alpine Tethyan margins 980 [START_REF] Sutra | Quantification and restoration 1650 of extensional deformation along the Western Iberia and Newfoundland rifted 1651 margins[END_REF], Mohn et al. 2013), they can be also recognized in most of the world 981 rifted margins, their importance and expression may change depending on the inheritance.

982

Therefore understanding which mode is controlled by what type of inheritance is a key to 983 describe and predict the evolution of rift systems (Fig. 14).

985

Stretching mode 986

During the stretching mode the total horizontal extension and the related crustal 987 and lithospheric thinning are minor, thus the thermal perturbation and the magmatic 988 production due to decompression melting are negligible. Therefore, in the absence of 989 magma (plume), the strength of the lithosphere is controlled by the strongest brittle layer, 990 which plays the role of a stress guider (see [START_REF] Lister | The origin of metamorphic core complexes and 1531 detachment faults formed during Tertiary continental extension in the northern 1532 Colorado River region[END_REF] on a larger scale it is not excluded that the newly formed weak material favours pure 1046 shear deformation to occur simultaneous to exhumation along localized detachment faults. 1047 Indeed, hydration of the crustal rocks (mainly feldspar to illite and chlorite) and mantle 1048 rocks (in particular olivine to serpentine) results in the formation of low friction mineral 1049 that may change in a significant way the rheology of the extending uppermost crust and 1050 exhumed mantle. It is important to note that these processes are a consequence of rifting 1051 rather than controlled by inheritance. This control by the change in mineralogy may be 1052 illustrated by the width of the hyperextended crustal wedge along the Iberian margin, 1053 which remains the same along the whole margin while the necking profiles vary strongly 1054 (e.g. Fig. 5c , 2013[START_REF] Direen | Architecture of volcanic 1458 rifted margins: new insights from the Exmouth -Gascoyne margin, Western 1459 Australia[END_REF]. Indeed, at all these margins, the final magmatic 1094 breakup overprinted older hyperextended domains, suggesting that the breakup was not 1095 related to the previous extensional event, but triggered by the excess of magma. Similar 1096 observations can be made in other margins, where the final breakup crosscuts previous rift 1097 structures (see southern North Atlantic, [START_REF] Lundin | Hyperextension, serpentinization, and weakening: A new 1534 paradigm for rifted margin compressional deformation[END_REF]. This begs the question 1098 of how much magmatic processes related to continental breakup are controlled by 1099

  Fig. 8d). The interaction of detachment faults and 775 decollements, the latter within the Keuper evaporites, can explain, as shown by Jammes et 776 al. (2010a), the unconformities observed and drilled in the southern Parentis and Mauléon 777 basins as well as the extreme thinning of the crust and local mantle exhumation observed 778 in these basins. 779 780

  and 12 show the results of the 3 different cases (1 to 3) with 842 variable Moho temperatures. The figures show the evolution of the viscosity and strain 843 rate over 30 Myr for each case (Fig.

1169Fig. 1 :

 1 Lithospheric scale sections showing (a) an idealized lithosphere made of a layer 1170 cake and thermally equilibrated) and (b) a "real" post-orogenic lithosphere showing 1171 inherited structural and compositional complexity. 1172 1173 1244 Fig. 6: (a) Map of the southern N-Atlantic showing the major structures such as oceanic 1245 fracture zones, OCT, necking zones, continental fracture zones and the magnetic 1246 anomalies 34 and J that are supposed to correspond to the continent ward limit of 1247 unequivocal oceanic crust and embryonic oceanic crust respectively. Maps compiled after 1248 Arenas et al. (2007), Ziegler and Dézes (2006), Mosar (2003) and Ryan et al. (2009). (b) 1249 Map of rift domains preserved in the Bay of Biscay and Western Pyrenees ; CIZ: Central 1250 Iberian zone. WALZ: West Asturian-Leonese zone. NPF: North Pyrenean fault (modified 1251 after Tugend et al. subm). 1252 1253 1254 Fig. 7: (a) Panoramic view of the Jurassic extensional Err and Jenatsch detachment faults 1255 exposed in the Err nappe in SE Switzerland. View from west towards the East. (b) Map of 1256 the Err nappe in the area of Piz Err -Piz Surgonda, showing relations between the 1257 Jurassic extensional detachment system (in blue and red colours), a pre-existing Permian 1258 normal fault (in green colour) and an Alpine thrust fault (in black). (c) NW-SE cross 1259 section across the Err nappe showing the present-day relation between the Jurassic 1260 extensional detachment faults, the Permian fault and the Alpine D1 thrust fault. (d) A 1261 restored section showing the relationships before onset of Jurassic rifting and the traces of 1262 the future faults. 1263 1264 1310 Fig. 10: The thermal structure is set such that the Moho temperature is 750ºC and the 1311 bottom temperature is 1330ºC. This situation is the case of a hot lithosphere 1312 corresponding to conditions similar to that past orogenic collapse. The lower and middle 1313 crusts are very weak and the mantle is mostly ductile. A) Evolution of the viscosity over 1314 100 km or 20 Myr of extension. The lithosphere has a low average viscosity (~10 23 Pa.s). 1315 The deformation is mostly viscous and crust and mantle deformation are decoupled for 1316 most of the evolution. Three small basins form and accommodated part of the extension in 1317 a brittle manner. As more extension is accommodated a shallow layer of hydrated middle 1318 crust forms and link the different basin. The crust and mantle thin uniformly while the 1319 middle crust is slowly exhumed close to the surface. After 75 km of extension the crust is 1320 uniformly thinned down to ~20 km. After 100 km, extension localizes in a very narrow 1321 basin or keystone block (near the distance 190 km), the deformation is still very 1322 distributed and the weak middle crust is exhumed on the flanks of the keystone. B) Until 1323 50 km of extension, the deformation is distributed in the middle to lower crust and mantle 1324 lithosphere with some normal faulting in the crust. After this period of ductile and 1325 distributed thinning, a large number of normal faults form across the zone of extension. 1326 This marks a period of embrittlement of the lithosphere that leads to the localization of the 1327 extension in one narrow keystone block (100 km) of extension. The thermal structure is set such that the Moho temperature is 600ºC and the 1332 bottom temperature is 1330ºC. This situation is the case of a hot lithosphere 1333 corresponding to conditions of a young basin (100 Myr past orogenic collapse). A) 1334 Evolution of the viscosity over 100 km or 20 Myr of extension at 5 mm.yr -1 . In that case a 1335 keystone block H forms between 50 and 75 km of extension. This follows a period of 1336 pure shear distributed extension in the upper crust accompanied by the formation of mid-1337 crustal layer of hydrated crust. The deformation in the crust and the mantle is decoupled 1338 until the keystone forms. The keystone is narrower (~40 km wide) than in the previous 1339 case. After 100 km of extension the crust is thinned to < 100 km and the mantle 1340 lithosphere upwelling in a narrow channel below the crust. B) Deformation is distributed 1341 in the upper crust from the onset of deformation suggesting that a large proportion of the 1342 lithosphere can behave in a brittle manner. After 25 km of extension multiple half grabens 1343 have formed in the upper and the mantle lithosphere starts to deform in a brittle manner 1344 along multiple normal shear zone. After 50 km of extension the deformation zone 1345 narrows until it localizes in a keystone block (75 km of extension). Between 75 and 100 1346 km of extension the crust and lithosphere thin in the keystone block (see previous case). The thermal structure is set such that the Moho temperature is 450ºC and the 1351 bottom temperature is 1330ºC. This situation is the case of a cold lithosphere. These are 1352 conditions in a very old orogeny (500 Myr) and a very strong lithosphere. The lower and 1353 middle crusts are very strong and brittle and the mantle lithosphere is brittle down to 50 1354 km. A) Evolution of the viscosity over 100 km or 20 Myr of extension at 5 mm.yr -1 . The 1355 models show localization of the deformation in a 90 km wide keystone (initial block H) 1356 that evolves in a very localized and narrow set of conjugate margins (100 km of 1357 extension). Zones of weak hydrated crust (outlined in with black lines in the middle crust) 1358 form at the base of the brittle shear zone in the crust. The rift flanks are very developed 1359 1372 Fig. 13: Final structure and topography of the lithosphere for each case after 150 km of 1373 extension. In all cases a localized thinned zone of crust forms. In the coldest case (Moho 1374 at 450ºC) the final basin is wide (~230 km) with a large zone of exhumed serpentinized 1375 mantle and a very pronounced topography (max ~10 km). For the hottest case (Moho at 1376 750ºC) the basin is narrower (<100 km) and the topography is very asymmetric. The 1377 margin has developed a very clear upper-lower plate structure. Moreover the hydrated 1378 middle crust is exhumed close to the surface. Finally the medium case (Moho at 600ºC) 1379 is akin to that of the Lavier and Manatschal (2006) case. The basin is asymmetric and 1380 some exhumation has occurred and the basin width is intermediate after the thinning 1381 phase (~160 km). 1382 1383 1384 Fig. 14: Conceptual model showing the tectonic, rheological and magmatic and thermal 1385 evolution of a magma-poor rifted margins and its relation to inheritance and rift induced 1386 processes. 1387 1388

The role of Variscan inheritance in structuring the Alpine Tethys and Iberian rift- 271

  or the Southern-North Atlantic (Péron-Pinvidic andManatschal, 256 2010). Why rift systems may fail or migrate and how far the localization/delocalization of 257 rift systems may be controlled by inheritance will be part of the discussion of this paper. Along strike variations, segmentation and 3D evolution of rift systems 260Most studies and models of rifted margins are focused on a 2D approach looking 261 at dip sections parallel to the kinematic transport direction. It is important to note that 262 these sections are not necessarily perpendicular to the margin, in particular in oblique or 263 shear margins. The along strike architecture of rifted margins characterized by 264 segmentation and strong lateral changes of crustal structures and/or the volume of 265 magmatic additions is not well understood. The hypothesis that inheritance and/or rift-266 induced processes may control the lateral evolution and segmentation of rift systems is 267 discussed further below, based on the example of the southern North Atlantic.

	258
	259
	2.2.3 268
	269
	270
	3. 284
	285
	3.1

systems 272

Here we use the term "Variscan inheritance" to refer not only to the orogenic 273 processes but, also to the late to post-orogenic events, distinguishing between crustal and 274 mantle inheritance. We focus on the Alpine Tethys and Iberian Atlantic rift systems that 275 formed over a lithosphere with a strong Variscan inheritance. Although the fate of these 276 two conjugate margins are different, i.e. one became part of a big ocean while the other 277 probably never developed into a steady state ocean, we consider the two systems as 278 belonging to the same Late Triassic to Early Cretaceous rift system that formed over 279 Variscan orogenic lithosphere. Until now, most discussions of the importance of 280 inheritance focused either on local or large-scale examples, without considering that the 281 resolution of the observations is strongly scale-dependent. Here we try to integrate this 282 idea by discussing the role of structural and compositional inheritance at different scales, 283 ranging from lithospheric-to outcrop-scale.

Influence of inheritance on a large-scale: the example of "Variscan" inheritance

	286
	287
	288	3.1.1 Variscan processes controlling crustal inheritance
	289
	290	3.1.1.1 Pre-and syn-Variscan structures

2 Control of orogenic inheritance on a regional-scale: crustal necking and rifted

	480	subcontinental
	481	mantle is heterogeneous.
	482
	3. 483
	484	margin segmentation
	485
	486	3.2.1 Control of inheritance vs. rift-induced processes on the morphology of necking
	487	profiles
	488	The Western Iberia margin is often considered as the archetype magma-poor
	489	hyperextended rifted margin (Péron-Pinvidic et al. 2013). The numerous studies
	490	performed along this rifted margin (e.g. (Tucholke et al., 2007), Péron-Pinvidic and
	491	Manatschal, 2009) benefited from the dense covering of refraction and reflection seismic
	492	and drill hole data. In particular the discovery of hyperextended continental crust begged
	493	the questions of how the crust thins and what controls extreme crustal thinning? In order
	494	to answer to these questions, Sutra and Manatschal (2012) defined and described necking
	495	zones using reflection and refraction seismic lines from the Iberia rifted margin. These
	496	necking zones correspond to areas where the crust is thinned from its initial thickness to a
	497	crustal thickness of < 10km.
	498
	499	Observations

467 Müntener et al. (2004)

suggested, using the examples of the Alps that magmas first 468 reacted with the previously depleted lithospheric mantle before they were extracted 469

[START_REF] Manatschal | A type sequence across an ancient magma-poor 1544 ocean continent transition: the example of the western Alpine Tethys ophiolites[END_REF]

. Assuming that these observations are generally valid, it would 470 mean that the subcontinental mantle was depleted during orogenic collapse and was 471 unable to produce a lot of magma during the following rifting. The first magma that was 472 formed by decompression reacted with the overlying depleted lithospheric mantle leading 473 to its fertilisation without any surface expression. This suggests that magma-poor rifted 474 margins may develop in the absence of plumes over a lithosphere that was depleted 475 during orogenic collapse. In contrast, magma-rich margins may occur over fertile mantle 476 lithosphere, either because it has never been depleted or because it was fertilized by a 477 plume or during previous rift-events. How important and how general these processes are, 478 is difficult to determine. However, it shows that linking ß-factors to volume of produced 479 magma is an oversimplification and may not apply to domains where the

  [START_REF] Tugend | Formation and deformation of hyperextended rift systems: insights from rift domain 1680 mapping in the Bay of Biscay -Pyrenees[END_REF] andJammes et al. (2010b) showed that 609 structuration of the hyperextended domain is strongly controlled by inheritance. Inherited 610 structures can either transfer extension (e.g. Pamplona fault), or control the polarity of 611 hyperextended margins, separating upper from lower plated margins (see Tugend et al.

	612
	613	2014).
	614
	615	Implications:

Control of inheritance on a local-scale 632

  In this section we use two examples to discuss how inherited structures and compositional 633 variations may control strain distribution on a local scale during hyperextension. The key 634 learning that we want to convey is the importance of inherited weak layers controlling 635 strain distribution and style of rifting. While the existence of pre-existing faults in the 636 brittle layer may be a common feature on rifted margins, the pre-rift salt example can be 637 used as an analogue for weaknesses such as weak layers in the crust or in the mantle able

	650	faulting and exhumation. The detailed mapping of this rift related detachment system
	651	(Froitzheim and Eberli, 1990; Manatschal and Nievergelt 1997; Masini et al. 2012) shows
	652	the importance of the pre-rift structures for the localization of this long offset low-angle
	653	extensional detachment structures.
	654
	655	Observations
	625	However, at
	626	present the process of breakup is not yet sufficiently constrained to understand the
	627	underlying controlling factors. The study of the Bay of Biscay -Pyrenean system shows
	628	that stretching, thinning and exhumation are strongly controlled by inherited structures,
	629	contrary to breakup and seafloor spreading. The latter may be controlled by heterogeneity
	630	within the lithospheric mantle underlying the hyperextended domains (e.g. section 3.1.2).
	631
	3.3 638
	639	to localize deformation.
	640
	641	3.3.1 Relation between pre-rift structures and detachment faulting (Err detachment
	642	system, SE Switzerland)
	643
	644	General situation
		The Err detachment system exposed in the Lower Austroalpine Err nappe in SE

645

Switzerland is one of the world's few exposed and preserved rift-related detachment 646 systems (e.g.,

[START_REF] Masini | The Alpine Tethys rifted margins: 1556 Reconciling old and new ideas to understand the stratigraphic architecture of 1557 magma-poor rifted margins[END_REF]

. This structure can be mapped over about 200 647 km 2 exhibiting primary relationships with pre-rift Permian structures and Alpine thrust 648 faults. The Err detachment system formed during the final stage of rifting by sequential 649

  These observations show that reactivation 699 of inherited structures or lithological boundaries with a major rheological contrast are 700 important and can control, on a local scale, the structures in hyperextended systems.

	701
	702
	703
	704	3.3.2 Relationships between pre-rift salt and detachment faulting (Mauléon and Parentis
	705	basins)
	706
	707	General situation
	708	The Mauléon and Parentis basins in the Western Pyrenees -Bay of Biscay formed
	709	as hyperextended basins during Late Aptian to Albian time (Fig. 8a). Within these basins
	710	salt structures have been drilled and are seismically imaged (e.g. Ibis glacier in Parentis
	711	basin: Mathieu, (1986); Biteau et al. (2006)) and observed in the field (e.g. massif des
	712	Arbailles, Roquiague diapir in Arzacq-Mauléon basin; Canérot (1989)) (Fig. 8b). These
	713	basins offer therefore the opportunity to study the role of Triassic pre-rift evaporites
	714	during the subsequent Late Aptian to Albian hyperextension using seismic and drill hole
	715	data and field observations. Jammes et al. (2010a) and Lagabrielle et al. (2010) showed
	716	that the pre-rift salt played an important role controlling the final structure of the
	717	hyperextended basins. We consider the pre-rift salt as inherited compositional weaknesses,
	718	similar to weaknesses in the crust and/or mantle, with the only difference that evaporites
	719	are better exposed and their initial geometry better understood.
	720
	721	Observations
	722	Within the Mauléon basin, the subsalt (basal Triassic Muschelkalk and Buntsandstein
	723	units) and the supra-salt units (Upper Triassic to Lower Cretaceous carbonates) are never
	724	juxtaposed within one continuous stratigraphic section. The sub-salt Lower Triassic
	725	sediments (Buntsandstein and Muschelkalk) are associated with the Paleozoic meta-
	726	sediments, while the supra-salt Upper Triassic to Lower Cretaceous carbonates form
	727	massive tilted blocks, interpreted as extensional allochthons that overlay Triassic
	728	evaporites and either continental or mantle rocks (e.g. Sarrance, Mail Arrouy blocks
	729	(Jammes et al., 2009). Thus, in the Mauléon basin, the Triassic evaporites served as a
	730	decoupling horizon separating sub-salt from supra-salt units. Although the present-day
	731	position of these blocks is the result of Pyrenean compression, the occurrence of growth
	732	structures of Early Aptian age (e.g. Urgonian carbonates in the Arbailles) indicates that
	733	these blocks tilted during Early Cretaceous rifting using the Triassic evaporites as
	734	decollement horizons. The observation of mantle rocks in salt diapirs within Upper

  see previous paragraph). About 2 km eastward 753 of the seismic section, two wells, SGM1 and CTS1, were drilled across the strong basal 754 reflector and penetrated into basement. Well SGM1 penetrated beneath Tertiary and 755 Upper Cretaceous sediments, 1737 m of Keuper salt series and 800 m of Triassic 756 Muschelkalk dolomite and Buntsandstein sandstones, and tectonized quartzite dated as 757 Carboniferous (Serrano et al., 2006). About 15 km to the north, located further basinward, 758 well CTS1 penetrates Tertiary and Upper Cretaceous sediments, 1189 m of salt, 500 m of 759 Lower Triassic sediments, and tectonized dolomites dated as Devonian (Serrano et al., 760 2006). The large thickness of the salt series penetrated by these two wells supports the 761 presence of diapiric structures. The strong continuous reflector identified in the seismic 762 data corresponds to the top of the Triassic Muschelkalk dolomite, which lies at the base of 763 the Keuper salt. The continuity of this reflector under tilted blocks suggests that a 764 decollement occurs in the overlying salt series, an interpretation similar to that proposed 765 for the Mauléon area (see previous paragraph). The fact that the top of the basement is 766 tectonized shows that this deformation is not only due to gravitational deformation. A 767 major decollement located in the evaporitic Keuper is required to decouple deformation in 768

  Onset of the thinning mode corresponds to the failure of the brittle layers in the 1002 lithosphere and terminates with the attenuation of the last ductile layer (Pérez-Gussinyé 1003 and[START_REF] Pérez-Gussinyé | Rheological evolution during extension at 1602 nonvolcanic rifted margins: Onset of serpentinization and development of 1603 detachments leading to continental breakup[END_REF]. Thus, the location of necking may be controlled by weaknesses in the 1004 strongest brittle layer. For low to moderate thermal gradients this may correspond to the 1005 upper mantle (see deep necking depths of[START_REF] Bassi | Contrasting styles of rifting: Models and examples 1402 from the Eastern Canadian Margin[END_REF]. Thus, inheritance within the 1006 strong upper mantle may be important and control the location of the necking zone and 1007 ultimately the area of breakup. The necking process is mainly controlled by the capacity 1008 of the crust to decouple from the mantle lithosphere, i.e. the relative abundance of ductile 1009 crust over brittle crust, enabling to partition the deformation between the brittle and 1010 ductile layers. In contrast to the stretching mode where deformation is mainly limited to 1011 the crust, deformation is localized at a lithospheric scale during the thinning mode. How 1012 deformation is partitioned, how much pure-shear thinning occurs in the ductile layers and 1013 what processes really control the thinning of the crust are still a matter of debate. At this 1014 stage of rifting, it is very important to understand how brittle and ductile layers are 1015 distributed within the crust and lithosphere. Compositional and structural inheritance may 1016 control onset of thinning, while magma and fluids may become more important during the 1017 subsequent hyperextension. Therefore knowing the bulk rheological structure of the crust 1018 is essential. This may either be controlled by weak phases (hydrated phases, quartz rich 1019 rocks or metasediments in the crust) or strong phases (mafic bodies or granulites). The 1020 underthrusting or accretion of sediments during subduction or the existence of multiple 1021 sutures during accretion may create strong inherited weaknesses that may control the 1022 crustal thinning process. In systems with either high geothermal gradients or high 1023 percentages of ductile layers coupling may never be accomplished and Moho topography 1024 remains subdued despite of high amounts of extension (e.g. Basin and Range Province; 1025Fig.10). The overall thermal state is mainly controlling the brittle ductile transition. This 1026 becomes important because it ultimately controls the relative amount of ductile vs. brittle 1027 material in a crustal section. The colder the thermal state the more brittle the crust and 1028 lithospheric mantle and the faster coupling between crust and mantle will occur (Fig.12).

	1000
	1001	Thinning mode
	1029
	1030	Conversely, the more ductile material the more extension is needed to reach the coupling
	1031	point (see Figs. 10, 11 and 12 as well as Fig. 5).
	1032
	1033
	1034	Exhumation mode
	1035	During the exhumation mode, all the remaining continental crust is made of
	1036	residual brittle layers, either upper curst, mafic lower crust, or former ductile crust that
	1037	cooled and became brittle during thinning. At this stage, active brittle faults facilitate fluid
	1038	transport and hydration of the crustal and mantle rocks, resulting in the serpentinization of
	1039	the uppermost parts of the subcontinental mantle. Intense hydration modifies the original
	1040	mineralogy controlling rifting (quartz, feldspar and olivine) to a new one controlled by
	991 1041	. Brittle deformation is clays and serpentine within the uppermost 6 kilometres of the basement. Thus, the
	992 1042	decoupled along ductile crustal layers or ductile shear zones allowing brittle layers to rheology of the remaining crust is strongly controlled by weak, hydrated material.
	993 1043	deform independently from each other. Thus, the evolution of rift basins in the stretching However, field observations show that, even during this late stage, local inheritance is
	994 1044	mode is mainly controlled by inheritance/weaknesses in the uppermost brittle layer, which important and the presence of inherited low-friction material (e.g. salt) or inherited
	995 1045	is usually the upper crust. Inherited compositional weaknesses (pre-existing basins or structures may control the architecture of detachment systems on a local scale. However,
	996	suture zones) or inherited structures (faults or shear zones) may control the localisation
	997	during these initial stages of rifting. If magmatic activity occurs at this stage (often related
	998	to plumes), magma can control the localization of deformation at this early stage of rifting
	999	through thermal weakening (see East African rift).

  Although magmatic activity may occur at low ß (< 1.5) in rift systems, mostly as a 1067 consequence of magma input from a deep mantle source (e.g. East African and Rio 1068Grande rifts), rift-related magmatism is usually a consequence of decompression melting 1069 related to extreme lithospheric thinning. The production of magma in rift settings is 1070 controlled by a number of processes and parameters, of which the thermal state, the rate 1071 of thinning, the shape of the base of the lithosphere, the content of hydrous phases in the 1072 lithospheric mantle and its composition. Composition and amount of water in the 1073 lithosphere (e.g. inherited depleted or fertile lithospheric mantle) are inherited and the

	1055	and discussion in chapter 3.2.1). This suggests that the processes controlling
	1056	hyperextension and exhumation are different from those controlling necking. We assume
	1057	that, within the coupled domain (< 10km thick), the material is formed by hydrated
	1058	residual brittle parts of the continental crust, while further outboard, where the crust is
	1059	thinned to less than 6km, it may overlie a serpentinized mantle behaving like a Mohr-
	1060	Coulomb material with a low friction coeofficient.
	1061	Thus, the architecture of the hyperextended domain may be primarily controlled
	1062	by newly formed clay, serpentine or talc resulting from hydration of the crust and/or
	1063	mantle after their embrittlement, rather than by inheritance.
	1064
	1065
	1066	Magmatic mode
	1074
	1075	thermal state allowing for magma formation is usually reached after the lithosphere has
	1076	been thinned by at least 50%.
	1077	In most rift models, mantle is considered to be homogenous and magma
	1078	production controlled by strain (ß). However, the study of mantle rocks derived from the
	1079	Alpine Tethys and Iberia-Newfoundland rift systems showed that the continental
	1080	lithosphere involved in rifting is strongly heterogeneous. Although it may be difficult to
	1081	determine the composition and history of the subcontinental mantle rocks, observations
	1082	tend to show that two processes are important in controlling the composition of the
	1083	lithosphere, which are: 1) magmatic processes associated with former subduction zones
	1084	(supra-subduction vs. subducted mantle); and 2) the post-orogenic thinning and melting
	1085	(cleaning) of the mantle. It is interesting to note that the Variscan orogenic collapse
	1086	resulted in strongly depleted subcontinental mantle in the Alps and in Iberia before the
	1087	onset of rifting. This subcontinental mantle was subsequently fertilized by the melt
	1088	produced during lithospheric thinning pre-dating lithospheric breakup. This "buffer
	1089	effect" may explain the delay in the major magmatic pulse during rifting at these two
	1090	magma-poor rifted margins. In this context it is also interesting to note that some of the
	1091	classical volcanic rifted margins (magma-rich margins) may have initiated as magma-
	1092	poor rift systems that underwent hyperextension prior to magmatic breakup. Examples are
	1093	the North Atlantic Vøre and Møring margins or the NW Shelf off NW Australia (Péron-
		Pinvidic et al.

1309