Keywords: modelling, retro-wedge, shortening distribution, Eastern Pyrenees, evolution

From rit system to doubly vergent orogen: An evolutionary model based on a case study of the Eastern Pyrenees and controlling factors from numerical models Abstract he doubly vergent nature of some natural orogens is classically understood as two opposing thrust wedges (pro and retro) that comply with critical taper theory. he evidence that retro-wedges and their associated basins behave diferently from their pro-wedge counterparts has been steadily increasing over the past few decades. However, what causes an orogen to become doubly vergent is currently not well understood. Nor is the relationship between the pro-and retro-wedge during the evolution of a doubly vergent orogen. It is the aim of this work to improve our understanding of: 1) how the pro-and retro-wedges relate to each other during the orogenic process, 2) what factors control the evolution of a doubly vergent orogen and 3) a possible link between the pro-and retro-wedge.

Answering these questions requires an improved knowledge of the evolution of a doubly vergent orogen. We focussed on the Eastern Pyrenees as a type example of a doubly vergent orogen, due to the large amount of available data. We performed a detailed tectonostratigraphic study of the retro-foreland of the Eastern Pyrenees (European plate), updating the interpretation based on recent insights into its hyperextended rit origins. We link the evolution of the retro-foreland to that of the pro-foreland (Iberian plate) in order to derive insight into the crustal scale dynamics. Based on cross section restoration, reconstructed shortening rates and subsidence analysis, we subdivide the East Pyrenean evolution into four phases. he irst (Late Cretaceous) phase is characterised by closure of an exhumed mantle domain between the European and Iberian rited margins, and simultaneous inversion of a salt-rich, thermally unequilibrated rit system. Shortening was distributed roughly equally between both margins during this early inversion phase. Following inversion, a quiescent phase (Paleocene) was apparently restricted to the retro-foreland. his phase may record the period of transition between inversion and full collision in the Eastern Pyrenees. he main collision phase (Eocene) records the highest shortening rates, which was predominantly accommodated in the pro-wedge. Retro-wedge shortening rates were lower than during the rit inversion phase. During the inal phase (Oligocene) the retro-wedge was apparently inactive and shortening of the pro-wedge slowed. his demonstrates that the relationship between the pro-and retro-wedges changes through time.

We used lithosphere-scale thermo-mechanical numerical models to simulate the evolution of a doubly vergent orogen. Our results show a similar evolutionary pattern as observed in the Pyrenees: A roughly symmetrical rit inversion phase is followed by an asymmetric collision phase. Rit inheritance was found to be essential for enabling double vergence. Other factors, such as surface processes and thin-skinned deformation, were found to have a signiicant efect on the crustal structure and strain partitioning between both wedges. A salt décollement layer in the sedimentary cover promotes the formation of a crustal antiformal stack such as observed in the Pyrenees and Alps by forming a wide and low-taper thin-skinned fold-and-thrust belt that forces crustal deformation to focus in the hinterland. Finally, we show that the evolution of the pro-and retro-wedges is inextricably linked: events or conditions on one side of the doubly vergent orogen have an immediate efect on the other side of the orogen. his is clearly demonstrated in our models by constant variations in shortening rates of the pro-and retro-wedge in response to accretion of new pro-wedge thrust sheets. he High Atlas (Morocco) and Pyrenees can be seen as examples of symmetric rit inversion and later asymmetric collision phases, respectively. hank you to Sébastien, Paul, and others at the CRPG for being friendly, welcoming, and helping me with my French.

Lastly, I want to thank my friends. For avoiding asking about work when I didn't feel like it, and discussing it when I needed it.

Du système de rit à l'orogène à double vergence : Un modèle évolutif basé sur l'étude de cas des Pyrénées Orientales et une étude des facteurs de contrôle à partir des modèles numériques mots clés géomodélisation, rétro-prisme, répartition de raccourcissement, Pyrénées Orientales, évolution

Résumé

Les orogènes à double vergence sont classiquement déinis comme deux prismes critiques opposés (pro et retro) qui évoluent ensemble. Les études récentes montrent que les rétro-prismes et leurs bassins d'avantpays associés se comportent diféremment des pro-prismes. Cependant, ni les facteurs qui mènent un orogène à devenir doublement vergent, ni la relation entre le pro-et rétro-prisme ne sont bien compris. Le but de cette étude est d'améliorer notre connaissance 1) de la relation entre le pro-et le rétro-prisme pendant l' orogénèse, 2) des facteurs contrôlant l' évolution d'un orogène à double vergence, et 3) d'un lien dynamique possible entre le pro-et le rétro-prisme.

Répondre à ces questions nécessite une connaissance améliorée de l' évolution d'un orogène à double vergence. Nous nous sommes concentrés sur les Pyrénées Orientales, en raison de la grande quantité de données disponibles. Nous avons efectué une étude de terrain tectono-stratigraphique détaillée à l' est du Massif de Saint Barthelemy et dans l'avant-pays autour de Lavelanet (plaque Européenne). Notre interprétation d'une coupe restaurée intègre une coniguration crustale pré-orogenique en tant qu'une marge hyper-amincie. Nous relions l' évolution détaillée du rétro-prisme à celle du pro-prisme (plaque Ibérique), ain de mieux contraindre la dynamique à l' échelle crustale. Nous subdivisons l' évolution des Pyrénées Orientales en quatre phases. La première phase (Crétacé Supérieur) est caractérisée par la fermeture d'un domaine de manteau exhumé entre les plaques et l'inversion synchrone d'un système de rit riche en sel et thermiquement déséquilibré. Le raccourcissement était distribué de façon égale entre les deux marges pendant cette première phase d'inversion. Une phase de quiescence (Paléocène), limitée au rétro-prisme, enregistre la transition entre l'inversion et la phase de collision. La phase de collision principale (Éocène) enregistre le taux de raccourcissement le plus élevé, et était principalement accommodé dans le pro-prisme. Pendant la phase inale (Oligocène) le rétro-prisme était largement inactif et le raccourcissement du proprisme a ralenti. Cela démontre que la relation entre le pro-et rétro-prisme change avec le temps.

Nous avons utilisé des modèles numériques 2D thermomécaniques à l' échelle lithosphérique pour simuler l' évolution d'un orogène à double vergence s'initie après avec un rit. Nos résultats montrent un modèle évolutif similaire à celui observé dans les Pyrénées Orientales avec une phase d'inversion du rit approximativement symétrique suivie d'une phase de collision asymétrique. L'héritage du rit est essentiel pour permettre le développement d'un orogène à double vergence. Des autres facteurs, comme les processus de surface et la déformation de la couverture, ont un efet signiicatif sur la structure crustale et la répartition du raccourcissement entre les deux prismes. Un niveau de décollement (sel) à la base de la couverture favorise la formation d'un empilement antiformal d' écailles crustales, similaire à la géométrie observée dans la Zone Axiale des Pyrénées, en formant un prisme à faible pente qui force la déformation crustale à se concentrer dans l'arrière-pays. Enin, nous montrons que l' évolution des pro-et rétro-prismes est inextricablement liée : des événements ou des conditions d'un côté de l' orogène ont un efet direct sur l'autre côté de l' orogène. Ceci est clairement démontré dans nos modèles par des variations constantes des taux de raccourcissement du pro-et rétro-prisme en réponse à l'accrétion dans le pro-prisme. Le Haut Atlas (Maroc) et Pyrénées peuvent être respectivement considérés comme des exemples d'inversion de rit symétrique et de phases de collision asymétrique ultérieures.

Foreword and acknowledgements iii

Foreword and acknowledgements

Well, this took longer than expected. he famous quote goes: "I apologise for writing such a long letter, for I had not the time to write a shorter one". From Pliny the Younger. Or Blaise Pascal. Or Mark Twain. It depends who you ask, apparently. Not everything you hear is true, it seems. Seeing as I truly do not have the time to crat this foreword and acknowledge those who have supported me, I suppose the premise of that quote shall also be disproven by this foreword. Door in het buitenland te wonen ben ik het belang van familie steeds meer op prijs gaan stellen. Mijn bezoeken thuis waren gevoelsmatig te infrequent en te kort. Ik wil jullie bedanken voor de steun, zowel wanneer ik het nodig had als wanneer dat niet het geval was. De motivatiecampagne per post, hoe klein ook, heet me elke keer weer toch een hoop plezier bezorgd. Dank jullie wel, en tot gauw.

Repeatedly have I been amazed by the amount of energy and efort invested by Mary and Ritske in supporting me during these four years (and a few months). Whenever I received comments on something I wrote, oten given to you on short notice, the quality of the response has helped me improve the text by a signiicant amount. My blatant ignorance of administrative requirements must also have caused quite some headaches, but even so, the support I have received from both of you has surpassed my expectations on numerous occasions. hank you for your patience, knowledge, and occasional motivating deadlines. tectonique (0,53 mm/y au maximum) directement au sud du front de poussée. Ceci conduit à la formation d'un bassin marin profond. Le raccourcissement de l'avant-pays sud-pyrénéen ralentit considérablement après la in de l'Yprésien, chutant à 1,0 mm/an. Cependant, à partir de ~50 Ma, le raccourcissement interne de la Zone Axiale accommode en moyenne 1,1 mm/an de plus (~ 23 km au total), comme en témoigne le début de son exhumation. La sédimentation dans les Pyrénées méridionales devient continentale après ~36 Ma, après que le soulèvement des Pyrénées Occidentales ait isolé de l' océan Atlantique le Bassin d'avantpays de l'Èbre. Le raccourcissement inal est alors principalement accommodé par un plissement de détachement au-dessus d'un décollement évaporitique. Dans cette partie de l'avant-pays sud-pyrénéen, les plus jeunes sédiments préservés (Rupélien) sont déformés, suggérant une déformation continue jusqu'au moins ~28 Ma.

Modèle à l' échelle crustale La comparaison des évolutions de chacun des deux avant-pays, décrites ci-avant, nous permet de diviser l' évolution des Pyrénées Orientales en quatre phases. La première phase (Crétacé supérieur) se caractérise d'une part par la fermeture d'un domaine de manteau exhumé entre les plaques ibérique et européenne, et d'autre part par l'inversion d'un système extensif riche en sel et non équilibré thermiquement. Le raccourcissement global (~1 mm/an) est alors réparti à peu près équitablement entre les deux marges durant une vingtaine de millions d'années. Vient ensuite la deuxième phase, à savoir une phase de quiescence (Paléocène) apparemment restreinte au rétro-système d'avant-pays et marquée par une déformation lente et continue dans le pro-système d'avant-pays (~0,4 mm/an). Dans les Pyrénées Orientales, cette deuxième phase précède le début de la collision principale (Eocène) qui enregistre le taux de raccourcissement moyen le plus élevé (~3,1 mm/an). Durant cette troisième phase, le raccourcissement est principalement accommodé dans le pro-système d'avant-pays. Enin, au cours de la phase inale (Oligocène), le rétrosystème d'avant-pays apparait inactif et le raccourcissement du pro-système d'avant-pays ralentit (~2,2 mm/an). En ignorant le raccourcissement que pourrait engendrer la fermeture du domaine de manteau exhumé, le raccourcissement total des Pyrénées Orientales est d' environ 111 km. Le rétro-système d'avantpays accommode au total environ 20 km de raccourcissement, principalement le long du Chevauchement Frontal Nord-Pyrénéen.

Nous voyons ainsi que le raccourcissement, à l' origine réparti de manière équitable de part et d'autre de l' orogène, devient bien plus prononcé dans le pro-système d'avant-pays. Cette modiication de la distribution du raccourcissement coïncide avec le début de la subduction de la croûte inférieure ibérique et du manteau lithosphérique (collision). Par conséquent, le changement dans la distribution de raccourcissement peut être dû à l'atteinte de seuils internes intrinsèques aux systèmes extensifs inversés, sans pour autant dépendre d'un quelconque forçage externe tel que la cinématique des plaques.

Modélisation de l'héritage extensif, des processus de surface et du décollement salifère

Nous avons utilisé des modèles numériques pour tester l'hypothèse que les systèmes extensifs inversés évoluent en orogènes à double vergence avec un changement dans la distribution de raccourcissement. En outre, nous avons cherché à expliquer le mécanisme responsable de la création d'une pile crustale antiforme, telle que celle formant la Zone Axiale dans les Pyrénées. Nous avons utilisé des modèles numériques thermo-mécaniques de haute-résolution, ce qui nous a permis de modéliser la déformation à l' échelle lithosphérique tout en conservant une bonne résolution des déformations dans les ceintures de plissement et de chevauchement et dans les bassins sédimentaires associés.

Résultats

Les efets de l'héritage extensif, du décollement salifère et des processus de surface aboutissent à une évolution en trois phases. La phase 0 est la phase d' extension pré-orogénique, où 50 km d' extension résultent en un système extensif étroit, à peu près symétrique et délimité par deux zones de cisaillement frictionplastique d' échelle crustale qui s' enracinent dans la croûte moyenne, plus faible rhéologiquement. Deux zones de cisaillement friction-plastique conjuguées dans le manteau lithosphérique permettent l' extension dans ce dernier.

Pendant la phase 1, le raccourcissement lithosphérique initial conduit à une inversion symétrique de la zone extensive. Les zones de cisaillement extensives héritées sont préférentiellement réactivées, ce qui entraine la soulèvement d'un bloc central en clé de voûte. La distribution de la vitesse de déformation montre que les deux zones de cisaillement conjuguées dans la croûte supérieure et le manteau supérieur lithosphérique sont simultanément activées pendant cette première phase d'inversion.

La phase 2 comprend le développement d'un édiice orogénique asymétrique à l' échelle crustale. Après une inversion symétrique, la localisation de la déformation sur une seule zone de cisaillement à grande échelle déclenche une subduction asymétrique de la croûte inférieure et du manteau lithosphérique, ce qui donne lieu à l'abandon des autres zones de cisaillement précédemment actives. Les premières ceintures étroites de plissement et de chevauchement dites de couverture se développent des deux côtés, s' enracinant à la fois dans les zones de cisaillement profond d'arrière-pays et dans le glissement de couverture du bloc central en clé de voûte. Dans le pro-système d'avant-pays, la sédimentation syn-tectonique favorise une large ceinture de plissement et de chevauchement. La déformation de socle se propage vers l' extérieur de l' orogène au-dessous de la ceinture de couverture en accrétant de nouvelles nappes de croûte supérieure. Les plus anciennes nappes du pro-prisme migrent lentement sur la plaque supérieure, réactivant la zone de cisaillement du rétro-prisme. Après 230 km de convergence, un découplage eicace le long du décollement salifère et une sédimentation sur le prisme, en partie distale de ce dernier, facilitent la propagation de la ceinture de plissement et de chevauchement de couverture du pro-prisme vers l' extérieur de l' orogène. Les nappes de socle du pro-prisme forment un empilement antiforme dont la croissance se fait par accrétion basale.

Distribution du raccourcissement La répartition du raccourcissement entre le pro-prisme et le rétro-prisme est décrite ci-dessous. L'inversion précoce de la phase 1 est symétrique, les deux côtés accommodant environ 50% de la convergence totale. Après le début de la subduction de la croûte inférieure et du manteau lithosphérique durant la phase 2, le pro-prisme accommode ~80% de la convergence totale tandis que le rétro-prisme en accommode seulement 20%. Les modèles sans héritage extensif ne montrent pas ce changement dans la distribution du raccourcissement.

Mécanisme d' empilement antiforme

La formation d'un empilement antiforme est le résultat des efets combinés d'un décollement salifère, de la sédimentation syn-tectonique et de l' érosion. Le décollement salifère crée un système à deux prismes : 1) Un prisme de socle caractérisé par un biseau critique relativement élevé et contrôlé par le décollement en croûte moyenne (« prisme crustal »), et 2) un prisme de couverture caractérisé par un biseau critique faible (« prisme de couverture ») contrôlé par le décollement salifère. Puisque les volumes élémentaire de ces prismes ignorent leurs voisins et la géométrie globale du système, ils ne se déforment ou ne se déplacent que sous l'inluence des forces suivantes : (a) le frottement le long de leur décollement basal respectif, (b) les contraintes liées à la convergence et (c) la charge de la pile sédimentaire sus-jacente. Le prisme de couverture recouvre le socle, contribuant ainsi à cette charge de la pile sédimentaire sus-jacente, et devrait donc être considéré comme faisant partie du prisme crustal. Efectivement, le système s'agence comme un prisme au sein d'un prisme, plutôt que comme un prisme au sommet d'un autre. Cette disposition signiie que les deux prismes partagent le même angle alpha (inclinaison de la surface sommitale), à savoir celui dicté par le décollement le plus faible. Le prisme de couverture dicte ainsi l' ouverture du biseau en dessous de celle du biseau critique du prisme crustal. Le prisme crustal se déforme de manière interne (raccourcissement interne) pour augmenter l' ouverture du biseau. Cependant, le prisme salifère ne peut pas supporter cette ouverture élevée, et se déforme pour maintenir une ouverture faible, obligeant le prisme crustal à se déformer avec un raccourcissement plus interne. La sédimentation syn-tectonique réduit également l' ouverture dans l'avant-pays, tandis que l' érosion la réduit dans l'arrière-pays. Combinés, ces efets aboutissent inalement à la formation d'une pile antiforme de nappes de socle.

Modélisation des contrôles de décollement sur la structure orogénique D'après notre modèle, que nous avons décrit précédemment, l'héritage extensif semble permettre la formation d'un orogène à double vergence avec une évolution et une distribution des déformations similaires à celles observées dans les Pyrénées. Cependant, la répartition des contraintes entre le proprisme et rétro-prisme varie considérablement d'un système naturel à l'autre. En outre, notre modèle montre que la déformation de couverture peut contrôler la structure crustale. Par conséquent, nous avons étudié l'inluence de plusieurs facteurs sur la distribution des déformations et l' évolution des orogènes à double vergence. Nous nous sommes concentrés sur l'inluence de la rhéologie et de la distribution d'un niveau de décollement sédimentaire.

Résultats

Les modèles sans héritage extensif aboutissent à un orogène hautement asymétrique où ~95% du raccourcissement est accommodé dans un pro-prisme très large. Tous les modèles avec héritage extensif aboutissent à une évolution en trois phases que nous avons explicité ci-dessus. Un décollement schisteux engendre des ceintures de plissement et de chevauchement de couverture aussi bien dans le pro-système d'avant-pays que dans le rétro-système d'avant-pays. Ces ceintures de déformation ne sont actives qu'à proximité de la zone de cisaillement frontale qui implique le socle. Elles sont abandonnées dès qu'une nouvelle zone de cisaillement plus distale est formée dans le pro-prisme. Un décollement salifère, plus eicace dans le découplage de la couverture sédimentaire par rapport à son socle, permet le glissement gravitationnel, ce qui crée des ceintures de couverture plus larges. Ces dernières transportent la couverture sédimentaire du coeur de l' orogène vers l'avant-pays externe, ce qui conduit à un pro-prisme crustal plus étroit. La sédimentation syn-tectonique stabilise ces ceintures de plissement et de chevauchement de couverture, ajoutant encore à la charge d'avant-pays. L' érosion est plus eicace dans la zone interne de l' orogène et engendre un orogène plus étroit. Lorsque les processus de surface ne sont pas pris en compte, le décollement salifère favorise un raccourcissement dans le rétro-prisme plus important que ne le fait le décollement schisteux. Cependant, lorsque les processus de surface sont considérés dans la modélisation, quelle que soit la rhéologie du décollement, le pro-prisme accommode une fois encore ~80% de la convergence totale tandis que le rétro-prisme en accommode seulement 20%.

Une distribution asymétrique du décollement peut inluencer la polarité de la plaque, plaçant préférentiellement le décollement dans le pro-système. Cependant, d'autres facteurs dont nous n'avons pas tenu compte peuvent avoir une plus grande inluence sur la polarité des plaques.

Comparaisons avec les systèmes naturels Nos modèles reproduisent systématiquement une évolution en trois phases lorsque l'héritage extensif est considéré. Ceci nous permet de proposer un modèle évolutif générique pour les systèmes extensifs inversés comme suit : (1) la phase 0 est la phase d' extension, (2) la phase 1 voit une inversion à peu près symétrique des structures extensives héritées, et (3) la phase 2 est la phase de collision asymétrique, où la subduction continentale entraine une accommodation d' environ 80% de la convergence totale dans le pro-prisme. Nous avons comparé plusieurs systèmes naturels au modèle proposé.

Le Haut Atlas (Maroc) est un système extensif inversé d'âge alpin. Les structures extensives héritées ne sont pas complètement inversées et le raccourcissement global est très faible (~26 km). Les failles crustales du côté nord plongent vers le sud et vice versa. La répartition du raccourcissement est symétrique, avec ~13 km de raccourcissement accommodés de chaque côté d'un bloc central. Nous proposons que le Haut Atlas se conforme à la phase 1 de notre modèle évolutif. L' orogène pyrénéen, comme nous l'avons décrit précédemment, est également un système extensif inversé. Les estimations de raccourcissement total vont de ~100 km à ~165 km dans les Pyrénées Centrales. Dans le cas d'un raccourcissement total maximal (~165 km), le pro-prisme accommode ~128 km de raccourcissement tandis que le rétro-prisme en accommode ~37 km. Notre reconstruction dans les Pyrénées Orientales se traduit par ~ 111 km de raccourcissement total sans tenir compte de la fermeture du domaine de manteau exhumé. Le pro-prisme accommode alors ~92 km de raccourcissement alors que le rétro-wedge

Chapter 1

Introduction Chapter 1. Introduction 10 his thesis investigates the relationship between, and evolution of, pro-and retro-wedges in doubly vergent orogens and how this is inluenced by a number of factors, in particular extensional inheritance and the rheology of a décollement in the sedimentary cover. he approach uses numerical modelling coupled with a detailed study of both sides of the eastern Pyrenean orogen. he work is based around several concepts and techniques that have been developed to describe the processes involved in orogenesis. hese concepts and techniques are introduced in this chapter. A summary of the state of knowledge on the Pyrenean orogeny provides the regional context for the East Pyrenean case study (Chapter 2). Finally, several limitations of the concepts and techniques explained in this chapter are outlined, and the scientiic questions addressed in this thesis are presented.

Concepts

hrust wedges and critical taper

To create a mountain belt, the crust has to be shortened and thickened. In the brittle part of the crust, this shortening is accommodated along a series of thrust faults that stack slices of crust on top of each other. hese thrust systems oten result in an approximate wedge shape that tapers toward the foreland (Figure 1.1) [e.g., [START_REF] Suppe | A retrodeformable cross section of northern Taiwan[END_REF]. Critical taper theory is a mathematical model, developed in the 1980's, that describes in mechanical terms how these wedges form, and what factors inluence the shape of the wedge [e.g., [START_REF] Davis | Mechanics of fold-and-thrust belts and accretionary wedges[END_REF][START_REF] Dahlen | Noncohesive critical Coulomb wedges: An exact solution[END_REF]Dahlen, , 1990]]. he wedge shape is deined by the slope of the top surface (angle ) and the slope of the basal décollement surface (angle ). he taper is the sum of these angles ( ; Figure 1.2a). A critically tapered wedge will grow self-similarly to maintain the same taper. Note that under most conditions, there are two values for critical taper, and a wedge with a taper between those two values is stable and will not deform internally (Figure 1.2b) [START_REF] Dahlen | Noncohesive critical Coulomb wedges: An exact solution[END_REF]. For frontal accretion, it suices to only regard the lower of the two critical tapers, as this is the irst taper a newly formed wedge will encounter. If a wedge has a subcritical taper, it will deform internally and not slide along the basal décollement until the taper has increased to become critical. If a wedge is supercritical, it will either extend internally (taper above stable ield) or slide along the décollement without deforming (taper in stable ield) until the taper is lowered through frontal accretion and/or surface erosion. Assuming a Mohr-Coulomb rheology, the critical taper depends on the relative magnitudes of the coeicient of friction along the basal décollement and the angle of internal friction of the wedge material [Dahlen, 1990]. Lowering the coeicient of friction along the basal décollement will reduce critical taper, and increasing it will increase critical taper. Lowering the angle of internal friction of the wedge material (making it weaker) will increase critical taper, because that makes it easier to deform the wedge than to slide along the basal décollement. he wedge will then deform internally until the new, steeper critical taper has been reached. However, if the wedge material is weakened to the point it can no longer support the topographic slope ( ), the wedge will collapse. Increasing the angle of internal friction of the wedge material lowers the critical taper. he wedge material is now strong enough to withstand the horizontal stress that drives convergence and it slides along the basal décollement instead. hese relationships are clear in the mathematical expression of this model [Dahlen, 1990]:

(1.1)
where is the angle of internal friction of the wedge material and is the friction coeicient along the basal décollement. his equation assumes noncohesive Coulomb failure of dry sand, and uses several simpliications for the case that [Dahlen, 1990]. Improvements on this basic mathematical model account for cohesion and pore luid pressure [START_REF] Davis | Mechanics of fold-and-thrust belts and accretionary wedges[END_REF]Dahlen, 1990]. In thrust wedges above a salt décollement, basal friction is not dependent on normal stress. he yield strength of salt instead depends on temperature and strain rate. If one assumes a constant yield strength , the critical taper of the wedge becomes [START_REF] Davis | he role of salt in fold-and-thrust belts[END_REF]:

(1.2)
where is rock density, the depth to the basal detachment, and the pore luid pressure relative to lithostatic pressure. If =1 MPa, appropriate for salt, the resulting taper is very low, ~1° [START_REF] Davis | he role of salt in fold-and-thrust belts[END_REF].

Following the development of critical taper theory, further investigations aimed to explain the inluence of other factors, such as syn-tectonic sedimentation [e.g., [START_REF] Storti | Synchronous and velocity-partitioned thrusting and thrust polarity reversal in experimentally produced, doubly-vergent thrust wedges: Implications for natural orogens[END_REF]Ford, 2004], erosion [e.g., [START_REF] Persson | Analogue models of orogenic wedges controlled by erosion[END_REF][START_REF] Konstantinovskaya | Accretionary orogens: Erosion and exhumation[END_REF], and a ductile crustal root [Willett et al., 1993]. Sedimentation efectively strengthens the distal wedge and undeformed foreland, focusing deformation in the interior of the wedge [START_REF] Storti | Synchronous and velocity-partitioned thrusting and thrust polarity reversal in experimentally produced, doubly-vergent thrust wedges: Implications for natural orogens[END_REF]Fillon et al., 2013b]. Erosion reduces the topographic slope ( ), making the wedge subcritical. Internal deformation to restore taper then results in a smaller wedge that exhumes deep materials [START_REF] Persson | Analogue models of orogenic wedges controlled by erosion[END_REF]. For an overview of wedge models, see Buiter [2012].

1.1.2 Flexural foreland basins he association of foreland basins with orogens has long been recognised [e.g., [START_REF] Abouin | Geosynclines[END_REF]. A lexural model of foreland basins was developed in which the topographic load of an adjacent orogen creates a downward lexure of the crust, thus forming a foreland basin (Figure 1.3) [START_REF] Walcott | Flexure of the lithosphere at Hawaii[END_REF][START_REF] Price | Large-scale gravitational low of supracrustal rocks, southern Canadian Rockies[END_REF][START_REF] Beaumont | Foreland Basins[END_REF]. his link between orogen growth and foreland lexure means that investigating the foreland basin can reveal information about the growth of the orogen. In the lexural foreland model, the width and depth of a foreland basin depend on two things: the magnitude of the load, and the lexural rigidity of the plate (oten expressed as 'efective elastic thickness' [START_REF] Watts | he efective elastic thickness of the lithosphere and the evolution of foreland basins[END_REF]). he larger the load, the greater the downward delection of the plate, thus the deeper the basin. he more rigid the plate, the shallower the basin, but also the wider the lexed zone (increased diameter of the curve in the lexed plate). his model has been reined and adapted for a broken plate that is unconnected at the edge below the load [START_REF] Walcott | Flexure of the lithosphere at Hawaii[END_REF][START_REF] Turcotte | Geodynamics: Applications of continuum physics to geological problems[END_REF], a load that is distributed along the lexed plate instead of placed at the edge [START_REF] Beaumont | Foreland Basins[END_REF], lateral variations in lexural rigidity, and buried loads such as subducted oceanic slabs [START_REF] Royden | he tectonic expression slab pull at continental convergent boundaries[END_REF].

Classical foreland basins form on a plate that is overthrust by an orogenic wedge. Any record of the downward lexure through time (subsidence) is thus transported toward the orogen during convergence. In the distal basin the plate is almost horizontal, thus a given amount of convergence will result in relatively little subsidence. In the proximal basin, the dip of the plate is much steeper (max. 9°), thus the same point on the mobile lower plate will experience gradually increasing subsidence. It is this hinterlandward transport of a point along a curved plate that gives foreland basin subsidence its classical accelerating pattern [e.g., [START_REF] Allen | Foreland Basins: An Introduction[END_REF]Vergés et al., 1998;[START_REF] Haddad | Subsidence history, gravity anomalies, and lexure of the northeast Australian margin in Papua New Guinea[END_REF]. 1.1.3 Pro-and retro-wedges and foreland basins Following the development of orogenic wedge models with a rigid backstop came the realisation that the upper plate, represented by the backstop in the experiments, should be deformable, just as the lower plate is. A new category of models was developed, where the substrate of a continuous layer of material is split into two halves (Figure 1.4) [Willett et al., 1993]. One half, representing the lower plate, is dragged towards the central singularity or S-point where the plates meet and the substrate of the lower plate is subducted. he other half, representing the upper plate, remains stationary. his asymmetrical boundary condition results in two wedges joined at the back: he wedge on the lower plate (pro-wedge) grows by frontal accretion and has a low critical taper (the lower of two critical tapers, see above). he wedge on the upper plate (retrowedge) grows outward by receiving material from the pro-wedge, resulting in a higher critical taper (the highest of two critical tapers, see above). he original description of a doubly vergent orogen with a proand retro-wedge was based on numerical models [Willett et al., 1993;[START_REF] Beaumont | Styles of crustal deformation in compressional orogens caused by subduction of the underlying lithosphere[END_REF][START_REF] Beaumont | Mechanical model for subduction-collision tectonics of Alpine-type compressional orogens[END_REF], but the same geometry has been observed in analogue models with a similar setup [e.g., [START_REF] Storti | Synchronous and velocity-partitioned thrusting and thrust polarity reversal in experimentally produced, doubly-vergent thrust wedges: Implications for natural orogens[END_REF]McClay et al., 2004;Hoth et al., 2007Hoth et al., , 2008]]. Modelling of doubly vergent orogens quickly evolved to account for plate lexure [e.g., [START_REF] Simpson | Inluence of the mechanical behaviour of brittle-ductile fold-thrust belts on the development of foreland basins[END_REF], sedimentation [Duerto and McClay, 2009], asymmetric erosion [Willett et al., 1993[START_REF] Co | Uplit , Shortening[END_REF][START_REF] Willett | Orogeny and orography: he efects of erosion on the structure of mountain belts[END_REF], friction along the base [START_REF] Beaumont | Styles of crustal deformation in compressional orogens caused by subduction of the underlying lithosphere[END_REF][START_REF] Naylor | A discrete element model for orogenesis and accretionary wedge growth[END_REF], partial subduction of the lower plate [START_REF] Beaumont | Styles of crustal deformation in compressional orogens caused by subduction of the underlying lithosphere[END_REF][START_REF] Beaumont | Mechanical model for subduction-collision tectonics of Alpine-type compressional orogens[END_REF], and viscous rheologies and convergence rates [START_REF] Rossetti | he inluence of backstop dip and convergence velocity in the growth of viscous doubly-vergent orogenic wedges: Insights from thermomechanical laboratory experiments[END_REF]. Eventually, numerical experiments simply replaced the S-point boundary condition at the base with full modelling of lithospheric plates resting on top of the asthenosphere and imposing convergence at the side boundaries [e.g., Jammes and Huismans, 2012;Erdős et al., 2014Erdős et al., , 2015]]. he asymmetry in doubly vergent orogens led to the realisation that the lexural foreland basins also record diferent subsidence histories (Figure 1.5) [Naylor and Sinclair, 2008;Sinclair and Naylor, 2012]. he pro-foreland basin, on the lower plate, records the classical accelerating subsidence pattern as a result of being transported towards the orogen. However, the retro-foreland basin is not transported toward the orogen, because the upper plate remains stationary relative to the plate boundary (S-point). he retroforeland basin thus only records subsidence as the orogen grows and lexure increases, and retro-foreland subsidence tends to be slow and long-lasting as a result [Naylor and Sinclair, 2008;Sinclair and Naylor, 2012]. If the orogen is in a lux steady state (accretional inlux = erosional outlux), both foreland basins remain the same size. Points in the retro-foreland basin thus do not record subsidence during steady state, and points in the pro-foreland record only subsidence due to convergence toward the orogen. his diference between pro-and retro-foreland basins also implies they preserve diferent parts of the orogenic history. Both basins record the early growth history of the orogen, but the early pro-foreland record is eventually destroyed. During steady state, only the pro-foreland records subsidence. In a suiciently mature orogen, combining the subsidence histories of the pro-and retro-foreland should thus result in a more complete picture of orogenic history [Naylor and Sinclair, 2008].

Techniques

Cross-section balancing

In the analysis of natural foreland fold-and-thrust belts, an important technique is the construction of cross sections. he available data in natural systems is usually incomplete, and thus permits several diferent interpretations that imply diferent amounts of shortening accommodated along the section. For this reason, it is important to ensure that the cross section is geologically valid, and that it is physically possible to achieve the proposed geometry. Balancing of cross sections is a technique where a deformed cross section is restored to its undeformed state using set rules (Figure 1.6), and then corrected for impossible geometries [START_REF] Dahlstrom | Balanced cross sections[END_REF][START_REF] Hossack | he use of balanced cross-sections in the calculation of orogenic contraction: A review[END_REF][START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF]. In constructing a valid balanced cross section, one has to take into account the following considerations: 1) Choosing the right orientation of the section. Restoring a 2D cross section assumes plane strain, thus material is transported parallel to the section line. A section that is oblique to the transport direction results in an overestimation of shortening. he transport direction can be determined based on geological evidence. his can be simply perpendicular to the strike of large-scale folds and faults [START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF]. For numerous small-scale folds, the strike is not a reliable guide for transport direction, but the transport direction can still be estimated using stereonet analysis [START_REF] Hansen | Methods of Deducing Slip-Line Orientations from the Geometry of Folds[END_REF]. Mineral stretching lineations in metamorphic terranes can also be used as a guide to the (syn-metamorphic) tectonic transport direction [START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF].

2) Conserving area between the deformed and restored section [START_REF] Dahlstrom | Balanced cross sections[END_REF]. Since the section is parallel to the direction of transport and the restoration assumes plane strain, the volume of material in the deformed and restored sections should be the same. Only in exceptional cases where out-of-plane transport, compaction, or dissolution play a signiicant role can the volume change between sections [START_REF] Rowan | Cross-section restoration of salt-related deformation: Best practices and potential pitfalls[END_REF]. In the case of compaction and dissolution, the volume of the restored section should always be greater than that of the deformed section. 3) Conserving bed length [START_REF] Dahlstrom | Balanced cross sections[END_REF]. Competent layers can be folded and faulted, but tend to keep the same length. his serves as an important constraint on the restored cross section as it will determine how long the restored section is going to be. 4) Ensuring that the restored section displays viable dips and geometries for faults and stratigraphy, in line with well-documented analogous structures from a diferent location [START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF]. his includes eliminating voids in the restored section, which may result from restoring an incorrectly interpreted deformed section. Adhering to geologically sound geometries requires a knowledge of processes and geometries involved in the deformation of strata, such as fault-bend folds [e.g., [START_REF] Rich | Mechanics of low-angle overthrust faulting as illustrated by Cumberland thrust block, Virginia, Kentucky, and Tennessee[END_REF][START_REF] Medwedef | Multibend fault-bend folding[END_REF], fold-propagation folds and trishear [e.g., [START_REF] Suppe | Geometry and kinematics of fault-propagation folding[END_REF][START_REF] Erslev | Trishear fault-propagation folding[END_REF], and forced folds [e.g., Mitra and Mount, 1998;[START_REF] Miller | Deformation and secondary faulting associated with basementinvolved compressional and extensional structures[END_REF].

5) It is good practice to keep shortening within the section to a minimum, and keep the geometry as simple as the data will allow. When constructing multiple sections, consistency between adjacent sections is preferable [START_REF] Dahlstrom | Balanced cross sections[END_REF][START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF].

Taking into account these guidelines, the process of balancing a cross section is simply iteratively restoring and correcting the deformed section until it is valid [START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF]. he process of balancing sections was originally developed for use in thin-skinned foreland fold-and-thrust belts, but these principles can also be applied to crustal scale sections that involve basement in the deformation. he balancing technique remains useful and is still used today [e.g., [START_REF] Suppe | A retrodeformable cross section of northern Taiwan[END_REF]Roeder, 1992;Vergés, 1993;[START_REF] Mcquarrie | Crustal scale geometry of the Zagros fold-thrust belt, Iran[END_REF][START_REF] Labaume | Tectono-thermal history of an exhumed thrust-sheet-top basin: an example from the south Pyrenean thrust belt[END_REF].

Subsidence analysis

As described above, foreland basins record subsidence that relates to the evolution of the adjacent orogen. Subsidence analysis is the process that extracts this information from the sedimentary basin. he aim is to relate the accumulation of sediments and vertical movements of the top of the sedimentary column relative to the eustatic sea level, in order to obtain the vertical movements of the basement, and eventually isolate the subsidence signal that is purely due to tectonic forces.

Because subsidence analysis involves the thickness of a column of sediments through time, it is important to understand compaction. he thickness of a sedimentary layer decreases as it is buried, and the weight of the overburden literally squeezes it into a thinner layer. his reduction of volume is accommodated by a reduction in pore space, lowering porosity (Figure 1.7a). For normally pressured sediments, the relationship between depth and porosity is:

(1.3)
where is the porosity at depth, the initial porosity (a material constant), the compaction coeicient (another material constant), and is depth [START_REF] Allen | Basin Analysis, Second Edi[END_REF]. In overpressured sediments, the equation becomes: ( 1.4) where is the bulk density of the layer (grains + pore ill), the density of the pore luid (water), and the ratio of pore luid pressure relative to lithostatic pressure [START_REF] Allen | Basin Analysis, Second Edi[END_REF]. Material constants and are highly lithology-speciic [START_REF] Sclater | Continental stretching: an explanation of the post-mid-cretaceous subsidence of the central north sea basin[END_REF][START_REF] Halley | High-porosity Cenozoic carbonate rocks of south Florida: progressive loss of porosity with depth[END_REF][START_REF] Sonnenfeld | Brines and evaporites[END_REF]. Each lithology (or mix of lithologies) thus has its own porosity-depth curve depending on initial porosity, lithology, pore luid pressure, diagenesis and other factors. Using the porosity-depth relationship, it is possible to calculate the thickness of a given layer at any depth. With the assumption that the thickness change is entirely the result of reducing pore volume during compaction, the thickness of a decompacted layer takes the form:

(1.5) with representing thickness of the layer and pore volume. Together with equation 1.3, this gives the general decompaction equation [START_REF] Allen | Basin Analysis, Second Edi[END_REF]:

(1.6)
where is the thickness of the decompacted layer (depth of base minus depth of top), is the thickness of the compacted layer, the second-to-last term represents the pore volume of the compacted layer, and the inal term represents the pore volume of the decompacted layer. here is no exact solution for this equation. Instead, it is solved by iteratively adjusting the values and until both sides of the equation are equal.

Using the general decompaction equation, a decompaction curve can be constructed. his is done by removing the top layer and recalculating the decompacted thickness of all layers underneath, then repeating the process until there is no sediment let on top of basement (Figure 1.7b). When correcting for paleobathymetry and eustatic sea level variations, the thickness of the sedimentary column (i.e. depth to basement) can be plotted against time to give total subsidence [START_REF] Allen | Basin Analysis, Second Edi[END_REF].

Total subsidence represents the 'true' vertical movements of basement through time, under the inluence of tectonic forces and isostatic adjustments to the sediment mass, and changes in the mass of the water column due to sea level variations. To isolate only the tectonic signal, it is necessary to correct for the isostatic response to sediments, paleobathymetry and eustatic sea level, a process called backstripping [Steckler and Watts, 1978;[START_REF] Allen | Basin Analysis, Second Edi[END_REF]:

(1.7)
where is depth to basement relative to present-day sea level, is the decompacted sediment column, is the density of the mantle, is the paleobathymetry, and the paleo-sea level relative to present-day. he backstripping equation assumes Airy isostasy to compensate for the load of the sediment column and part of the water column. he result is tectonic subsidence, which represents the depth of basement if there had been no sediment, and sea level would have been the same as today (Figure 1.8). It is a measure of the tectonic forcing in a basin's evolution [Steckler and Watts, 1978].

Tectonic subsidence is a valuable tool in the analysis of foreland basins [e.g., Desegaulx and Brunet, 1990;Vergés et al., 1998]. It is equally useful in other tectonic settings, as long as there is accumulation of sediment relative to a known datum (i.e. sea level) [e.g., [START_REF] Xie | Plate tectonics and basin subsidence history[END_REF].

Numerical mechanical modelling

Numerical modelling of deformation has proven to be a very efective tool in geoscience. With the availability of ever increasing computational power, numerical models show a clear trend of increasing complexity and resolution. he earliest numerical investigations applied purely thermal models to mantle convection in subduction zones [e.g., [START_REF] Richter | Simple plate models of mantle convection[END_REF][START_REF] Schubert | Numerical Models of Mantle Convection[END_REF]. Mechanical models of continental collision followed [e.g., [START_REF] Bird | Finite element modeling of lithosphere deformation: the Zagros collision orogeny[END_REF][START_REF] Daignières | Modèle de type Norton-Hof généralisé pour l' étude des déformation lithosphériques[END_REF]. Since then, model codes have become steadily more capable, including thermo-mechanically coupled deformation, higher resolutions, 3D capable codes, coupled surface processes, etc. [e.g., [START_REF] Fullsack | An arbitrary Lagrangian-Eulerian formulation for creeping lows and its application in tectonic models[END_REF][START_REF] Braun | DOUAR: A new threedimensional creeping low numerical model for the solution of geological problems[END_REF]hieulot, 2011]. and Watts [1978]. hree main categories of numerical implementation have been developed: inite element (Figure 1.9a) [e.g., [START_REF] Fullsack | An arbitrary Lagrangian-Eulerian formulation for creeping lows and its application in tectonic models[END_REF][START_REF] Braun | DOUAR: A new threedimensional creeping low numerical model for the solution of geological problems[END_REF], inite diference [e.g., [START_REF] Waltham | Finite diference modelling of hangingwall deformation[END_REF][START_REF] Waltham | he velocity description of deformation. paper 1: theory[END_REF][START_REF] Gerya | Introduction to Numerical Geodynamic Modelling[END_REF], and discrete element (Figure 1.9b) [e.g., Finch et al., 2004;[START_REF] Naylor | A discrete element model for orogenesis and accretionary wedge growth[END_REF]. Finite element and inite diference models are both implementations of deformation in a continuum, whereas discrete element models consist of discrete, rigid particles that are loosely packed together, analogous to sand.

Mechanical models that are based on continuum mechanics model deformation by solving the Stokes equation. When assuming incompressibility and negligible inertial forces, it can be written as [e.g., [START_REF] Currie | hinning of continental backarc lithosphere by lowinduced gravitational instability[END_REF]: (1.8) where is the stress tensor, the spatial dimensions, is density and the gravitational acceleration. Materials in the model deform by lowing as a viscous luid, following a power law rheology that relates the strain rate to the creep stress to the power :

(1.9) where is the strain rate tensor, is a pre-exponential material constant, is the creep stress, the exponent, is the activation energy, is pressure, is the activation volume, is the gas constant and is temperature. Material constants , , , and are determined from laboratory data [Karato and Wu, 1993;Gleason and Tullis, 1995]. Extrapolating the mechanical properties of rocks from laboratory experiments to geological timescales and rates may introduce errors in the mechanical model. Another limitation of numerical mechanical models using continuum mechanics is that they cannot model 'true' frictional (brittle) deformation. A method to simulate frictional deformation is to calculate the theoretical yield stress and adjust the viscosity to ensure the creep stress does not exceed the yield stress [START_REF] Fullsack | An arbitrary Lagrangian-Eulerian formulation for creeping lows and its application in tectonic models[END_REF]hieulot, 2011]. For a more in-depth explanation of the governing equations of the numerical models used in this work, see appendix B.

Mechanical models (numerical and analogue) can be used to test the inluence of various factors on a certain deformational process. An advantage of numerical models over analogue models is that all experiments can be performed at full scale with geological deformation rates. Numerical models do not have to rely on suitably scaled analogue materials. In addition, the control over factors such as sedimentation and erosion is more reined in numerical models. However, because numerical models can simulate processes at the spatial and temporal scales of geological processes, material properties obtained from laboratory experiments must be extrapolated, introducing some uncertainty. his demonstrates the importance of benchmarking a numerical model against known analogue and numerical experiments.

Numerical deformation models have been applied to a wide variety of problems: subduction zones [e.g., [START_REF] Schellart | Evolution and diversity of subduction zones controlled by slab width[END_REF][START_REF] Currie | hinning of continental backarc lithosphere by lowinduced gravitational instability[END_REF], continental collision [e.g., Willett et al., 1993;Pifner et al., 2000;Sinclair et al., 2005;Jammes and Huismans, 2012;Erdős et al., 2014;[START_REF] Ueda | Geomorphological-thermo-mechanical modeling: Application to orogenic wedge dynamics[END_REF], riting [e.g., [START_REF] Lesne | Dynamics of intracontinental extension in the north Baikal rit from two-dimensional numerical deformation modeling[END_REF]Huismans and Beaumont, 2007], salt décollements [e.g., [START_REF] Ruh | Numerical investigation of deformation mechanics in fold-andthrust belts: Inluence of rheology of single and multiple décollements[END_REF][START_REF] Ghazian | Numerical modelling of the role of salt in continental collision: An application to the southeast Zagros fold-and-thrust belt[END_REF], and many more.

In this work, we use a modiied version of the numerical code FANTOM, a high-resolution arbitrary Eulerian-Lagrangian inite-element code for modelling thermo-mechanically coupled deformation [hieulot, 2011]. his code is described in more detail in Chapters 3, 4 and Appendix B. 1) he Aquitaine retro-foreland basin, a mildly deformed sedimentary basin. In the east, this basin comprises successive sub-basins of Late Cretaceous to Eocene age that overlie Paleozoic basement [START_REF] Plaziat | Late Cretaceous to late Eocene paleogeographic evolution of southwest Europe, Palaeogeogr[END_REF]Tambareau et al., 1995;[START_REF] Martin-Martin | Tectonic controls on the deposits of a foreland basin: an example from the Eocene Corbières-Minervois basin, France[END_REF]Christophoul et al., 2003]. In general, younger subbasins are situated to the north of older sub-basins. In the west, syn-orogenic sediments overlie Triassic salt to Lower Cretaceous rit sediments [Puigdefàbregas and Souquet, 1986;Desegaulx and Brunet, 1990].

Pyrenees

2) he Sub-Pyrenean Zone, a more deformed (folded and faulted) part of the foreland basin immediately north of the North Pyrenean Frontal hrust, usually limited to the north by a blind thrust [Bilotte et al., 1988;[START_REF] Deramond | Relationships between thrust tectonics and sequence stratigraphy surfaces in foredeeps: model and examples from the Pyrenees (Cretaceous-Eocene, France, Spain)[END_REF].

3) he North Pyrenean Zone, a narrow belt between the North Pyrenean Frontal hrust to the north and the North Pyrenean Fault to the south. he intensity of deformation in this zone is signiicantly higher than in the zones further north, and mainly north-vergent [e.g., Choukroune, 1974]. his zone comprises inverted Early Cretaceous rit basins [Souquet et al., 1985;Debroas, 1990], allochthonous basement massifs [Souquet et al., 1977;Souquet and Peybernès, 1987;Baby, 1988;Baby et al., 1988;de Saint Blanquat et al., 1990], and the Metamorphic Internal Zone immediately north of the North Pyrenean Fault. he Metamorphic Internal Zone mostly consists of Mesozoic sediments that have undergone high temperature, low pressure metamorphism and are highly deformed and brecciated [Golberg and Leyreloup, 1990;Debroas et al., 2010aDebroas et al., , 2010b;;de Saint Blanquat et al., 2016]. Blocks of mantle peridotites are embedded within the metamorphosed sediments and breccias (Figure 1.11) [Lagabrielle et al., 2010]. hese ultrabasic mantle rocks are named lherzolites, ater Lherz, a locality south of the Trois Seigneurs Massif, where they were irst described [e.g., [START_REF] Bonney | he Lherzolite of Ariège[END_REF]. Due in part to the presence of lherzolites, the North Pyrenean Fault that forms the southern border of this zone is generally regarded as the suture between the Iberian and European plates. he North Pyrenean Fault is a major vertical fault, whether it accommodated a large strike-slip displacement or not was diicult to constrain, and thus debated [e.g., [START_REF] Casteras | Recherches sur la structure du versant nord des Pyrénées centrales et orientales[END_REF]Choukroune, 1974;Souquet et al., 1977;Choukroune and Mattauer, 1978]. Due to the lack of plate kinematic consensus, part of this debate is ongoing [e.g., Olivet, 1996;Jammes et al., 2009;Vissers and Meijer, 2012b]. 4) he Axial Zone comprises Paleozoic rocks deformed mainly by the Variscan orogeny. he ECORS Pyrenees deep seismic relection proile across the Central Pyrenees constrains the deep structure [START_REF] Choukroune | he Ecors Pyrenean deep seismic proile relection data and the overall structure of an orogenic belt[END_REF]Roure et al., 1989;Muñoz, 1992;Roure and Choukroune, 1998;Beaumont et al., 2000]. he Iberian lower crust subducts below the European plate, and Iberian upper crustal thrust sheets form a south-vergent antiformal stack [Muñoz, 1992]. To the west and to the east, the stack of Iberian thrust sheets is interpreted to form a less well developed antiformal stack [Vergés et al., 1995;Teixell, 1998].

5) he South Pyrenean Zone, a south verging thin-skinned fold-and-thrust belt of variable width that is delimited by the Vallfogona thrust in the east, the Sierras Marginales in the Central Pyrenees, and the external Sierras in the west. he thrust sheets in the South Pyrenean Zone were thrust southward along an evaporite décollement. For the upper, older thrust sheets this décollement consists of Triassic Keuper evaporites, while for the lower, younger structures this décollement consists of two Eocene evaporite layers (Beuda Fm. and Cardona Fm.) [START_REF] Puigdefàbregas | hrusting and foreland basin evolution in the Southern Pyrenees, in hrust Tectonics[END_REF]Vergés et al., 1992;Vergés, 1993]. 6) he Ebro pro-foreland basin. his basin consists of Eocene and Oligocene stratigraphic units that have been mildly deformed by detachment folding above the upper Eocene Cardona evaporite formation [Vergés et al., 1992;Vergés, 1993]. he Ebro foreland basin is bordered to the south by the north-vergent Catalan Coastal Ranges.

Hyperextended rit

Extension and subsidence between the Iberian and European plates is recorded since Triassic times as the Iberian plate lay between the opening North Atlantic and Tethys oceans [START_REF] Biteau | he Aquitaine Basin[END_REF]. he principal rit phase occurred in the Early Cretaceous. he plate kinematics of this phase remain diicult to constrain, resulting in several competing models (Figure 1.12) [Olivet, 1996;Sibuet et al., 2004;Jammes et al., 2009]. hese models predict transtension, strike-slip or orthogonal extension, broadly categorised into three competing models: 1) A strike slip model that implies several hundred kilometres of sinistral strikeslip accommodated along the North Pyrenean Fault [Olivet, 1996]. 2) A scissor type model that implies opening of the Bay of Biscay in the west contemporaneous with oceanic subduction in the east [Srivastava et al., 1990;Roest and Srivastava, 1991;Sibuet et al., 2004;Vissers andMeijer, 2012a, 2012b], and a hybrid model that combines an anticlockwise rotation and eastward movement of Iberia with transtension between Iberia and Europe, followed by orthogonal extension during the Aptian and orthogonal convergence from the Late Cretaceous onwards [Jammes et al., 2009].

Regardless of the plate kinematics, the presence of lherzolites in the Metamorphic Internal Zone shows that mantle was exhumed during the pre-orogenic rit phase. he metamorphosed Mesozoic sediments are assumed to have covered the exhumed mantle during this rit phase (Figure 1.13). his is supported by roughly coeval ages for riting and metamorphism [Golberg and Leyreloup, 1990;[START_REF] Clerc | Amincissement crustal extreme et thermicité élévee: un analogue pour les marges passives[END_REF] and syn-rit breccias with Lherzolite clasts [Jammes et al., 2009;Lagabrielle et al., 2010;de Saint Blanquat et al., 2016]. Triassic Keuper salt may also have provided a décollement for the Mesozoic cover to slide into the rit [START_REF] Jammes | Interaction between prerit salt and detachment faulting in hyperextended rit systems: he example of the Parentis and Mauléon basins (Bay of Biscay and western Pyrenees)[END_REF]Teixell et al., 2016;Saura et al., 2016]. 1.3.3 Convergent evolution he onset of convergence is recorded by lexural foreland basins and syn-depositional thrusts in both forelands starting around 85-80 Ma [Desegaulx et al., 1990;[START_REF] Deramond | Relationships between thrust tectonics and sequence stratigraphy surfaces in foredeeps: model and examples from the Pyrenees (Cretaceous-Eocene, France, Spain)[END_REF]. he onset of convergence is usually placed at late Santonian to early Campanian (~83 Ma) [Puigdefàbregas and Souquet, 1986;Mouthereau et al., 2014]. he Metamorphic Internal Zone was emplaced northward (and likely also southward) during closure of the exhumed mantle domain [Jammes et al., 2009;Lagabrielle et al., 2010]. Marine lexural basins formed on both the Iberian and European margins. In the northern Pyrenees, paleocurrent data and the westward prograding geometry of Upper Cretaceous lysch indicates that sediments were supplied from the east (Figure 1.14) [Bilotte, 1985]. In the southern Pyrenees, sediments were supplied from the south and east [START_REF] Rosell | El "Garumniense" Prepirenaico[END_REF]. Paleocene sedimentation was continental throughout the Pyrenees, and tectonic activity appears to have been slow or nonexistent [START_REF] Plaziat | Late Cretaceous to late Eocene paleogeographic evolution of southwest Europe, Palaeogeogr[END_REF]Desegaulx and Brunet, 1990;Vergés et al., 2002]. his phase of slow shortening may be due to a temporary halt in plate convergence, an idea based on plate kinematic reconstructions [Rosenbaum et al., 2002], but this has been contested by other authors [Vissers and Meijer, 2012a]. Following a marine transgression from the west in the Ypresian, the south Pyrenean foreland developed as underilled and marine from 55 to 37 Ma [Puigdefàbregas and Souquet, 1986;[START_REF] Puigdefàbregas | hrusting and foreland basin evolution in the Southern Pyrenees, in hrust Tectonics[END_REF]Vergés et al., 1995Vergés et al., , 1998]. he transition to continental conditions in the southern Pyrenees occurred because the Ebro foreland basin was isolated from the Atlantic Ocean ater uplit of the Western Pyrenees [START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF]. In the northeastern Pyrenees, the Ypresian marine transgression was followed by a westward regression and deposition of continental sandstones and conglomerates [START_REF] Martin-Martin | Tectonic controls on the deposits of a foreland basin: an example from the Eocene Corbières-Minervois basin, France[END_REF]Christophoul et al., 2003]. Syn-depositional deformation was recorded in both the north and south Pyrenean forelands [Ford et al., 1997;[START_REF] Martin-Martin | Tectonic controls on the deposits of a foreland basin: an example from the Eocene Corbières-Minervois basin, France[END_REF]Ramos et al., 2002;Christophoul et al., 2003;[START_REF] Labaume | Tectono-thermal history of an exhumed thrust-sheet-top basin: an example from the south Pyrenean thrust belt[END_REF]. Uplit and exhumation of the Axial Zone starting around 50 Ma is recorded by both in-situ and detrital thermochronology data [Fitzgerald et al., 1999;Sinclair et al., 2005;Metcalf et al., 2009;Rahl et al., 2011;Whitchurch et al., 2011]. he peak of exhumation of the Axial Zone appears to have occurred in the Early Oligocene (~36 Ma), simultaneous with a large inlux of sediment into the south Pyrenean foreland [Fitzgerald et al., 1999;Sinclair et al., 2005;Fillon et al., 2013a]. Exhumation migrated both southward and westward in the Pyrenees [Sinclair et al., 2005;Jolivet et al., 2007;Metcalf et al., 2009]. he end of deformation in the Pyrenees is placed in the Late Oligocene (~24.7 Ma) by magnetostratigraphic dating of growth strata along the south Pyrenean thrust front [START_REF] Meigs | Sequential development of selected Pyrenean thrust faults[END_REF]. Following the end of convergence, post-orogenic exhumation of the Ebro foreland basin remained relatively minor until 9 Ma, most likely because of a drop in base level ater a connection to the Atlantic Ocean was re-established [Fillon and van der Beek, 2012]. 1.3.4 Shortening estimates Estimates of the amount of shortening in the Pyrenees vary. Before the acquisition of deep seismic data, minimum shortening was estimated between 55 km in the west to 80 km in the east [START_REF] Seguret | Crustal scale balanced cross-sections of the Pyrenees; discussion[END_REF]. Following the irst results of the ECORS deep seismic proile, shortening was estimated at a minimum, but poorly constrained, 120 km [Roure et al., 1989]. Later, better constrained balanced sections along the ECORS line placed shortening at 147 km and 165 km, using upper crustal imbricates in the Axial Zone and subducting Iberian lower crust [Muñoz, 1992;Beaumont et al., 2000]. he latter estimate is partitioned between 128 km for the southern wedge and 37 km for the northern one [Beaumont et al., 2000]. More recently, a much lower estimate of 92 km of shortening (excluding closure of the exhumed [START_REF] Ricateau | Evolution au Crétacé supérieur de la pente séparant le domaine de plate-forme du sillon sous-pyrénéen en Aquitaine méridionale[END_REF] and Bilotte [1985].

Chapter 1. Introduction 24 mantle domain) was obtained, resulting from full crustal imbrication and a revised shortening estimate of the southern Pyrenees [Mouthereau et al., 2014].

Shortening in the Eastern Pyrenees is estimated at 125 km [Vergés et al., 1995]. his is partitioned as 32 km in the northern Pyrenees [Baby et al., 1988], ~23 km of internal shortening of the Axial Zone, and 70 km in the south Pyrenean foreland [Vergés, 1993]. In the Western Pyrenees, shortening is estimated to be lower, at 75 to 80 km [Teixell, 1998], and more recently 115 km including closure of a 15 km wide exhumed mantle zone [Teixell et al., 2016].

he evolution of shortening rates in the Pyrenees remains debated. Several diferent shortening histories have been proposed, some of them contradicting others (Figure 1.15). In the Central Pyrenees, Beaumont et al., [2000] show an overall acceleration of the shortening rate, in particular at 36 Ma. his is based around a peak in exhumation rates of the Axial Zone around that time [e.g., Fitzgerald et al., 1999;Metcalf et al., 2009]. In contrast, Mouthereau et al., [2014] propose a deceleration of shortening, starting at ~3 mm/y in the Late Cretaceous and Paleocene and gradually decelerating until the Miocene. In the Western Pyrenees, shortening rates are estimated to have peaked in the Eocene at 3-2.5 mm/y [Teixell et al., 2016]. In the Eastern Pyrenees, a shortening rate history exists only for the southern foreland. his also places the peak of shortening during the Eocene, at ~4.5 mm/y [Vergés et al., 1995].

Limitations and scientiic questions

Critical taper theory is useful to describe the average shape and deformation of an orogenic wedge. However, it does not describe the episodic nature of deformation along discrete fault planes as observed in modelled thrust wedges [Hoth et al., 2007;[START_REF] Simpson | Mechanics of non-critical fold-thrust belts based on inite element models[END_REF]Buiter, 2012]. In doubly vergent models, it has been shown that pro-and retro-wedges and pro-and retro-foreland basins behave diferently [Willett et al., 1993;Naylor and Sinclair, 2008]. Analogue models have hinted at a possible mechanical coupling between the pro-and retro-wedge, where shortening alternates between them [Hoth et al., 2007]. his behaviour cannot be explained with critical taper theory alone, and is currently poorly understood. he relationship between the pro-and retro-wedge, whether that relationship changes through time, and what would cause such a change remains poorly constrained in natural systems. Detailed comparisons of the two wedges and their foreland basins in terms of structure and evolution are required to investigate these issues. From these limitations of our understanding of doubly vergent orogens come the following main scientiic questions that are investigated in this thesis.

• How does a doubly vergent orogen evolve?

• Do the pro-and retro-wedge interact?

• What factors inluence shortening distribution, and how?

Associated with these questions are secondary questions: What causes the Paleocene quiescence in the Pyrenees? Can thin-skinned deformation inluence the crustal structure of an orogen?

In this work, a dual approach is used that combines detailed study of the Eastern Pyrenees and numerical modelling of controlling factors to constrain the evolution of doubly vergent orogens. Both methods were used in parallel, so that insights from one may beneit the other. he results of the East Pyrenean case study are described in Chapter 2 (in revision ater review at Tectonics). he results of numerical modelling are described in Chapters 3 and 4 (destined for submission to Geology and Journal of Geophysical Research: Solid Earth, respectively). Based on the results of these studies, a generic evolutionary model for inverted rit systems is developed and compared against several natural doubly vergent orogens, also in Chapter 4.

Introduction

It is common in foreland basin modelling and fold-and-thrust belt studies to analyse only one side of a doubly vergent orogen [e.g., [START_REF] Ershov | Northern Caucasus basin: hermal history and synthesis of subsidence models[END_REF]Fillon et al., 2013;[START_REF] Decarlis | Salt tectonics in the SW Alps (Italy-France): From riting to the inversion of the European continental margin in a context of oblique convergence[END_REF]. However several studies, mainly based on mechanical modelling, suggest that pro-and retro-wedges interact in accommodating plate convergence [Willett et al., 1993;Sinclair et al., 2005;Hoth et al., 2007Hoth et al., , 2008;;Erdős et al., 2015]. Correlating the behaviour of the two wedges in doubly vergent orogens can therefore provide important insight into orogen dynamics. How deformation and subsidence of the pro-and retro-foreland of a natural system relate to each other, whether this relationship changes through time, and what would cause such a change remains unclear. Detailed comparisons of the two wedges and their foreland basins in terms of structure and evolution are required to investigate these issues.

In this paper we investigate the Eastern Pyrenees, comparing and correlating the northern retro-foreland and southern pro-foreland along two sections directly opposite each other to better understand full orogen evolution (Figure 2.1). he South Pyrenean fold-and-thrust belt and Ebro foreland basin system is extremely well documented owing to excellent outcrop conditions and well-deined stratigraphy [e.g., Burbank et al., 1992b;Vergés et al., 1992;[START_REF] López-Blanco | Tectonic and climatic controls on the development of foreland fan deltas: Montserrat and Sant Llorenç del Munt systems (Middle Eocene, Ebro Basin, NE Spain)[END_REF][START_REF] Gómez-Paccard | Tectonic and climatic controls on the sequential arrangement of an alluvial fan/fan-delta complex (Montserrat, Eocene, Ebro Basin, NE Spain)[END_REF]. It developed on the subducting Iberian plate starting in the Late Cretaceous. Subsidence records show locally short-lived and accelerating subsidence owing to constant outward basin migration and incorporation of the proximal basin into the thrust wedge. his corresponds to simulated behaviour of kinematic and mechanical models of pro-forelands [Beaumont et al., 2000;Naylor and Sinclair, 2008;Sinclair and Naylor, 2012;Fillon et al., 2013;Erdős et al., 2014]. In contrast, owing to poorer outcrop conditions and structural and stratigraphic complexity, our understanding of the North Pyrenean retro-wedge is less well developed despite a large body of work spanning several decades [e.g., Choukroune, 1974;Souquet et al., 1977;Puigdefàbregas and Souquet, 1986;Lagabrielle et al., 2010;[START_REF] Clerc | High-temperature metamorphism during extreme thinning of the continental crust: A reappraisal of the North Pyrenean passive paleomargin[END_REF]. he retro-wedge preserves an inverted hyperextended rit system of Early Cretaceous age, as recorded by the presence of mantle blocks and synextension high-temperature metamorphism in the internal, southern part of the retro-wedge [Jammes et al., 2009;Lagabrielle et al., 2010;Mouthereau et al., 2014;[START_REF] Clerc | High-temperature metamorphism during extreme thinning of the continental crust: A reappraisal of the North Pyrenean passive paleomargin[END_REF][START_REF] Vacherat | Rit-to-collision transition recorded by tectonothermal evolution of the northern Pyrenees[END_REF]. In general, retro-wedges can be expected to accommodate less shortening than pro-wedges, resulting in slow and long-lived foreland basin subsidence [Sinclair and Naylor, 2012]. Previous work proposes that the western Aquitaine Basin follows this predicted pattern [Desegaulx and Brunet, 1990;Naylor and Sinclair, 2008], however more recent work shows that the eastern retro-foreland basin has a more complex history that challenges these models, including a temporary stagnation in subsidence for which the cause remains unclear [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF].

To be able to make a detailed comparison of the pro-and retro-forelands, we studied the tectonic and stratigraphic evolution of the North Pyrenean Zone and Aquitaine foreland basin in the Eastern Pyrenees (Ariège, France; Figure 2.1). Our work is based on a structural analysis and cross section balancing that integrated newly acquired ield data, seismic relection proiles, borehole data, and published stratigraphic and structural data [e.g., Choukroune, 1974;Baby, 1988;Bilotte et al., 1988b]. his allowed us to decipher the geometry, timing, rates, and kinematics of deformation and to identify and characterize the various phases, styles, and amounts of Santonian to Oligocene shortening in the eastern Pyrenean retro-wedge and its foreland basin. We then correlated this complete retro-wedge history with its pro-wedge counterpart in the Spanish Pyrenees using the work of Vergés [1993], Vergés et al. [1995Vergés et al. [ , 1998Vergés et al. [ , 2002]], and [START_REF] Vergés | Mesozoic evolution and Cainozoic inversion of the Pyrenean rit, in Peri-Tethys Memoir 6: Peri-Tethyan Rit/Wrench Basins and Passive Margins[END_REF]. We identiied key diferences and point out possible interactions between the pro-and retro-foreland fold belts. hen, we investigate how shortening was partitioned in time and space between pro-and retro-wedge and discuss what controls the sequence of deformation during orogeny by presenting a new crustal-scale model for the evolution of the Eastern Pyrenees.

Geological Setting

2.2.1 Plate Kinematics he Pyrenees formed as part of the Alpine belt due to the collision of the Iberian microplate with the European continent. Before the collision, an Aptian to early Cenomanian oblique rit system associated with exhumation of mantle rocks formed between the two plates [Debroas, 1990;Jammes et al., 2009;Lagabrielle et al., 2010]. Riting has been related to the opening North Atlantic Ocean [Srivastava et al., 1990;Olivet, 1996;Sibuet et al., 2004]. No consensus on plate kinematics around Iberia yet exists and there are several competing models. Most models are based on reconstructions of magnetic sea-loor anomalies of the Bay of Biscay, Central, and North Atlantic Ocean [Olivet, 1996;Sibuet et al., 2004;Vissers andMeijer, 2012a, 2012b]. hese models predict diferent extensional histories for the Pyrenean realm, and, ater the onset of convergence (~84 Ma), predict diferent degrees of obliquity for N-S shortening. Pyrenean tectonic structures predominantly record a N-S shortening direction and a minor strike-slip component that was mainly focussed along the North Pyrenean Fault [Souquet et al., 1977;Choukroune and Mattauer, 1978].

Tectonic Zones and Major Boundaries

In the Eastern Pyrenees the orogen is divided into six tectonostratigraphic zones (Figure 2.1b). From north to south, the Aquitaine foreland basin is followed by the sub Pyrenean Zone, a folded and faulted zone north of the North Pyrenean Frontal hrust, usually delimited by a blind thrust [Souquet et al., 1977;[START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF] Zone, a narrow, north verging fold-and-thrust belt. In the study area this zone comprises: inverted rit basins (e.g., Fougax and Roquefeuil blocks; Figure 2.2), inverted basement massifs (e.g., Saint-Barthélémy Massif; Figure 2.2), and the Metamorphic Internal Zone (Figure 2.2) [Souquet et al., 1977]. he North Pyrenean Fault separates the North Pyrenean Zone from the Axial Zone and is traditionally seen as the suture between Europe and Iberia. In the classic interpretation of the ECORS section lying some 50 km to the west (see below) [e.g., Muñoz, 1992] crustal thickening in the Axial Zone is related to a south verging, antiformal stack of Iberian crust. Interpretations in the eastern and western Pyrenees show almost no Alpine folding of Paleozoic basement rocks of the Axial Zone [Vergés et al., 2002;Teixell et al., 2016]. South of the Axial Zone lies the South Pyrenean Zone, a south verging thin-skinned fold-and-thrust belt with syn-orogenic thrust sheet-top basins. he Ebro Basin forms the pro-foreland basin of the Pyrenees. It is an asymmetric double lexural basin with a central intrabasinal high due to loading by the Pyrenees in the north and the Catalan Coastal Ranges in the south.

he deep crustal structure of the Central Pyrenees is imaged by the ECORS Pyrenees deep seismic study (Figure 2.1b) [Roure et al., 1989;Muñoz, 1992;Roure and Choukroune, 1998] and more recently by seismic tomography studies [START_REF] Souriau | A new tomographic image of the Pyrenean lithosphere from teleseismic data[END_REF][START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF][START_REF] Wang | he deep roots of the western Pyrenees revealed by full waveform inversion of teleseismic P waves[END_REF]. hese data show the Iberian Moho underthrusting the European plate, deepening northward from 35 km to 50 km depth [Beaumont et al., 2000]. Most authors show the European Moho at a fairly constant 30 km depth along the ECORS line [Beaumont et al., 2000;[START_REF] Díaz | Crustal structure beneath the Iberian Peninsula and surrounding waters: A new compilation of deep seismic sounding results[END_REF] and some show a gentle southward dip [Roure et al., 1989;Roure and Choukroune, 1998]. Recent seismic tomography data show the subduction interface dipping 20° northward and show the European Moho rising southward [START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF].

he nature and extent of Alpine deformation in the Axial Zone, and the role of inherited Variscan shear zones are currently subject to renewed debate and analysis with modern tools [e.g. [START_REF] Laumonier | he Alpine southeastern Pyrenees (France, Spain); a synthesis[END_REF][START_REF] Cochelin | Structural inheritance in the Central Pyrenees: the Variscan to Alpine tectonometamorphic evolution of the Axial Zone[END_REF]. In this paper we focus on the fold-and-thrust belts and foreland basins on both sides of the Eastern Pyrenees. As a detailed analysis of the Axial Zone is beyond the scope of this paper, we use the Axial Zone model of Vergés et al. [1995Vergés et al. [ , 2002] ] and [START_REF] Vergés | Mesozoic evolution and Cainozoic inversion of the Pyrenean rit, in Peri-Tethys Memoir 6: Peri-Tethyan Rit/Wrench Basins and Passive Margins[END_REF], which remains valid.

Methods

In order to investigate the structural evolution of the retro-foreland, we constructed two cross sections, the main section S3 and an auxiliary section S3b (for location see Figures 2.1a, 2.2). hese sections were positioned directly opposite cross section J3 of Vergés [1993] in the pro-foreland. Standard balancing techniques were used to restore the sections [START_REF] Woodward | Balanced Geological Cross-Sections: An Essential Technique in Geological Research and Exploration[END_REF]. While most sedimentary strata were line-length balanced, weak units with signiicant internal deformation were also area balanced. We used the restored cross sections combined with stratigraphic constraints on timing of activity to estimate average shortening rates for each map-scale structure across the pro-and retro-forelands of the Eastern Pyrenees. he minimum overall shortening rate for a given period was derived by summing the shortening rates for all unlinked structures active during that period. A complete high resolution shortening history across the whole orogen can thus be constructed. Such an analysis integrating detailed map-scale structural history is especially important in areas that record very slow and distributed shortening like the Pyrenees. his shortening history was then used to constrain the crustal scale restoration model, together with subsidence histories, thermochronology data, and deep geophysical data.

he retro-foreland database includes surface geology from 1:50,000 geological maps of the French Geological Survey (BRGM; 1036 Castelnaudary [Cavaillé et al., 1975a]; 1058 Mirepoix [Cavaillé, 1976a]; 1075 Foix [Bilotte et al., 1988a]; 1076 Lavelanet [Bilotte et al., 1988b]), an unpublished geological map [START_REF] Marty | Relations géologiques entre le massif du Saint-Barthélémy et les séries post-hercyniennes du pays de Sault (Pyrénées ariégeoises)[END_REF], publicly available borehole data (Bureau Exploration-Production des Hydrocarbures; beph.net), limited seismic relection data provided by the BRGM, a 90 m SRTM Digital Elevation Model processed and distributed by the CGIAR Consortium for Spatial Information (srtm.csi.cgiar.org) [START_REF] Jarvis | Hole-illed SRTM for the globe Version 4, available from the CGIAR-CSI SRTM 90m Database[END_REF], and our own ield data. he stratigraphic framework used for the retro-foreland was established by the ANR PYRAMID research project [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF] and linked to the 2013/01 chronostratigraphic chart published by the International Commission on Stratigraphy [START_REF] Cohen | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF]. Details on stratigraphic nomenclature and equivalent BRGM units are presented in Table A.1 (Supplementary material). he litho-and chronostratigraphy for the pro-foreland are based on Vergés [1993] and Vergés et al. [1998Vergés et al. [ , 2002]].

Total and tectonic subsidence curves for four wells in the retro-foreland are compared with those for six wells or sections in the pro-foreland [Vergés et al., 1998], all calibrated to present-day sea level. We used standard methods of decompaction and backstripping assuming Airy isostasy [Steckler and Watts, 1978]. Decompaction calculations integrate the values for initial porosity and compaction coeicient for each lithology as used in Vergés et al. [1998]. For the backstripping process each unit was given bulk values for those constants based on the percentages of each lithology in that particular unit (Table A.2). hese percentages were determined from the highest resolution sections available, ranging in scale between 1:500 and 1:2000. Estimates for paleo-bathymetry were based on depositional environments. Continental deposits were considered to be equivalent to water-illed basins with 0 or -50 m bathymetry. We used a mantle density of 3300 kg/m 3 and a density of 1000 kg/m 3 for water. he original published subsidence [START_REF] Marty | Relations géologiques entre le massif du Saint-Barthélémy et les séries post-hercyniennes du pays de Sault (Pyrénées ariégeoises)[END_REF], and de Saint Blanquat et al. [2016]. Boreholes are located between the North Pyrenean Frontal hrust and the Lavelanet Anticline. he blue lines indicate partial section traces of S3 and S3b. NPFT, North Pyrenean Frontal hrust; 3MF, 3M Fault; 3MF(1), lat of 3M Fault; 3MF(2), hanging wall shortcut of 3M Fault; NPF, North Pyrenean Fault. calculations for the pro-foreland did not take into account eustatic sea level changes [Vergés et al., 1998]. hese data were therefore converted using the same eustatic data as used for the retro-foreland [START_REF] Snedden | A Compilation of Phanerozoic Sea-Level Change , Coastal Onlaps and Recommended Sequence Designations[END_REF]. In displaying total and tectonic subsidence relative to present-day sea level, many curves start above 0 m because paleo-sea level was between 160 and 210 m higher than present day and paleobathymetry in the forelands was relatively minor. In the case of the pro-foreland, including the long term eustatic sea level curve may introduce artefacts in the subsidence. Brief periods of low bathymetry in the Eocene are shown as tectonic uplit events in our converted curves, but the observed reduction in bathymetry may have been caused by a short-lived deviation from the long term sea level curve instead [Burbank et al., 1992b]. Errors in the estimation of bathymetry have the greatest efect on calculated subsidence. Errors in sea level and sediment ages have less impact. South of the North Pyrenean Frontal hrust, Triassic evaporites (Keuper Group; unconstrained thickness) lie at the base of the Roquefeuil and Fougax block successions [START_REF] Marty | Relations géologiques entre le massif du Saint-Barthélémy et les séries post-hercyniennes du pays de Sault (Pyrénées ariégeoises)[END_REF]. Jurassic limestones and dolomites occur across the northern North Pyrenean Zone (Black Dolomite Group; 100 to 200 m; Figure 2.2). Berriasian to Barremian limestones are present in the Roquefeuil block (1000 m, Mirande Group, Figure 2.3b) but absent in the Fougax and Nalzen blocks (Figures 2.4,2.5). hese strata record littoral environments, with occasional lagoonal and distal platform deposits [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF].

North Pyrenean Foreland

In the Metamorphic Internal Zone (Figure 2.2) [Choukroune, 1974] protolith carbonates and siliciclastic strata are interpreted as Jurassic to Early Cretaceous in age [Golberg and Leyreloup, 1990;Lagabrielle et al., 2010] but their true age remains unclear owing to Albian to Cenomanian high-temperature, low-pressure metamorphism and strong internal deformation [Golberg and Leyreloup, 1990]. Widespread breccias predominantly comprise clasts of marble and hornfels, but also rare clasts of basement rocks and mantle lithologies [de Saint Blanquat et al., 2016].

Syn-rit deep marine shales of the Black Flysch Group (~1000 m; Aptian to Lower Cenomanian) are in stratigraphic continuity with the Mirande Group limestones on the southern Fougax block (Figures 2.2, 2.5), while on the northern side they rest unconformably on the Jurassic Black Dolomite Group [Bilotte et al., 1988b]. he term 'lysch' is traditionally used in the name of the Mid-Albian to Lower Cenomanian 'Black Flysch' Formation, despite its pre-orogenic origin [Souquet et al., 1985;Debroas, 1990]. Here we apply the name 'Black Flysch Group' to a wider age range due to very similar lithologies of Aptian to Lower Albian age. Black Flysch shales transition northward into rudist platform limestones (Urgonian facies) of equivalent age (Figure 2.3b) and of variable thickness (maximum ~600 m) [Bilotte et al., 1988b]. he Black Flysch Group thins westward where it is preserved in the Nalzen block north of the Saint-Barthélémy Massif (~600 m; Figure 2.2), thus recording stratigraphic continuity between the Fougax and Nalzen blocks. here, these shales contain marine conglomerates with locally sourced basement clasts (Conglomérat de Freychenet) [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF].

he post-rit Grey Flysch Group succession (Cenomanian to Santonian) records a deepening upward trend from platform carbonates to deep-water marls (borehole Dr4; Figure 2.3b). South of the North Pyrenean Frontal hrust, it is only preserved in the Nalzen block (Figure 2.2; 1000 to 1500 m). To the north it directly overlies a partially inverted half-graben and Variscan basement (Figure 2.5) and thins northward to pinch out north of the Lavelanet Anticline (Figure 2.4c). here, Santonian strata are either missing or indistinguishable from the overlying syn-orogenic sediments. Santonian sediments are unlikely to have been eroded completely, given the marine depositional environment. herefore, we interpret this as local non-deposition, recording southward olap of the upper Grey Flysch Group.

he Petites Pyrénées Group represents the early syn-orogenic succession (Lower to Middle Campanian; Figures 2.3b, 2.4, 2.5). South of the North Pyrenean Frontal hrust the succession reaches a minimum thickness of ~500 m (Nalzen block) and records shallower and more sandy facies than further north. North of the Frontal hrust, the group thins northward from 1500 -2500 m to pinch out in the Tréziers Anticline (Figures 2.3b,2.4). Internal deformation is responsible for thickening of the Petites Pyrénées Group below the Lavelanet Anticline. he northern series records a shallowing upward trend from deep marine shales (with secondary siltstones and sandstones) representing delta progradation from east to west [START_REF] Ricateau | Evolution au Crétacé supérieur de la pente séparant le domaine de plate-forme du sillon sous-pyrénéen en Aquitaine méridionale[END_REF]Bilotte, 1985]. It is overlain by and laterally equivalent to the luviodeltaic Plantaurel Group (Late Campanian to middle Maastrichtian) that was supplied from the east with rare marine incursions from the west [Bilotte, 1985;[START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF]. he Plantaurel Group thickens northward from 300 m to 1300 m across the syn-depositionally inverted Benaix Fault and then gradually thins northward to 700 m in the Lavelanet Anticline, reaching 430 m in the Tréziers Anticline before pinching out around Orsans (Figures 2.3b,2.4).

North of the North Pyrenean Frontal hrust, red continental muds, sandstones, and intercalated conglomerates of the Aude Valley Group (Garumnian facies; middle Maastrichtian to hanetian) unconformably overlie the Plantaurel (Lavelanet Anticline; Figure 2.2) [Bilotte et al., 1988b] and Petites

Pyrénées Groups (near Benaix; Figure 2.2). he Maastrichtian conglomerates contain pebbles of Upper Cretaceous limestones [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF], recording exhumation and erosion in the hinterland. In borehole Dr4 locally developed Danian to Selandian lacustrine limestones and marls are only ~95 m thick (Table A .3). A brief hanetian marine incursion is recorded in the southern half of the basin by internal platform limestones of the Rieubach Group, which is the western equivalent of, and interingers with, the continental Aude Valley Group [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF]. Growth strata only appear in the upper hanetian portion of the succession (Figure 2.4c). Overall, the Aude Valley Group thins northward from 700 m south of the Lavelanet Anticline to pinch out just south of borehole MRL1 (Figures 2.3b,2.4).

A more signiicant but brief marine incursion in the early Ypresian (Ilerdian) is recorded across the entire eastern foreland basin at the base of the Coustouge Group [START_REF] Cavaillé | Notice explicative de la feuille Castelnaudary à 1/50 000 (Feuille №1036)[END_REF]Cavaillé, 1976a;[START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF]. his group then records a shallowing upward trend with marine marls and limestones passing up into mudstones and nummulitic sandstones [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF]Tambareau et al., 1995]. hicknesses vary considerably due to syn-sedimentary deformation: from ~700 m immediately north of the Lavelenet Anticline to ~450 m in the Tréziers Anticline, ~500 m in the footwall of the Tréziers hrust, to ~100 m at the northern limit of the preserved foreland basin (borehole MRL1). he northern pinchout is not preserved (Figure 2.1a). he Late Ypresian to Rupelian Carcassonne Group covers the entire foreland north of the Lavelanet Anticline, with the main depocentre in the footwall of the Tréziers hrust (Figure 2.4). hese sediments comprise mainly luvial sandstones. Proximal conglomerates along the thrust front include clasts of crystalline basement, scapolite-bearing limestones, impure marble with garnets, and nonmetamorphic Mesozoic sediments, sourced from the south [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF][START_REF] Crochet | Molasses syntectoniques du versant nord des Pyrénées: la série de Palassou[END_REF].

2.4.2 North Pyrenean Structure he structure of the eastern North Pyrenean Zone in the Saint-Barthélémy -Fougax area is here described along two restored cross-sections (S3 and S3b; Figures 2. 4, 2.5). Restoration provides estimates of N-S shortening across the whole northern foreland of 13.3 km (S3) and 13.2 km (S3b), excluding emplacement of the Metamorphic Internal Zone (minimum ~5 km). his is lower than previous estimates of around 25-30 km that also excluded emplacement of the Metamorphic Internal Zone [Souquet and Peybernès, 1987;Baby, 1988;Baby et al., 1988]. Most of this diference comes from a basement duplex proposed below the Saint-Barthélémy Massif that is not necessary to it our data, and thus is absent in our interpretation.

In the Aquitaine Basin the sedimentary inill is gently deformed by the basement-rooted Tréziers and Orsans hrusts (Figures 2.4a,d). Both structures are the western terminations of thrusts accommodating large anticlinal basement uplits in the eastern foreland: the Mouthoumet [Cavaillé, 1976a;Bilotte et al., 1988b;[START_REF] Crochet | Carte géologique de la France au 1/50 000: Quillan (Feuille №1077)[END_REF] and Alaric thrusts respectively [START_REF] Ellenberger | Carte géologique de la France au 1/50 000: Capendu (Feuille №1060)[END_REF]Christophoul et al., 2003].

he WNW-ESE trending Lavelanet Anticline is the most prominent surface feature of the sub Pyrenean Zone (Figure 2.2). Boreholes (Dr3, Dr4) reveal a 1150 m diference in basement depth below the southern lank (Figure 2.4c), interpreted here as due to a south-verging thrust (Lavelanet Backthrust) [Baby, 1988]. Repetition of stratigraphy in borehole Bx1 implies the existence of two other blind, north-verging thrusts (Benaix Fault and Lavelanet hrust; Figure 2.4c). he Benaix Fault is a partially inverted normal fault delimiting a Lower Cretaceous rit basin that is equivalent to the Camarade basin to the west [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF].

he North Pyrenean Frontal hrust marks an abrupt southward increase in deformation intensity (Figures 2.4,2.5). he fault is here interpreted as an inverted normal fault with the deformed Nalzen-Fougax basins in its hanging wall. he Nalzen succession (S3; Figure 2.4) forms the vertical to overturned limb of a major north verging anticline cored by the Saint-Barthélémy basement (Figure 2.4c). We interpret this structure as a basement-cored forced fold, similar to Laramide structures with a larger fault ofset [e.g., Mitra and Mount, 1998]. Several hanging wall splays of the main North Pyrenean Frontal hrust cut the Nalzen block. his basement-cored forced fold transitions into a recumbent fold with no basement ~5 km to the east, on section S3b (Figure 2.5), where the Fougax block is thrust over the overturned frontal limb. In the Fougax block, Black Dolomite and Black Flysch strata are intensely folded and cleaved (Figure 2.5) [START_REF] Marty | Relations géologiques entre le massif du Saint-Barthélémy et les séries post-hercyniennes du pays de Sault (Pyrénées ariégeoises)[END_REF]. he E-W trending, tight, upright folds show a short wavelength of ~750 to 850 m and relatively low amplitudes of ~500 to 600 m. Using methods described in [START_REF] Bulnes | Estimating the detachment depth in cross sections involving detachment folds[END_REF] we use the fold geometry to infer a depth to décollement of approximately 1 to 1.5 km. his suggests that the Fougax and Nalzen successions were detached from their basement (probably along Keuper evaporites [Marty,13.2 1976]), folded and thrust northward above the North Pyrenean Frontal hrust. Baby [1988] proposes that the Nalzen basin is immediately underlain by a basement high, implying the same for the Fougax Basin. Given the diference between the estimated depth of the Fougax decollement and the observed depth of basement in boreholes further north, this model requires a throw of 3 km on an unknown north vergent thrust. he two blind thrusts observed in borehole Bx1 do not accommodate suicient shortening to produce the required displacement. In addition, our interpretation of the Nalzen Block means it was cut from the footwall of the Roquefeuil Fault, thus Black Flysch Group strata should also be present at depth in the footwall of the North Pyrenean Frontal hrust. We therefore extrapolated the foreland basin succession in borehole Bx1 southward (~4 km) to underlie the allochthonous Fougax and Nalzen blocks and overlie a partially inverted rit basin that is equivalent to the Camarade basin to the west (Figures 2.4c, 2.5c) [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF]. 2.2) [Bilotte et al., 1988b] across the same lateral ramp. his lateral ramp is interpreted here as a reactivated Cretaceous extensional structure that may also be responsible for the westward thinning of the Lower Cretaceous sedimentary cover of the Saint-Barthélémy Massif. he massif has previously been interpreted as an allochthonous unit that originally formed a basement high south of the Nalzen block [Souquet and Peybernès, 1987;Baby, 1988;Baby et al., 1988]. Apatite ission track data from the adjacent Arize Massif imply that the massifs were buried before the Eocene [START_REF] Vacherat | Rit-to-collision transition recorded by tectonothermal evolution of the northern Pyrenees[END_REF]. Assuming a standard geothermal gradient of 30 °C/km and closure temperature of 100-120 °C [START_REF] Reiners | Zircon (U-h)/He hermochronometry[END_REF], the massifs were buried to at least 3-4 km by Mesozoic sedimentary cover (Black Flysch Group, Grey Flysch Group, and possibly Petites Pyrénées Group). his is consistent with our restoration that places the massif below the ~4 km thick Nalzen block, which requires less shortening than earlier restorations [Baby, 1988;Baby et al., 1988].

Folds of the

he amount of Pyrenean shortening accommodated in the Metamorphic Internal Zone and its displacement along the 3M Fault (Figures 2.2, 2.4, 2.5, 2.6) [START_REF] Marty | Relations géologiques entre le massif du Saint-Barthélémy et les séries post-hercyniennes du pays de Sault (Pyrénées ariégeoises)[END_REF] A .3). Syn-orogenic subsidence curves display a coherent trend across the foreland with two clear periods of subsidence separated by a quiescent phase (Figure 2.7c, d). Overall syn-orogenic tectonic subsidence is characterised by very low rates with only 0.06 mm/y on average for Dr4, which records the fastest subsidence. A sharp increase in subsidence rate at ~84 Ma marks the onset of convergence, reaching a maximum of 0.14 mm/y during the early Campanian. Subsidence slowed down ater 75 Ma to 0.02 mm/y and 0.04 mm/y in boreholes Dr4 and T1, respectively. Little or no subsidence was recorded throughout the foreland from ~66 to ~59 Ma (quiescent phase). he subsidence shown in the synthetic borehole syn1 during the irst half of this phase is probably an artefact due to poor age control. Renewed subsidence started in the more distal basin during the hanetian (~59 Ma, Figure 2.7c, d) and is characterised by lower tectonic subsidence rates than during the Late Cretaceous (0.03 mm/y and 0.09 mm/y on average, respectively). hese long-term averages are based on tectonic subsidence of 1530 m in 18 My (borehole Dr4) and 244 m in 7.3 My (borehole T1). Borehole syn1 shows a gradual slowing of subsidence towards the end of the Eocene. 2.4.4 Evolution of the North Pyrenean Foreland he model presented here for the evolution of the North Pyrenean foreland in the Saint-Barthélémy -Fougax area is based on the restorations of the cross sections S3 and S3b (Figures 2.4, 2.5), foreland basin subsidence history, and the construction of three lateral cross sections linking the two present-day cross sections (Figure 2.6). he age boundaries of the evolutionary phases correspond to rounded boundaries of international chronostratigraphic ages due to limitations in the available stratigraphic dating [START_REF] Cavaillé | Notice explicative de la feuille Mirepoix à 1/50 000 (Feuille №1058)[END_REF][START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF]. [Gouache, 2017]) North Pyrenean Zone represents the European margin of the Apto-Cenomanian hyperextended rit system. Early Cretaceous extension was focussed mainly further south, where thinning of the crust exhumed mantle between the Iberian and European plates. he Metamorphic Internal Zone is thought to have overlain the zone of exhumed mantle [Lagabrielle et al., 2010;Mouthereau et al., 2014;[START_REF] Tugend | Characterizing and identifying structural domains at rited continental margins: application to the Bay of Biscay margins and its Western Pyrenean fossil remnants[END_REF][START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF]. he Roquefeuil and Benaix Faults (Figures 2.5b, 2.6b) represent the proximal northern margin of the rit system and accommodated ~1.5 km of N-S extension. he Roquefeuil Fault was active from Berriasian to Barremian. In the Aptian, the rit deepened and its margin migrated north to the Benaix Fault whose footwall was eroded to basement. A N-S lateral ramp delimited the Saint-Barthélémy high to the west, although the Mirande Group may still [1993] and northern Pyrenees (S3 section; Figures 2.4 and 2.5). Each line represents either the horizontal position of the leading edge of a principal fault or the position of the hinge zone at the surface of a major anticline. Displacement and fault length are assumed to increase linearly for each fault. For emergent faults, we assumed the fault reached maximum length during deposition of youngest cut strata. Fault timings for the southern Pyrenees fromfault-strata relationships [START_REF] Vergés | Corte compensado del Pirineo oriental: Geometría de las cuencas de antepaís y edades de emplazamiento de los mantos de corrimiento[END_REF]Vergés et al., 1995Vergés et al., , 1998Vergés et al., , 2002;;Ford et al., 1997;Ramos et al., 2002;[START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF][START_REF] Carrigan | Fault-related fold kinematics recorded by terrestrial growth strata, Sant Llorenç de Morunys, Pyrenees Mountains, NE Spain[END_REF]. (b, d) Horizontal position of the deformation front, the main depocentre(s) and foreland pinchout of the basin as a function of time in the southern and northern Pyrenees, respectively. he deformation front tracks the location of the most frontal structure through time. Depocentres were measured from balanced cross sections. Pinchout locations were determined from borehole and seismic data. Shortening rates through time are shown by grey bars, derived from the fault activity in (a, c). Shortening rates for the south include internal shortening in the Axial Zone (AZ). Dashed bars have a high uncertainty for the early convergence.

have covered it. Deep marine strata (Black Flysch Group) were deposited from the Aptian to Cenomanian, thinning westward across the lateral ramp. he post-rit basin progressively deepened as its margin migrated northward, onlapping unrited basement, from Middle Cenomanian to approximately Middle Turonian times, followed by progressive olap until end Santonian (Figure 2.8d). Following [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF] and [START_REF] Vacherat | Rit-to-collision transition recorded by tectonothermal evolution of the northern Pyrenees[END_REF], we propose that the post-rit Grey Flysch extended south to cover the whole rit zone.

2.4.4.2 First Orogenic Phase: 84-66 Ma, Early Inversion During early convergence, the Metamorphic Internal Zone was emplaced northward along the 3M Fault. Gentle inversion afected inherited normal faults of the European margin and began folding Mesozoic basin successions (Nalzen, Fougax). Estimated shortening rates were 0.3-0.6 mm/y, the latter value including a minimum of shortening related to emplacement of the Metamorphic Internal Zone. Early inversion of the margin is not recorded by low temperature thermochronological data from the adjacent Arize Massif [Fitzgerald et al., 1999;[START_REF] Vacherat | Rit-to-collision transition recorded by tectonothermal evolution of the northern Pyrenees[END_REF] suggesting that early basement exhumation was limited. A deep marine lexural basin was supplied with sediment mainly from the east [Bilotte, 1985]. Tectonic subsidence was relatively rapid though still only 0.09 mm/y close to the thrust front (Table 2.1). We relate this early subsidence to emplacement of the Metamorphic Internal Zone and early inversion and crustal thickening along the European margin. Campanian strata preserved in the Nalzen block indicate that the foreland basin continued south of the thrust front implying that the early orogenic ediice had low topographic relief and remained largely submarine. Shallower water facies in the Nalzen block suggest a syn-sedimentary uplit along the North Pyrenean Frontal hrust within the subsiding basin. he foreland basin records a deepening and then shallowing upward trend, ending in luvial conditions across the entire foreland. Aude Valley Group strata contain clasts of Upper Cretaceous limestones that were either sourced from the east, the south, or from local uplits [START_REF] Bilotte | Notice explicative de la feuille Lavelanet à 1/50 000 (Feuille №1076)[END_REF].

2.4.4.3 Second Phase: 66-59 Ma, Northern Quiescence During the early Paleocene, no fault activity was recorded. Tectonic subsidence slowed to near zero in the foreland with deposition of only a thin unit of ine-grained continental deposits. his quiescence is also documented further west [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF], suggesting it was a widespread phase in the northern Pyrenees.

2.4.4.4 hird Phase: 59-34 Ma, Slow Northern Activity Tectonic activity returned in the hanetian with distributed thick-skinned shortening across the North Pyrenean foreland. he shortening rate was lower than in the Late Cretaceous, at 0.3 mm/y on average (Table 2.1). A slight acceleration in foreland basin subsidence was associated with a minor marine incursion across the proximal foreland (Rieubach Group; Figure 2.3b). Continental facies were then deposited across the whole foreland. More widespread deformation returned in the early Ypresian (~56 Ma), with activity on the Tréziers hrust and Lavelanet Anticline. A basin-wide marine incursion can be related to a eustatic sea level rise (Coustouge Group) [Christophoul et al., 2003]. Widespread continental sedimentation (Carcassonne Group) was re-established at the end of the Ypresian and continued until at least Rupelian times (Figure 2.3b). Post-hanetian activity of the North Pyrenean Frontal hrust until at least lower Ypresian is recorded north of the Nalzen block. Uplit and erosion of the Metamorphic Internal Zone, Saint-Barthélémy Massif, and/or the Axial Zone is recorded by pebbles of metamorphosed limestone and crystalline basement in Carcassonne Group conglomerates [START_REF] Crochet | Molasses syntectoniques du versant nord des Pyrénées: la série de Palassou[END_REF]. Subsidence was slower during this phase at ~0.03 mm/y near the basin depocentre. he deformation front migrated northward, progressively activating basement-involved thrusts in the foreland (Tréziers and Orsans thrusts). he basin depocentre followed, and usually lay behind the deformation front (Figure 2.8d). he youngest evidence for deformation in the North Pyrenean foreland is slight folding of the upper Carcassonne Group above the Tréziers hrust, dated as approximately Priabonian (Figure 2.4d). [1993]. Boreholes and stratigraphic sections used for subsidence calculations are indicated as semitransparent white boxes with blue letters. M, Montserrat; C, Castellfollit; S, Santpedor; P-r, Puig-reig; J, Jabalí; G, Gombrèn (split in two); Pf, Pedraforca. Formation names are indicated by numbers in circles as in Figure 2.3a.

that the Tréziers and Orsans thrusts were abandoned at the end of the preceding phase, as they are sealed by the upper Carcassonne Group.

South Pyrenean Foreland

We performed a similar analysis of the South Pyrenean foreland along transect J3 (Figure 2.9) [Vergés, 1993]. his 120 km long balanced cross section stretches from the Axial Zone in the north to the Catalan Coastal Ranges in the south, traversing the eastern South Pyrenean Zone and Ebro Foreland Basin (Figure 2.1). he section was constrained by borehole data, surface geology, and seismic lines. A detailed description of this area can be found in Vergés et al. [1995Vergés et al. [ , 1998Vergés et al. [ , 2002] ] and only a summary is given here. Total shortening in the South Pyrenean foreland is 69.2 km, not accounting for internal deformation in the Axial Zone (~23 km) [Vergés et al., 1995] and Catalan Coastal Ranges (Figure 2.9).

Stratigraphy of the South Pyrenean Foreland

Upper Triassic Keuper evaporites serve as a basal detachment for the Upper and Lower Pedraforca thrust sheets in the South Pyrenean Zone but is absent further south below the Ebro Basin, except close to the Catalan Coastal Ranges (Figure 2.9). A thin Liassic limestone and dolomite succession is only found in the Pedraforca thrust sheets. Cretaceous strata are mostly found in the Pedraforca thrust sheets and thin southward over a distance of 10 km (Figure 2.9b) [Vergés, 1993]. he Upper Pedraforca thrust sheet, which restores to the most northerly position, mainly comprises a ~1500 m thick inverted Lower Cretaceous extensional basin. he Lower Pedraforca thrust sheet contains ~1700 m of post-rit and syn-orogenic Upper Cretaceous conglomerates, marls and limestones, capped by ~700 m of Maastrichtian to Paleocene red beds (Garumnian facies) of the Tremp Formation (Figure 2.3a) [START_REF] Vergés | Corte compensado del Pirineo oriental: Geometría de las cuencas de antepaís y edades de emplazamiento de los mantos de corrimiento[END_REF][START_REF] Vergés | Mapa Geológico de España[END_REF].

A regional unconformity marks the base of Eocene marine sedimentation in the South Pyrenean foreland [START_REF] Pujalte | Late Cretaceous-Middle Eocene sequence stratigraphy and biostratigraphy of the SW and W Pyrenees (Pamplona and Basque basins, Spain)[END_REF][START_REF] Serra-Kiel | Cronoestratigrafía de los sedimentos marinos del Terciario inferior de la Cuenca de Graus-Trem (Zona Central Surpirenaica)[END_REF]Vergés et al., 1998]. hese marine sediments overlie the Tremp Formation in the north and onlap basement further south. hey record four transgressive-regressive cycles from Ypresian to Bartonian [Puigdefàbregas and Souquet, 1986;[START_REF] Martinez | Stratigraphic framework of the thrust geometry and structural inversion in the southeastern Pyrenees : La Garrotxa area[END_REF]Vergés et al., 1998]. Within each cycle, terrigenous clastic deposits were deposited in front of the advancing South Pyrenean hrust Front (Corones, Bellmunt and Milany Formations, Figure 2.3a). hese depositional units grade southward into moderate to deep-water marine carbonates and calcareous mudstones (Sagnari, Armáncies, Campdevánol and Banyoles Formations; Figure 2.3a) that, in turn, grade into shallow marine platform carbonates (Cadí, Penya and Tavertet Formations; Figure 2.3a) towards the southernmost border of the foreland basin. During the fourth cycle, the basin was characterised by prodelta ofshore marls (Igualada Formation; Figure 2.3a) bounded by northward prograding fan-delta deposits sourced from the upliting Catalan Coastal Ranges in the south (Montserrat Formation; Figure 2.3a) and southward prograding Milany Formation in the north. At the end of cycles two and four, thick evaporitic units are deposited in the central marine basin, the marine Beuda and transitional Cardona Formations, respectively (Figure 2.3a). In the footwall of the Vallfogona thrust, the Beuda evaporites reach 2000 m thickness, including a 200 m thick package of salt [START_REF] Martinez | Stratigraphic framework of the thrust geometry and structural inversion in the southeastern Pyrenees : La Garrotxa area[END_REF]. he Ebro foreland basin became fully continental during the deposition of the uppermost part of the Cardona evaporitic Formation at ~36 Ma (middle Priabonian) [START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF]. Widespread non-marine deposition was then established until the latest Oligocene comprising the alluvial and luvial Solsona Formation and equivalent lacustrine deposits in the centre of the foreland basin (Figure 2.3a) [START_REF] Meigs | Ten-million-year history of a thrust sheet[END_REF][START_REF] García-Castellanos | Interplay between tectonics, climate, and luvial transport during the Cenozoic evolution of the Ebro Basin (NE Iberia)[END_REF][START_REF] Valero | 20 Myr of eccentricity paced lacustrine cycles in the Cenozoic Ebro Basin[END_REF][START_REF] Garcia-Castellanos | Quantifying the post-tectonic topographic evolution of closed basins: he Ebro basin (northeast Iberia)[END_REF].

2.5.2 Structure of the South Pyrenean Foreland he South Pyrenean foreland fold-and-thrust belt, here called the South Pyrenean Zone, lies between the antiformal Axial Zone to the north and the mildly deformed Ebro foreland basin to the south (Figure 2.1b) [Vergés et al., 2002]. In the Eastern Pyrenees, three stacked thrust sheets compose the South Pyrenean Zone, namely from top (oldest) to bottom (youngest), the Upper Pedraforca, the Lower Pedraforca and the Cadí thrust sheets (Figure 2.9). he Upper Pedraforca and Lower Pedraforca thrust sheets are mostly constituted by Jurassic to Palaeocene units detached above the Keuper evaporites. he Vallfogona thrust transports the Cadí thrust sheet with the Upper and Lower Pedraforca thrust sheets in its hanging wall. Following Vergés [1993], and Vergés et al. [1995Vergés et al. [ , 2002]], the Pedraforca thrust sheets root below the Nogueres basement thrust sheet of the Axial Zone whereas the Vallfogona thrust links back to the Orri basement thrust (Figure 2.1b). [START_REF] Laumonier | he Alpine southeastern Pyrenees (France, Spain); a synthesis[END_REF] alternatively suggests that the Pedraforca thrust sheets may root north of the North Pyrenean Fault, forming part of a 'supra-axial unit' . his model depends on the interpretation of the eastern Axial Zone as a single block. he kinematics of the supra-axial unit proposed by [START_REF] Laumonier | he Alpine southeastern Pyrenees (France, Spain); a synthesis[END_REF] implicitly require coeval closure of the exhumed mantle domain and emplacement of the Lower Pedraforca thrust sheet, which is inconsistent with our data (see sections 5.4.2 and 6.2).

he Puig-reig anticline within the Ebro foreland basin is the surface expression of a thrust ramp that steps up from the Beuda evaporite horizon to the Cardona salt horizon (Figure 2.9). his thrust links back into the Rialp basal thrust in the Axial Zone. Further south, the Cardona and Súria anticlines are detached and thrust along the Cardona evaporite horizon [e.g., Vergés et al., 1992]. he southern margin of the Ebro basin was the result of the thick-skinned inversion of the Tethyan Mesozoic basin of the Catalan Coastal Ranges from Lutetian to Chattian [START_REF] References Anadón | Las facies lacustres del olioceno de campins (Vallès Oriental, provincia de Barcelona)[END_REF][START_REF] López-Blanco | Tectonic and climatic controls on the development of foreland fan deltas: Montserrat and Sant Llorenç del Munt systems (Middle Eocene, Ebro Basin, NE Spain)[END_REF][START_REF] Gómez-Paccard | Tectonic and climatic controls on the sequential arrangement of an alluvial fan/fan-delta complex (Montserrat, Eocene, Ebro Basin, NE Spain)[END_REF]. he southern and eastern borders of the Ebro foreland basin were regionally uplited during the Cenozoic extensional phase that opened the Western Mediterranean, thus tilting the entire southern foreland [e.g., [START_REF] Lewis | High mountains in a zone of extended crust: Insights into the Neogene-Quaternary topographic development of northeastern Iberia[END_REF]. A .4). However, the Gombrèn stratigraphic section, now located in the Cadí thrust sheet, records slightly faster tectonic subsidence in the Paleocene (0.05 mm/y; Figure 2.7b; Table A .4).

Subsidence accelerates in the Eocene, forming the classic convex-up subsidence pattern expected for proforeland basins [Sinclair and Naylor, 2012]. Ypresian tectonic subsidence was high in the north of the Ebro foreland basin (Gombrèn section; 0.53 mm/y around 50 Ma; Table A.4), while rates were much lower to the south (0.03 to 0.07 mm/y; Table A .4). he onset of subsidence acceleration (>0.1 mm/y) migrated south at a rate of about 10 mm/y between 50.7 Ma and 40 Ma [Vergés et al., 1998]. his outward migration is also typical of pro-foreland basins [Sinclair and Naylor, 2012]. Increases in tectonic subsidence can be correlated with tectonic activity in the Axial Zone and South Pyrenean Zone (Figure 2.8a). For example, at 43 Ma tectonic subsidence of the Jabalí and Puig-reig boreholes increased to 0.26 and 0.14 mm/y respectively in response to loading by the Orri basement thrust sheet and emplacement of the Cadí thrust sheet [Vergés et al., 1998]. Tectonic subsidence in the south (Montserrat section) sees a brief acceleration to 0.21 mm/y at 41.5 Ma that can be correlated to activity of the Catalan Coastal Ranges [Vergés et al., 1998]. Ater ~36 Ma, tectonic subsidence slowed down throughout the southern foreland basin (Figure 2 2.5.4 Evolution of the South Pyrenean Foreland he summary of South Pyrenean foreland evolution given here is based on a compilation of subsidence and stratigraphic data, and a restored cross section [Vergés, 1993]. he boundaries of each evolutionary phase were chosen to relect signiicant changes in the whole orogenic system, and therefore do not necessarily correspond perfectly to distinct changes in the South Pyrenean foreland itself.

Chapter 2. Case study of the Eastern Pyrenees 51 2.5.4.1 First and Second Phase: 84 -59 Ma, Rit Inversion he timing of the onset of shortening in the southern Pyrenees cannot be distinguished in the preserved subsidence signal and is not clear from the rest of our data. Growth strata related to the Boixols thrust, carrying the western equivalent of the Upper Pedraforca thrust sheet, place the onset of shortening around ~72 Ma [Puigdefàbregas and Souquet, 1986;[START_REF] Bond | Inversion of a Lower Cretaceous extensional basin, south central Pyrenees, Spain, in Basin Inversion[END_REF]. However, many authors place the onset of shortening in the southern Pyrenees at Late Santonian, or ~84 Ma [START_REF] Ardèvol | Depositional sequence response to foreland deformation in the upper Cretaceous of the Southern Pyrenees, Spain[END_REF]McClay et al., 2004;Saura et al., 2016]. In the western Pyrenees, [START_REF] Teixell | he Anso transect of the southern Pyrenees: basement and cover thrust geometries[END_REF] also places the onset of convergence at Late Santonian. A Late Cretaceous lexural basin is preserved in the Lower Pedraforca thrust sheet (Figures 2.7a,b,2.9b). We interpret this subsidence as a response to early inversion of the distal Iberian margin, during closure of the exhumed mantle domain. he Tremp Formation of latest Maastrichtian to Upper Paleocene age seals the Upper Pedraforca thrust sheet [START_REF] Vergés | Corte compensado del Pirineo oriental: Geometría de las cuencas de antepaís y edades de emplazamiento de los mantos de corrimiento[END_REF]Vergés et al., 1995], however, tectonic subsidence appears to have continued very slowly throughout the Paleocene (Figure 2.7b; 2.3a). Meanwhile, further south, tectonic subsidence rates and sediment supply were much lower (around 0.04 mm/y), allowing the growth of carbonate platforms. Emplacement of the Lower Pedraforca thrust sheet occurred during early Eocene until ~47 Ma, as shown by syn-orogenic strata [START_REF] Solé Sugrañes | Nota sobre la edad y posición tectónica de los conglomerados eocenos de Queralt (Prepirineo oriental, Prov. de Barcelona)[END_REF][START_REF] Vergés | Corte compensado del Pirineo oriental: Geometría de las cuencas de antepaís y edades de emplazamiento de los mantos de corrimiento[END_REF]Burbank et al., 1992b].

Shortening slowed down signiicantly ater the end of the Ypresian to 1.0 mm/y (Figure 2.8b). A new basement thrust sheet (Orri) became active, as constrained by the Vallfogona thrust, marking the start of true thick-skinned frontal accretion (Figure 2.8a). Crustal thickening and exhumation in the Axial Zone started at ~50 Ma, recorded by low-temperature thermochronology data [Fitzgerald et al., 1999;Sinclair et al., 2005;Metcalf et al., 2009;Rahl et al., 2011;Whitchurch et al., 2011]. Axial Zone deformation accommodated an average of 1.1 mm/y of extra shortening until the end of convergence (~23 km since 50 Ma; Figure 2.8b) [Vergés et al., 1995]. his was accompanied by an increase in tectonic subsidence rates throughout the foreland basin to around 0.10 mm/y (Figure 2.7b). he Gombrèn subsidence record in the proximal foreland shows a much lower rate starting in the late Ypresian owing to deformation ahead of a propagating Vallfogona thrust, which reached the surface in the mid-Lutetian [Vergés et al., 1998;Ramos et al., 2002]. . Widespread growth strata within the Berga conglomerates (Figs, 2.3, 2.9) constrain the latest activity of the Vallfogona thrust to at least ~31 Ma (Figure 2.8a) [Ford et al., 1997;Vergés et al., 1998;[START_REF] Carrigan | Fault-related fold kinematics recorded by terrestrial growth strata, Sant Llorenç de Morunys, Pyrenees Mountains, NE Spain[END_REF]. Most of the foreland basin remained marine to brackish, but sedimentation along the thrust front was continental from Lutetian onwards (Figure 2.3a). he Cardona salt horizon was deposited around 36 Ma as a result of the isolation from the Atlantic Ocean ater uplit of the Western Pyrenees [Burbank et al., 1992b;[START_REF] Vergés | Eocene-Oligocene hrusting and Basin Coniguration in the Eastern and Central Pyrenees (Spain), in Tertiary Basins of Spain. he Stratigraphic Record of Crustal Kinematics[END_REF][START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF]. Growth strata on the lanks of the Puig-reig anticline date the onset of folding to around the same time [Vergés et al., 2002]. his was related to the Rialp basement thrust and increased the shortening rate to 1.6 mm/y (Figure 2.8b).

2.5.4.3 Fourth Phase: 34 -28 Ma, Final Deformation During this inal phase, the shortening rate in the foreland decreased to 0.6 mm/y (Figure 2.8b). Most shortening in the foreland was accommodated by detachment folding above the Cardona evaporites and the inal activity of the Vallfogona thrust [Vergés et al., 1992;[START_REF] Carrigan | Fault-related fold kinematics recorded by terrestrial growth strata, Sant Llorenç de Morunys, Pyrenees Mountains, NE Spain[END_REF]. he deformation front thus rapidly migrated to the south along the Cardona evaporite horizon, while the depocentre remained approximately stationary (Figure 2.8b). We have no subsidence records for this phase, but sedimentation in the foreland remained continental (Solsona Formation and equivalents; Figure 2.3a). In this part of the southern foreland, the youngest preserved sediments (Rupelian) are deformed, suggesting deformation continued until at least ~28 Ma. Shortening in the Central Pyrenees is estimated to have ended around 30 Ma based on Apatite ission track data from the Axial Zone [Fitzgerald et al., 1999], or ~24.7 Ma based on magnetostratigraphic dating of the South Central Unit and its footwall [START_REF] Meigs | Ten-million-year history of a thrust sheet[END_REF].

Discussion

2.6.1 Crustal-scale Model of Orogenic Evolution Geodynamic models for the Pyrenees have evolved over the last 30 years, as have estimates of total N-S shortening. Most crustal-scale restorations have focussed on the Central Pyrenees along the Ariège ECORS deep seismic line. Using upper crustal imbricates in the Axial Zone and restoring to normal crustal thickness, Muñoz [1992] proposes 147 km of N-S shortening while Beaumont et al. [2000] prefer 165 km over several other possible geometries. More recent reconstructions restore to the pre-orogenic rit system [e.g., Mouthereau et al., 2014;Teixell et al., 2016], most notably integrating the hypothesis that the Iberian and European plates were separated by a zone of exhumed mantle due to Aptian-Cenomanian hyper-extension [e.g., Jammes et al., 2009;Lagabrielle et al., 2010]. However, the original width of this zone and hence the convergence necessary to subsequently close it are unknown. Proposed widths vary from 15 km in the Western Pyrenees [Teixell et al., 2016] to 50 km of exhumed mantle in the Central Pyrenees [Mouthereau et al., 2014] or several hundred km of oceanic crust [Vissers and Meijer, 2012b]. However, deep seismic tomography imaging below the Pyrenees rules out oceanic subduction prior to collision, as no subducted oceanic slab can be identiied in the mantle [START_REF] Chevrot | High-resolution imaging of the Pyrenees and Massif Central from the data of the PYROPE and IBERARRAY portable array deployments[END_REF]. Using full crustal imbrication, as proposed by Roure et al. [1989], Mouthereau et al. [2014] obtain 92 km of total shortening along the ECORS line, excluding closure of the exhumed mantle domain. his markedly lower value is principally related to a revised shortening estimate for the southern Pyrenees compared to Beaumont et al. [2000]. In the western Pyrenees along ECORS-Arzacq, [START_REF] Teixell | he Anso transect of the southern Pyrenees: basement and cover thrust geometries[END_REF]Teixell [ , 1998] ] estimates a total shortening of about 80 km while more recently, Teixell et al. [2016] propose 99 km along a section 15 km to the east of ECORS-Arzacq, excluding closure of the exhumed mantle domain.

In the Eastern Pyrenees, Vergés et al. [1995] integrate surface geology with the deep structure of the ECORS line projected onto the same section line as this paper. By using the work of Baby et al. [1988] for the northern Pyrenees, assuming imbrication of Iberian upper crust only and subduction of Iberian lower crust, restoring to normal crustal thickness, and integrating estimates of erosion and sedimentation during main orogenic phases, they ind a minimum total shortening of 125 km. [START_REF] Vergés | Mesozoic evolution and Cainozoic inversion of the Pyrenean rit, in Peri-Tethys Memoir 6: Peri-Tethyan Rit/Wrench Basins and Passive Margins[END_REF] revise the restoration (but not the estimated shortening) by proposing two possible geometries for the rited crust. We further update this reconstruction by integrating the detailed deformation history presented above (Figure 2.8) with subsidence (Figure 2.7), thermochronology [e.g., Metcalf et al., 2009], and deep geophysical data [e.g., [START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF]. his allowed us to build a new and more detailed evolutionary model for the Eastern Pyrenees in four phases, chosen to relect signiicant changes in the dynamics of the orogenic system (Figure 2.10). he main diferences compared to Vergés et al. [1995Vergés et al. [ , 2002] ] and [START_REF] Vergés | Mesozoic evolution and Cainozoic inversion of the Pyrenean rit, in Peri-Tethys Memoir 6: Peri-Tethyan Rit/Wrench Basins and Passive Margins[END_REF] are summarised in the following paragraphs.

First, the deep structure of the present day section integrates major lithospheric boundaries identiied by recent P wave tomography and migration of converted P waves along the ECORS line [START_REF] Chevrot | High-resolution imaging of the Pyrenees and Massif Central from the data of the PYROPE and IBERARRAY portable array deployments[END_REF][START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF]. [START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF] identify the interface between the top of the Iberian crust and the European mantle, dipping 20° N. heir Iberian Moho (at 32 km depth) corresponds well with ECORS data, but when projected to our section and aligned along the North Pyrenean Fault, the Moho inlexion to dip north at 20° is positioned below the Puig-reig anticline (Figure 2.10a). his Moho geometry leads to an overthickened Iberian crust (>50 km) below the Axial Zone, which can be neither isostatically justiied nor volumetrically balanced. We therefore continue the Iberian lower crust northward to an inlexion point below the southern Axial Zone (as imaged on the ECORS proile), where it dips 20°N into the mantle. his creates a slight mismatch between the projected Moho and our model, a result of the slightly narrower Axial Zone along our section. he apparent southward thinning of the European crust observed by [START_REF] Chevrot | he Pyrenean architecture as revealed by teleseismic P-to-S converted waves recorded along two dense transects[END_REF] is integrated into our model. Second, in line with recent indings, the composite fully restored section (end Santonian; Figure 2.10e) shows the rited margins separated by a zone of exhumed mantle with a hypothetical width of 50 km. he basins are illed with a syn-rit to post-rit succession that is ~4 km thick above mantle. Following Lagabrielle et al. [2010], we position the metamorphosed syn-rit succession (future Metamorphic Internal Zone) directly above mantle. he geometry of the pre-orogenic extensional fault system is still a matter of much debate as discussed above. Restored geometries on the Iberian plate follow those of [START_REF] Vergés | Mesozoic evolution and Cainozoic inversion of the Pyrenean rit, in Peri-Tethys Memoir 6: Peri-Tethyan Rit/Wrench Basins and Passive Margins[END_REF] and Vergés et al. [2002] while the European margin is based on new work in this paper. he inferred geometries of the Iberian and European rited margins are constrained by the observed remnants of extensional Mesozoic basins, preserve the area needed for a precise crustal balancing through time, are isostatically stable, and are guided by new seismic interpretations of the Atlantic Iberian margins [e.g., [START_REF] Sutra | Quantiication and restoration of extensional deformation along the Western Iberia and Newfoundland rited margins[END_REF]. hird, the lower estimate of shortening for the European foreland (~20 km) leads to a new minimum overall shortening of 111 km. All estimations of shortening given here exclude closure of the exhumed mantle domain. We cannot constrain the original volume of the Metamorphic Internal Zone nor its tectonic and erosional history. he largest uncertainties associated with the amount and distribution of shortening are therefore for the irst phase.

2.6.1.1 First Phase: 84-66 Ma, Rit Inversion During the Late Cretaceous (Figure 2.10e to d) slow and distributed shortening (~18 km) was accommodated by inversion of normal faults across the whole rit system while the mantle domain closed. An overall shortening rate of ~1 mm/y is estimated with deformation distributed roughly equally across the two margins (Figure 2.8b,d;Table 2.1). As this estimate excludes closure of the mantle zone, true convergence rates would be considerably higher. For example, assuming a 50 km wide mantle zone implies a convergence rate of 3.8 mm/y. he Metamorphic Internal Zone, of unknown dimensions, was thrust northward and probably also southward if we assume that it covered the whole zone of exhumed mantle (Figure 2.10d) [Mouthereau et al., 2014;Teixell et al., 2016]. Marine lexural foreland basins developed on both plates and may have formed a continuous basin across the submarine, low relief orogen. Depocentres became overilled and continental in the Maastrichtian (Figure 2.3). Average tectonic subsidence rates are 0.09 mm/y and 0.06 mm/y for the north and south, respectively. Low temperature thermochronological data to the east suggest the onset of uplit of the North Pyrenean Massifs during this early phase [START_REF] Ternois | Zircon ( U-h )/ He evidence for pre-Eocene orogenic exhumation of eastern North Pyrenean massifs , France[END_REF].

2.6.1.2 Second Phase: 66-59 Ma, Slowdown and Northern Quiescence he total shortening rate slowed to ~0.4 mm/y during the Danian and Selandian (Figure 2.10c). Subduction polarity was established during this phase, the asymmetry of which resulted in a temporary stop in foreland subsidence and deformation of the upper plate (Figures 2.7d, 2.8d; Tables 2.1, A.3). However, very slow tectonic subsidence (0.05 mm/y) and shortening (0.4 mm/y) continued on the lower plate, accommodating some 2 km of shortening. Continental sedimentation prevailed in both foreland basins.

2.6.1.3 hird Phase: 59-34 Ma, Main Collision and Southern Dominance he principal phase of collision (approximately 78 km of shortening) took place between ~59 and ~34 Ma with shortening and tectonic subsidence rates revealing a strong asymmetry, being much higher in the south (2.7 mm/y and 0.12 mm/y on average, respectively) than in the north (0.3 mm/y and 0.03 mm/y on average, respectively; Figures 2.7, 2.8; Table 2.1). he peak in overall shortening rate (~4.5 mm/y) began at ~56 Ma and was followed by a gradual deceleration to 2.3 mm/y until ~36 Ma, when accretion of a new basement thrust sheet (Rialp) increased the overall rate to 2.9 mm/y (Figures 2.8a, A1c). he South Pyrenean wedge accommodated ~70 km of shortening by frontal accretion and thickening of the Axial Zone. In the north, shortening was accommodated along several thrusts simultaneously, with the North Pyrenean Frontal hrust accounting for the majority of shortening (Figure 2.8c). Ater the early Ypresian marine incursion, continental conditions were quickly re-established in the northern foreland while the south sustained a carbonate platform and deep marine basin until ~36 Ma (Figure 2.3) [Vergés et al., 1998;[START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF]. Low-temperature thermochronology records uplit and exhumation of the Axial Zone from ~50 Ma [e.g., Metcalf et al., 2009], which is corroborated by alluvial sedimentation along the thrust fronts of both wedges.

2.6.1.4 Fourth Phase: 34-28 Ma, Final Shortening In the inal stage, ater ~34 Ma, orogenic processes gradually died out. he eastern northern foreland appears to have been abandoned at around 34 Ma but this is unclear due to later uplit and erosion (Fig 2 .8c; Table 2.1). In contrast, during the Oligocene ~13 km of shortening was accommodated in the south by detachment anticlines above the Cardona salt horizon (~7 km) and by internal deformation and uplit of the Axial Zone (~6 km; Figures 2.8a, b, 2.10a). While the exact timing of the end of convergence is not well known, we estimate an average shortening rate of ~2.2 mm/y (Table 2.1) up to at least ~28 Ma. Postorogenic exhumation of the southern Axial Zone continued into the Miocene [Rushlow et al., 2013].

2.6.2 Improved detail in orogen deformation history Crustal deformation histories previously proposed for the Pyrenees are similar in broad terms to our results, with inversion of the rit followed by pro-wedge propagation and basement stacking in the Axial Zone [e.g., Sinclair et al., 2005]. However, along our section in the Eastern Pyrenees, the distribution of displacement on faults through time (Figure 2.8) has revealed a more detailed deformation history.

While recent models include the closure of an exhumed mantle domain, how this event relates to inversion of the rited margins remains diicult to constrain. Mouthereau et al. [2014] show the mantle domain closing irst, followed by thick-skinned inversion of rit structures ater ~75-70 Ma. We propose instead that closure of the mantle domain took place progressively and simultaneously with thick-skinned inversion of the rited margins, based on timing of the North Pyrenean Frontal hrust as constrained by stratigraphy. his is corroborated by the foreland tectonic subsidence signal that we link to emplacement of the Metamorphic Internal Zone and loading of the margins.

None of the published crustal-scale restorations recognise the Paleocene slowdown described in this paper, either because it is below the temporal resolution of the sequential restoration steps used [Muñoz, 1992;Beaumont et al., 2000;Teixell et al., 2016], or because it cannot be detected by time-temperature models from apatite ission track data used to constrain vertical motions [Mouthereau et al., 2014]. Furthermore, previous attempts at estimating convergence rates through time have given conlicting results. In the Central Pyrenees, Beaumont et al. [2000] show an overall acceleration, in particular at 36 Ma, related to late stacking of crustal nappes in the Axial Zone. In contrast, Mouthereau et al. [2014] propose a high initial rate of convergence (~3 mm/y) in the Late Cretaceous and Early Paleocene, followed by gradual deceleration until the Miocene. For the Western Pyrenees, Teixell et al. [2016] estimate very low rates (<1 mm/y) during the Late Cretaceous and Paleocene, followed by the principal peak of convergence in the Eocene (3-2.5 mm/y). Along our section in the Eastern Pyrenees, we show three peaks of relatively rapid shortening (83, 56, and 36 Ma), each followed by a period of slower and decelerating shortening (Figures 2.8b, d, A1; Table 2.1). We discuss the main diferences in the following paragraphs. he highest shortening rates in the Eastern Pyrenees, according to our reconstruction, occurred during the early Eocene, which can be considered as the onset of main collision. his coincides with relatively rapid tectonic subsidence of the Aquitaine Basin starting at ~55 Ma (Figure 2.7) [Desegaulx and Brunet, 1990;[START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF], and the onset of crustal stacking in the Axial Zone around 50 Ma [e.g., Metcalf et al., 2009]. Teixell et al. [2016] show that the main pulse of shortening in the Western Pyrenees also occurred during the Eocene. In the Central Pyrenees however, neither Beaumont et al. [2000] nor Mouthereau et al. [2014] place the highest shortening rates during the Eocene. he initial high shortening rate of Mouthereau et al. [2014] is based on Late Cretaceous to Paleocene cooling detected by detrital lowtemperature thermochronology, from foreland basin sediments that the authors assume were derived from the Pyrenean orogen. However, paleocurrent and facies analyses [Bilotte, 1985] show that these sediments were sourced from an area to the east of the Pyrenees. herefore, these data do not constrain exhumation rate in the Pyrenees along the ECORS line. Furthermore, the gradually decelerating convergence proposed by Mouthereau et al. [2014] cannot be reconciled with the increased Eocene tectonic subsidence and shortening rates for the forelands. Beaumont et al. [2000] place the highest shortening rates ater 36 Ma, in accordance with rapid exhumation in the Axial Zone, as shown by numerous thermochronological studies [START_REF] Morris | Exhumation of the Pyrenean Orogen: implications for sediment discharge[END_REF]Fitzgerald et al., 1999;Jolivet et al., 2007;[START_REF] Maurel | he Meso-Cenozoic thermo-tectonic evolution of the Eastern Pyrenees: An 40Ar/39Ar ission track and (U-h)/He thermochronological study of the Canigou and Mont-Louis massifs[END_REF]Metcalf et al., 2009;Fillon and van der Beek, 2012;Rushlow et al., 2013]. However, thrusting along the eastern border of the South Central Unit (~25 km east of ECORS; Figure 2.1a) appears to have slowed down ater 36 Ma [Burbank et al., 1992a]. Furthermore, foreland tectonic subsidence slows down around 36 Ma in both foreland basins. his is well constrained in the south by magnetostratigraphic dating (Figure 2.7b) [Vergés et al., 1998]. However, the inal deceleration is less well constrained in the north as rates were already very low and the youngest strata are poorly dated and preserved (Figure 2.7d) [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF]. Although the peak in exhumation rate of the Axial Zone may have occurred around 36 Ma, there is evidence for exhumation of the Axial Zone since ~50 Ma [Fitzgerald et al., 1999;Sinclair et al., 2005;Metcalf et al., 2009;Rahl et al., 2011;Whitchurch et al., 2011]. Because exhumation rate cannot be directly related to shortening rate, we assume a constant shortening rate for internal shortening of the Axial Zone since ~50 Ma, thus yielding a lower shortening rate ater 36 Ma than shown by Beaumont et al. [2000] and providing a better it with foreland evolution. 2.6.3 Linked Pro-and Retro-wedge Dynamics As described in section 2.6.1, Late Cretaceous shortening was distributed across both the Iberian and European rited margins, before the convergent system evolved into an asymmetrical, pro-wedge-dominant orogen (Figure 2.8b,d). Models of doubly vergent orogens with distinct lower and upper plates typically show a pro-wedge-dominant strain partitioning [Beaumont et al., 2000;Sinclair et al., 2005;Hoth et al., 2007;[START_REF] Cruz | Deformation, exhumation, and topography of experimental doubly-vergent orogenic wedges subjected to asymmetric erosion[END_REF]Duerto and McClay, 2009;[START_REF] Hardy | Deformation and fault activity in space and time in highresolution numerical models of doubly vergent thrust wedges[END_REF]Jammes and Huismans, 2012;Erdős et al., 2014Erdős et al., , 2015;;[START_REF] Jammes | Lateral variation in structural style of mountain building: controls of rheological and rit inheritance[END_REF]. However at the onset of Pyrenean convergence, no such distinction can be made. Instead, the template comprises two opposing rited margins separated by exhumed mantle [Jammes et al., 2009[START_REF] Jammes | Lateral variation in structural style of mountain building: controls of rheological and rit inheritance[END_REF]Lagabrielle et al., 2010;Mouthereau et al., 2014;[START_REF] Clerc | High-temperature metamorphism during extreme thinning of the continental crust: A reappraisal of the North Pyrenean passive paleomargin[END_REF][START_REF] Vacherat | Rit-to-collision transition recorded by tectonothermal evolution of the northern Pyrenees[END_REF]. Numerical models that include rited margins show early inversion that is roughly symmetrical, until full collision and onset of subduction of continental mantle lithosphere forces a change to an asymmetric geometry at depth [Jammes and Huismans, 2012;Erdős et al., 2014;[START_REF] Jammes | Lateral variation in structural style of mountain building: controls of rheological and rit inheritance[END_REF]. Whether this change to an asymmetric deep geometry in the Pyrenees occurred during closure of the exhumed mantle domain or at the beginning of main collision is unclear. Our model suggests that the transition to a prowedge-dominant strain partitioning and related abandonment of the retro-wedge around 66 Ma occurred shortly ater closure of the exhumed mantle domain. he change in strain partitioning may thus record the change from early rit inversion to main collision (i.e. continental subduction) in the Pyrenees, implying this major change does not necessarily rely on external forcing. In the Western Pyrenees, Teixell et al.

[2016] infer that closure of the mantle domain did not occur until the Early Eocene. his suggests that early convergence was less signiicant in the west than in the east, and/or that the mantle domain was wider in the west than in the east as proposed by Jammes et al. [2009].

he Paleocene retro-foreland quiescence must be explained on a regional scale, as it has also been recognised in the Central Pyrenees [START_REF] Ford | Retro-wedge foreland basin evolution along the ECORS line, eastern Pyrenees, France[END_REF][START_REF] Rougier | Stratigraphic and tectonic studies in the central Aquitaine Basin, northern Pyrenees: Constraints on the subsidence and deformation history of a retroforeland basin[END_REF]. One potential cause for this north Pyrenean quiescence is a pause in plate convergence. While some plate reconstructions based on magnetic sea loor anomalies show a complete stagnation of convergence in the Paleocene [Roest and Srivastava, 1991;Rosenbaum et al., 2002], others show continuous convergence [Vissers and Meijer, 2012a]. Other potential contributing factors for the north Pyrenean quiescence can be related to redistribution of shortening within the surrounding plates [e.g., Mouthereau et al., 2014], or a change in the deep structure of the orogen as it evolves to full collision, as discussed below.

Late Cretaceous subsidence in both the pro-and retro-forelands was mainly driven by loading related to emplacement of the Metamorphic Internal Zone and onset of inversion of the rited margins. hanetian to Oligocene pro-foreland behaviour can be directly related to crustal thickening and loading during main collision. We here propose that contemporaneous retro-foreland subsidence was driven by backthrusting of the thickened pro-wedge onto the European plate, similar to predictions from numerical and analogue models [e.g., Hoth et al., 2007;Erdős et al., 2014]. he deep structure imaged by the ECORS line indicates that the northern part of the Axial Zone overlies the European plate (Figure 2.1b) [Roure et al., 1989;Muñoz, 1992;Teixell, 1998;Teixell et al., 2016]. hermochronological data generally place the onset of exhumation in the Axial Zone at ~50 Ma [Fitzgerald et al., 1999;Sinclair et al., 2005;Metcalf et al., 2009;Rahl et al., 2011;Whitchurch et al., 2011], but in the eastern Pyrenees this may have started as early as ~55-60 Ma [START_REF] Maurel | he Meso-Cenozoic thermo-tectonic evolution of the Eastern Pyrenees: An 40Ar/39Ar ission track and (U-h)/He thermochronological study of the Canigou and Mont-Louis massifs[END_REF]. Tectonic subsidence of the retro-foreland resumed around 59 Ma, suggesting that the period between the onset of pro-dominant strain partitioning (~66 Ma) and the onset of Axial Zone backthrusting (~60 Ma at the earliest) may have contributed to the Paleocene quiescence observed in the retro-foreland.

Using subsidence analysis and detailed tectonic reconstructions to compare the evolution on both sides of a natural system has given insight into crustal-scale dynamics of doubly vergent orogens. Changes in the strain distribution over time may be intrinsic to the system and independent of external factors. So far, these changes have not been suggested by models [e.g. Sinclair et al., 2005;Hoth et al., 2007]. Previous foreland basin subsidence models simplify the orogen into a uniformly growing double wedge and therefore do not predict multiple periods of retro-foreland subsidence, nor a quiescent phase [Sinclair et al., 2005;Naylor and Sinclair, 2008;Sinclair and Naylor, 2012].

Conclusions

We present new stratigraphic and tectonic data for the north Pyrenean fold-and-thrust belt and Aquitaine Basin in the Eastern Pyrenees, which, when integrated with published data, provide new insight into retrowedge evolution. By integrating these new results with equivalent data on a section through the southern Pyrenees and other published data (thermochronology, seismic tomography data) we derive an updated full crustal history of the East Pyrenean double-wedge orogen.

Calculated minimum shortening for the East Pyrenean retro-foreland fold-and-thrust belt and Aquitaine Basin was ~20 km during the Pyrenean orogeny. Late Cretaceous shortening and tectonic subsidence rates in the northern Pyrenees (0.6 mm/y and 0.09 mm/y, respectively) were followed by a Paleocene quiescence and lower rates during the Eocene (0.3 mm/y and 0.03 mm/y). A minimum overall N-S shortening of ~111 km was estimated for the Eastern Pyrenees, not including closure of an inferred exhumed mantle domain of unknown width. he evolution of the double wedge orogen is divided into four phases: (1) Late Cretaceous shortening was characterised by closure of the exhumed mantle domain and inversion of the Iberian and European margins. An overall shortening rate of ~1 mm/y on average was distributed roughly equally between the two margins. Signiicant early foreland basins developed on both margins well before the creation of signiicant topography. (2) In the Paleocene slow shortening continued in the southern Pyrenees (~0.4 mm/y on average) while the north was temporarily inactive. ( 3) he highest overall shortening rate was reached in the Eocene (~3.1 mm/y on average) recording main collision. his was predominantly accommodated by shortening in the southern Pyrenees and Axial Zone. However, continued thickening of the Axial Zone then led to backthrusting onto the upper plate, reactivating the retro-wedge and its foreland basin. ( 4) Overall shortening slowed in the Oligocene to ~2.1 mm/y on average. he northern Pyrenees were probably inactive in this phase. According to our revised evolutionary model for the Eastern Pyrenees, the evolution from rit inversion to main collision caused a change from equal shortening in both rited margins to a pro-wedge dominant strain distribution, accompanied by a temporary quiescence of the retro-wedge.

Comparative analysis of pro-and retro-wedge behaviour reveals the self-organisation of a complex orogenic system. Major changes or events (change in strain distribution) may be due to internal intrinsic thresholds being reached (e.g. onset of continental collision) and therefore may not rely on external forcing (plate kinematics). We believe that the application of the same detailed approach in other orogens can potentially identify similar self-organisation and intrinsic thresholds involved in orogenesis.

Introduction

he distribution of shortening between the pro-wedge (lower plate) and retro-wedge (upper plate) varies greatly between orogens around the world. Distributions range from the highly asymmetric Zagros [Molinaro et al., 2005;[START_REF] Mouthereau | Tertiary sequence of deformation in a thin-skinned/thick-skinned collision belt: he Zagros Folded Belt (Fars, Iran): Tectonics[END_REF], to the symmetric High Atlas [START_REF] Beauchamp | Inversion tectonics and the evolution of the High Atlas Mountains, Morocco, based on a geological-geophysical transect[END_REF]Arboleya et al., 2004]. Shortening distribution can also vary strongly along strike, as in the Alps [Rosenberg and Kissling, 2013]. he factors that control the distribution of shortening and how it evolved through time are currently poorly understood. However, models have shown qualitatively that rit inheritance can inluence orogen crustal structure, promoting propagation of deformation into the retro-wedge [e.g., Jammes and Huismans, 2012;Erdős et al., 2014]. Determining whether this alters the distribution of shortening between the pro-and retro-wedge requires further quantitative analysis.

Another poorly understood aspect of the crustal structure of orogens is the formation of crustal scale antiformal stacks: long crustal thrust sheets with large ofsets that are stacked and folded into a crustal scale anti-form, such as the Lepontine dome in the Alps [e.g., Schmid and Kissling, 2000] and Axial Zone in the Pyrenees [e.g., Muñoz, 1992]. Proposed mechanisms include increasing the critical taper through strength changes along the mid-crustal décollement [Beaumont et al., 2000], or triggering underplating by reducing taper with erosion of the internal thick-skinned pro-wedge [Rushlow et al., 2013] and/or by sediment blanketing of the external thick-skinned pro-wedge [Sinclair et al., 2005]. Considering that surface processes can change the distribution of loads within an orogenic ediice, can ongoing thinskinned deformation inluence the deeper thick-skinned structure? Although thin-skinned deformation is generally understood to be driven by thick-skinned deformation, the potential for a feedback efect on the crustal structure warrants investigation.

Here we build on previous work [Erdős et al., 2014] that documents the efect of rit inheritance and investigate the combined roles of rit inheritance, a highly mobile cover and surface processes on the interaction between thin-and thick-skinned deformation across a doubly vergent orogen. We use highresolution 2D thermo-mechanical numerical models to study the efect of fault-to lithosphere-scale dynamics on the crustal-scale shortening distribution including the formation of a crustal scale antiformal stack. We quantify the shortening distribution through time and compare our results with observations from the Pyrenees.

Methods

We use a modiied 2D version of FANTOM, a high resolution, thermo-mechanically coupled, arbitrary Lagrangian-Eulerian, inite element code [hieulot, 2011]. he high resolution allowed us to model deformation at the lithospheric scale while still resolving good detail in the fold-and-thrust belts and associated sedimentary basins. Localisation of deformation is incorporated through strain weakening of frictional-plastic materials. Surface processes were included with simple algorithms for moderate elevation dependent erosion and sedimentation (up to 0 m base level) [e.g., Erdős et al., 2014;Appendix B]. Erosion and deposition are not mass balanced. Well-developed retro-wedge, narrow thick-skinned pro-wedge and wide thin-skinned pro-foreland foldand-thrust belt.

」 セ ᄋ M ᄋ ⦅ ᄋ ⦅ ᄋ ⦅ M ⦅ N ⦅ N j ヲ ュ t ヲ ュ セ Q セ i セ A ヲ セ セ セ セ Z Z ュ Z Z M M [ [ ↑ Z [ M M N M イ N イ ⦅ ᄋ ⦅ ᄋ ⦅ セ M M セ ᄋ M ᄋ⦅ ᄋ Z Z ⦅ ᄋ @ セ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ Z Z Z Z Z Z Z @
.......... : --; " " ' \" : -, . , . ::.

•. -. •. -. • セ [ [ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ ᄋ A @ Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z Z セ Z Z Z Z Z セ @
. ..:;::::=:::::: :

Results

Model 1:

Efects of extensional inheritance, salt, and surface processes. Model 1 is characterised by a three-phase evolution. Phase 0 is the pre-orogenic rit phase, where at t = 5 My, 50 km of extension results in a narrow, roughly symmetrical rit bounded by two conjugate crustal scale frictional-plastic shear zones that root in the weak mid-crust (Figure 3.1a). Two conjugate frictionalplastic shear zones in the lithospheric mantle accommodate extension in the mantle. he sedimentary cover has gravitationally slid of the rit margins. No sediments were deposited in the rit.

During Phase 1 initial lithospheric shortening leads to symmetric inversion of the rit zone. Inherited extensional shear zones are preferentially reactivated during the irst 25 km of convergence, restoring crustal thickness, followed by uplit of a symmetric central block (Figure 3.1b). he strain rate distribution shows that both conjugate shear zones in the upper crust and upper mantle lithosphere are simultaneously active during this irst inversion phase.

Phase 2 comprises the development of an asymmetric crustal scale orogeny (Figure 3.1c,d,e). Following symmetric inversion, strain localisation on a single large-scale shear zone initiates asymmetric subduction of lower crust and mantle lithosphere and abandonment of other, previously active shear zones (Figure 3.1c). Initial narrow thin-skinned fold and thrust belts root back into both the deep shear zones and the cover sliding of the upliting keystone block. At t = 15 My and 95 km of convergence, syn-tectonic sedimentation in a wedge-top basin promotes the creation of wide thin-skinned thrust sheets in the external pro-wedge (Figure 3.1d). Below and at the same time, thick-skinned deformation propagates outward in the pro-wedge to accrete a new upper crustal thrust sheet. Over-steepening of the antiformal stack trailing edge reactivates the retro-wedge (Figure 3.1d).

Ater 230 km of convergence (28 My; Figure 3.1e) eicient decoupling along the salt décollement together with wedge-top sedimentation facilitate outward propagation of the thin-skinned pro-wedge fold-andthrust belt. he internal antiformal stack continues to grow by basal accretion of crustal scale thrust sheets and the upliting internal zone is continually eroded. he contrasting behaviour resulting from a less eicient décollement horizon and absence of extensional inheritance is illustrated in models DR1 and DR2 [Appendix B; see also Erdős et al., 2014Erdős et al., , 2015]]. Model DR1 includes extensional inheritance and surface processes, but has a stronger shale décollement horizon as compared to Model 1. It exhibits a clear coupling between shallow and deep structures. As deformation migrates outward in the pro-wedge, thin-skinned leading edge structures root back into an active crustal scale shear zone. hese thin-skinned structures are then carried passively above the next crustal imbricate as deformation propagates forward. he décollement level is therefore active only at its leading edge at any one time, linking back to a single crustal structure. In contrast to Model 1 an antiformal stack does not form, however, retro-wedge development is very similar to Model 1 (Figure B.2). Supplementary model DR2 combines absence of rit inheritance with surface processes and a less eicient shale décollement. Subduction asymmetry is created with the irst shear zone, resulting in a highly asymmetric crustal scale pro-wedge orogen and negligible retro-wedge shortening (Figure B.3). he pro-wedge deformation style is the same as in DR1.

Shortening distribution

Quantitative analysis of shortening distribution in the pro-and retro-wedges reveal characteristic diferences between Model 1 and supplementary models DR1 and DR2. Models 1 and DR1 (with rit inheritance) show initial symmetric inversion during Phase 1 with 50% of total shortening in the future retro-wedge. During asymmetric orogenic growth in Phase 2, the models with extensional inheritance exhibit ~18-20 % shortening in the retro-wedge (Figure 3.2a). Retro-wedge shortening rates show a similar pattern with an initial high, on average 50% shortening rate corresponding to symmetric inversion during Phase 1, a short phase during which the retro-wedge is abandoned, followed by reactivation of the retrowedge at 20% of the total shortening rate (Figure 3.2b). In contrast, model DR2 without rit inheritance exhibits only minor retro-wedge shortening stabilising at about 6% of total shortening (Figure 3.2a) and correspondingly low retro-wedge shortening rate 3-9% ater 50 km of shortening (Figure 3.2b).

Discussion

he models presented here identify three main factors that control the structural style of mountain belt formation: 1) rit inheritance, 2) eiciency of decoupling between thin-skinned and thick-skinned deformation, and 3) syn-tectonic erosion and sedimentation.

Rit inheritance promotes more shortening in the retro-wedge and generates two phases, each with a distinct distribution of shortening. Phase 1 is characterised by reactivation of inherited upper crustal extensional frictional-plastic shear zones. In the model presented here the pre-existing rit symmetry results in symmetric inversion. Phase 1 structural styles of crustal shortening will depend on inherited extensional crustal geometries. Lithosphere scale asymmetric subduction and accretion is the primary characteristic of Phase 2 mountain building. Without extensional inheritance asymmetry at depth is immediately established and an asymmetric crustal orogen focuses most shortening on the pro-wedge with limited retro-wedge activity. Extensional inheritance leads to a two-phase shortening history with early inversion across both wedges. Further shortening in the thick-skinned retro-wedge leads to transport of crustal material onto the retro-wedge (Figure 3.3a).

Although décollement rheology has a negligible direct impact on the distribution of shortening between the pro-and retro-wedge, it does provide a primary control on deformation style. A weak salt décollement allows very eicient decoupling leading to the formation of a broad low taper thin-skinned fold and thrust belt in the pro-wedge (Figure 3.3a). his efectively reduces the pro-wedge taper to subcritical and thus promotes thick-skinned underthrusting in the core of the orogenic wedge, resulting in a narrower inner orogen zone. he impact of erosion and sedimentation on the crustal structure of orogens is well-known [e.g., Willett et al., 1993;[START_REF] Braun | Structural evolution of a three-dimensional, inite-width crustal wedge[END_REF]Erdős et al., 2014Erdős et al., , 2015]]. Erosion is most eicient in the core of the orogen, reducing critical taper and resulting in a narrower orogen (Figure 3.3a). Sedimentation increases the load in the foreland (Figure 3.3a), reduces critical taper and promotes longer thin and thick-skinned thrust sheets with greater ofset along each thrust [e.g., Erdős et al., 2015]. he efects of a highly mobile cover and surface processes work in tandem to create a wide pro-wedge fold-and-thrust belt and promote a crustal scale antiformal stack (Figure 3.3a), without needing additional complexities such as stacked crustal weakness zones or lateral strength variations of the mid-crustal décollement [Beaumont et al., 2000].

Pyrenean double vergent orogen

We next describe the characteristics of the Central Pyrenean type example of a mountain belt that resulted from inversion of a salt rich, narrow hyper extended rit (Figure 3.3a) [e.g., Muñoz, 1992;Jammes et al., 2009]. Features that require explanation include: (1) a two stage mountain belt evolution with initial symmetric inversion, followed by largely asymmetric shortening with about 80 % in the pro-wedge and 20% in the northern retro-wedge; (2) formation of a narrow crustal scale pro-wedge antiformal stack in the core of the orogen; (3) a broad pro-wedge fold and thrust belt with minor thick-skinned thrusting at depth and ( 4) a narrow thick-skinned retro-wedge. Previous work suggests that complex inherited weak zones and highly eicient erosion played a key role in the evolution of these geometries [Beaumont et al., 2000]. he numerical models presented here demonstrate that these key features can be explained by the combined efects of eicient erosion of the elevated internal zone of the orogen and a weak décollement layer at the base of the thin-skinned fold and thrust belt, promoting basal accretion of thick-skinned basement thrust sheets at depth in the core of the orogen, and a wide low taper pro-wedge basin. Deposition of material across the low taper foreland basin areas reinforces the interaction and feedback between thin and thickskinned deformation. Extensional inheritance is an essential element that allows for shortening in both pro-and retro-wedge. Notably the models do not require complex inherited weaknesses. Similarly, the combined efects of orogenic erosion and low taper thin-skinned fold and thrust belts with syn-tectonic wedge top sedimentation may explain the antiformal structure of the Lepontine dome in the Central Alps and the adjacent wide foreland basin in North Alpine foreland [Schmid et al., 1996]; and comparable [1992] and Beaumont et al. [2000].

structures in the frontal zone of the Himalayan mountain belt [START_REF] Beaumont | Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to focused surface denudation[END_REF][START_REF] Searle | he structural geometry, metamorphic and magmatic evolution of the Everest massif, High Himalaya of Nepal-South Tibet[END_REF][START_REF] Decelles | Along-strike continuity of structure, stratigraphy, and kinematic history in the Himalayan thrust belt: he view from Northeastern India: Tectonics[END_REF].

Conclusions

We have shown how forward dynamic models provide a general explanation for the evolution of Alpine type mountain belts as exempliied by the Central Pyrenean type example with formation of an crustal scale antiformal stack in the axial zone, adjacent wide low taper fold and thrust belts decoupled on weak salt, and eicient syn-orogenic erosion and deposition. We have also shown that rit inheritance promotes the two-phase evolution with early symmetric inversion followed by asymmetric mountain building. Rit inheritance is also an essential element for forming a doubly vergent orogenic wedge that accommodates >10% of total convergence in the retro-wedge. he models show that formation of an antiformal stack of crustal scale thrust sheets and adjacent low taper thin-skinned fold and thrust belts can be explained by the interaction and feedback between: (1) a weak décollement layer decoupling thin and thick-skinned deformation and resulting in low taper foreland basin deformation, (2) eicient erosion of the internal thick-skinned pro-wedge, and 3) wedge top deposition of sediments in the pro-wedge.

Introduction

Collisional orogens can accommodate convergence by shortening their lower plate (pro-wedge) and shortening their upper plate (retro-wedge) [e.g., Willett et al. 1993]. he strain partitioning between the proand retro-wedge varies greatly between natural systems. An orogen can be singly vergent, accommodating all shortening in the pro-wedge, such as the Zagros [e.g., Molinaro et al., 2005;[START_REF] Sherkati | Central Zagros fold-thrust belt (Iran): New insights from seismic data, ield observation, and sandbox modeling[END_REF]Allen et al., 2013], and most of the Greater Caucasus [e.g., [START_REF] Saintot | he Mesozoic-Cenozoic tectonic evolution of the Greater Caucasus[END_REF]. Alternatively, orogens can be doubly vergent, accommodating shortening in both the pro-and retro-wedge, such as the Pyrenees [e.g., Muñoz, 1992], and Tien Shan [e.g., [START_REF] Loury | Crustal-scale structure of South Tien Shan: implications for subduction polarity and Cenozoic reactivation[END_REF]. In doubly vergent orogens, the pro-retro strain partitioning can even vary along strike, such as in the Alps [Rosenberg and Kissling, 2013], and in time, as in the Pyrenees [Grool et al., 2018]. What controls this variation in space and time is currently poorly understood.

Most investigations into the controlling factors of this variation using forward models have been qualitative, focusing only on the distinction between singly and doubly vergent orogens. First order controlling factors that enable double vergence are rit inheritance [Jammes and Huismans, 2012;Erdős et al., 2014; see also Chapter 3], strength contrasts between the upper and lower plate [Vogt et al., 2017b], and decoupling of the upper and lower crust of the lower plate [Vogt et al., 2017a]. Other factors have been found to control the crustal structure but not orogen vergence, such as sedimentation and erosion resulting in a wider and narrower thick-skinned orogen, respectively [Erdős et al. 2015]. In combination with these surface processes, the crustal structure can also be inluenced by thin-skinned deformation, which depends on the décollement rheology (see Chapter 3).

hese factors that inluence the crustal structure but cannot themselves control orogen vergence may still inluence the pro-retro strain partitioning within a doubly vergent orogen. So far, this inluence has not been quantiied for these factors individually. Especially the inluence of thin-skinned deformation is currently poorly understood. Décollement rheology and distribution provide major controls on the thin-skinned deformation, as understood through critical taper theory, models, and natural examples [e.g., Dahlen, 1990;Ford, 2005;Buiter, 2012]. herefore, testing the inluence of décollement rheology and distribution on the thick-skinned structure and pro-retro strain partitioning may shed light on the role of thin-skinned deformation. Testing combinations of multiple factors may shed light on their interaction. he pro-retro strain partitioning may also be inluenced by interaction of the pro-and retro-wedge themselves. How the pro-and retro-wedge interact is not clear, and highly dependent on one's deinition of what constitutes the boundary between them. For practical purposes, we deine the pro-and retro-wedge as all material derived from the lower and upper plate respectively, rather than placing the boundary at the topographic drainage divide as in the deinition of Willett et al. [1993]. Using our deinition, there can be no material transfer between wedges. However, some studies suggest there is a link in the timing of deformation of both wedges [Sinclair et al. 2005;Hoth et al. 2007; see also Chapter 3].

We test various combinations of rit inheritance, décollement rheology, décollement distribution, and surface processes. We individually quantify the inluence of these factors on the pro-retro strain partitioning and how they interact. he quantiied strain partitioning also allows us to investigate the interaction between the pro-and retro-wedge through time. We then verify our quantiied pro-retro strain partitioning results by comparing to natural systems.

Methods

For our experiments we used a modiied version of FANTOM, an arbitrary Lagrangian-Eulerian initeelement code [hieulot, 2011]. We ran high-resolution experiments featuring 2D, thermo-mechanically coupled, visco-plastic low. Flow is governed by the semi-static Stokes equation, assuming incompressibility and negligible inertial forces. In the viscous regime, materials follow a power-law rheology:

(4.1)
where is the strain rate tensor, deined as the spatial derivatives of the velocity ield, is the preexponential material constant, the viscous creep stress, the exponent, the activation energy, is pressure, the activation volume, the universal gas constant and is temperature. Material constants , , , and are determined from laboratory data for wet quartzite, wet olivine and dry olivine rheologies [Karato and Wu, 1993;Gleason and Tullis, 1995]. In the plastic regime, the material's strength follows the Drucker-Prager yield criterion, equivalent to Coulomb failure in 2D. his is given by the yield stress (4.2) where is the second invariant of the deviatoric stress tensor , is cohesion, and is the efective angle of internal friction. For visco-plastic materials, the model automatically chooses the appropriate low regime by computing an efective viscosity with the lowest available stress ( or ):

(4.3)
where is the second invariant of the strain rate tensor. he thermal structure is governed by the heat equation (4.4) where is density, the heat capacity, is time, the spatial dimensions, the thermal conductivity, and is radiogenic heat production. Density variations due to temperature are taken into account. A scaling factor f is included in the pre-exponential material constant. he conversion is A=f -n A base .

Chapter 4. Numerical modelling of strain partitioning 76 Figure 4.1 shows the initial coniguration of our models. he model domain is 1200 km wide and 600 km deep, and contains 2400x1200 Eulerian cells. he resolution is varied along the vertical axis to achieve a resolution of 500x200 m in the top 25 km. Pre-orogenic sediment occupies the top 3 km, underlain by a 1 km thick décollement layer and 21 km thick upper crust basement. he lower crust is 10 km thick, placing the Moho at 35 km depth. he Lithospheric mantle lies below until 120 km depth. Material properties for each layer are listed in Table 4.1. A 6x6 km weak seed of pre-strain weakened material is placed in the lower crust at the centre of the model. Velocity boundary conditions are imposed on both sides, introducing lithosphere material at 5 mm/y on each side while material is extracted below to keep the volume constant. he top is a free surface and the bottom allows free slip. he initial thermal proile is laterally uniform, increasing parabolically from = 0° C at the surface to = 550° C at the Moho, and linearly to = 1330° C at the base of the Lithosphere. he top and bottom boundaries are kept at = 0° C and = 1330° C, respectively while the side boundaries are thermally insulated. he initial coniguration of our models is completely symmetrical unless speciied otherwise, and the plate polarity is random. In our igures we have mirrored some models so all appear with the same polarity for ease of comparison. In our models we have tested the inluence of rit inheritance, sedimentation and erosion, décollement rheology and décollement distribution (Table 4.2). Rit inheritance was implemented by inverting the velocity boundary conditions during the irst 5 My, leading to 50 km of extension [Jammes and Huismans, 2012]. We used a simple sedimentation algorithm that ills all available space up to a base level of 0 m, starting at 10 My [e.g., Erdős et al., 2014]. Erosion is altitude dependent, following , where is altitude, is time and is the inverse erosional timescale (in s -1 ). was chosen so that a 4 km high topography erodes by 1 km in 2 My [Erdős et al., 2014]. Eroded and deposited volumes were not linked. he inluence of décollement rheology was tested by using two possible décollement rheologies: a reference shale-like visco-plastic rheology with an efective angle of internal friction = 2° (M1-M5), or a weak salt-like linear viscous rheology with an efective viscosity = 10 19 Pa•s (M6-M8), similar to estimates for halite [e.g., van Keken et al., 1993;[START_REF] Chemia | Numerical modelling of rise and fall of a dense layer in salt diapirs[END_REF]. We tested the inluence of the décollement distribution by removing the décollement from half the model (M5), replacing it with regular pre-orogenic sediment.

We quantify the strain partitioning between the pro-and retro-wedge by tracking the horizontal position of a column of points that originate at the centre of the model (see also Chapter 3). his column is displaced towards the side that accommodates the most shortening, and the amount of displacement is a measure for the ratio of shortening accommodated on either side. Note that this does not distinguish between the pro-and retro-wedge as originally deined by Willett et al. [1993], where the boundary between the proand retro-wedge is a topographic drainage divide. Instead, our method tracks a tectonic boundary that originates at the plate boundary or S-point. his facilitates comparison with natural systems, where the position of this boundary can be reconstructed.

Results

Here we describe our model results in two sets (Table 4.2). Set 1 contains all models with a shale décollement (M1 to M5), and set 2 contains all models with a salt décollement (M6 to M8). Some of these models are identical or very similar to previously published models, and are used as reference against the new models to provide insight into the inluence of the décollement rheology and distribution. Amounts of convergence are given as 'convergence since irst shear zone' for M1, or 'convergence beyond full inversion of the extensional shear zones' for M2 to M8. skinned pro-wedge thrust sheets are accreted by in-sequence formation of new shear zones every ~31 km of additional convergence, forming a wide thick-skinned pro-wedge (Figure 4.2a). In contrast, only a single, poorly developed retro-ward shear zone that accommodates very little shortening is formed. his gives the orogen a highly asymmetric, one-sided appearance. hin-skinned deformation is mostly active in the footwall of the frontal thick-skinned thrust, and abandoned once a newer, more distal thick-skinned thrust is formed.

During the irst 5 My in M2, extension has resulted in a narrow, symmetric rit bounded by a conjugate pair of shear zones in the upper crust and lithospheric mantle. Inversion occurs symmetrically along these inherited shear zones, upliting the central block and over-inverting the shear zones past 6.25 My. A new shear zone is formed on one side ater 15 km of convergence beyond full inversion, creating an asymmetry that determines the plate polarity. he retro-wedge is abandoned and all shortening is accommodated by subduction of the lower plate. Ater 65 km of convergence, a new pro-wedge thrust sheet is accreted, upliting the older, upper pro-wedge thrust sheet and reactivating the retro-wedge. More pro-wedge thrust sheets are accreted in-sequence every ~33 km, while older, upper thrust sheets slowly migrate onto the upper plate. Most thick-skinned pro-wedge thrusts have a forelandward vergence, but one backthrust is created at 97.5 km of convergence. Continued convergence creates a doubly-vergent orogen with a wide thick-skinned pro-wedge and well-developed retro-wedge (Figure 4.2b). he hinterland of the pro-wedge rests on top of the upper plate. Like in M1, thin-skinned deformation is only active in the footwall of the frontal thick-skinned thrust. he irst 10 My of M3 are identical to M2, except the plate polarity happens to be reversed. he rit is inverted symmetrically, followed by an asymmetric phase with an abandoned retro-wedge ater 15 km of convergence beyond full inversion. Sedimentation is started at 10 My, illing the pro-and retro-foreland with ~4 and ~1 km of sediment, respectively. Ater 65 km of convergence, a new thick-skinned pro-wedge thrust sheet is accreted in sequence every ~34 km. Older, upper thrust sheets migrate onto the upper plate, reactivating the retro-wedge. Four out of a total of 10 thick-skinned pro-wedge thrusts develop with a hindward vergence, at 100, 120, 230, and 287.5 km of convergence. Ater 300 km of convergence, a doublyvergent orogen is formed, with a wide, thick-skinned pro-wedge overlain by a thick sedimentary cover and a well-developed retro-wedge (Figure 4.2c). hrust sheets, both thick-skinned and thin-skinned, appear longer than in M2 but this cannot be measured with conidence due to the high variability in thrust sheet length within each model. hin-skinned deformation is only active in the footwall of the frontal thickskinned thrust and is mostly abandoned once a new thick-skinned frontal thrust is formed.

Model M4 adds erosion. he evolution is similar to previous models, with symmetric rit inversion and the onset of asymmetry at 15 km of convergence beyond full inversion. At 10 My, the sedimentary cover of the uplited central block is almost fully eroded, and sedimentation ills the pro-and retro-foredeep with ~4 and ~1 km of sediment, respectively. From 67.5 km of convergence onwards, pro-wedge thrust sheets are accreted in sequence every ~36 km of convergence. Older-upper pro-wedge thrust sheets are uplited and migrate onto the upper plate, reactivating the retro-wedge, while also eroding slowly. Ater 300 km of convergence, this creates a narrower doubly vergent orogen than M2 and M3. Basement is exposed in the centre of the orogen. he thin-skinned wedges are also smaller.

M5 is the only asymmetric model, with a shale décollement on the let side and no décollement on the right side. he efects of this asymmetry are already apparent in the riting stage, where the rit becomes slightly asymmetric, accommodating more extension along the let-hand (décollement side) shear zone. Ater symmetry is almost fully restored at 6.25 My during inversion, the other shear zone (non-décollement side) becomes dominant. At 10 km of convergence beyond full inversion, a new shear zone is formed on the décollement side, creating the deinitive asymmetry with a décollement in the pro-wedge (Figure 4.3a) and no décollement in the retro-wedge (Figure 4.3b). Further shortening is accommodated by subduction of the lower plate while the retro-wedge extends slightly. Ater 70 km of convergence, a new pro-wedge thrust sheet is accreted every ~36 km, as the older, upper pro-wedge thrust sheets slowly migrate onto the upper plate and the retro-wedge is very slowly reactivated (Figure 4.3c). At 237.5 km, a new thick-skinned retro-wedge thrust is formed (Figure 4.3d). Ater 300 km of convergence, this creates a doubly vergent orogen with frontal accretion in both the pro-and retro-wedge. hin-skinned deformation is not present in the retro-wedge, owing to the lack of a décollement layer.

4.3.1.2 Set 2: Salt décollement models he salt décollement in M6 allows the sedimentary cover to slide into the rit during the irst 5 My, exposing the rit shoulders. he rit is then symmetrically inverted along the two inherited shear zones, past full inversion at 6.5 My. As the keystone is uplited, the cover slides of towards both sides under the inluence of gravity. At 15 km beyond full inversion, a new shear zone is created, establishing an asymmetry and abandoning the retro-wedge (Figure 4.4a). All shortening is accommodated by subduction of the lower plate until 80 km beyond full inversion. During this period, the sedimentary cover of the lower plate is stacked in the pro-foreland. A new thrust sheet is accreted into the thick-skinned pro-wedge at 80 km beyond full inversion, upliting the upper, older thrust sheet and keystone block, and reactivating the retro-wedge. he stack of thin-skinned thrust sheets slides of into the pro-foreland as the new thrust sheet is uplited (Figure 4.4b). As the keystone and upper pro-wedge thrust sheets slowly migrate onto the upper plate, thin-skinned thrust sheets are stacked in front of the basement blocks in the retro-wedge (Figure 4.4c). he process of new pro-wedge thrust sheets upliting and the cover sliding into the foreland repeats every ~43 km of convergence, creating long thick-skinned pro-wedge thrust sheets with large ofsets (Figure 4.4d). Two new thick-skinned retro-wedge thrusts are formed at 245 and 285 km of convergence, respectively. Ater 300 km of convergence, a doubly vergent orogen with a narrow thick-skinned prowedge and a wide thick-skinned retro-wedge is formed. he highly mobile cover has allowed wide and thick thin-skinned fold-and-thrust belts to form in both forelands, the pro-foreland belt being the largest.

During the irst 10 My, M7 develops very similar to M6. he rit is symmetrically inverted, with the sedimentary cover sliding of the keystone block ater reaching full inversion of the inherited shear zones at 6.25 My. At 17.5 km beyond full inversion, a new thrust is formed, creating an asymmetry and abandoning the retro-wedge. he sedimentary cover of the pro-wedge slides into the pro-foreland. However, as soon as sedimentation is started at 10 My (37.5 km), the pro-foreland thin-skinned deformation front jumps outward to a more distal position, forming a wide wedge-top basin. Ater 80 km of convergence beyond full inversion, a new pro-wedge thrust is formed, upliting the older, upper thrust sheet. he retro-wedge is reactivated and the sedimentary cover of the pro-wedge slides downwards into the pro-foreland. his is associated with another forelandward jump of the pro-foreland thin-skinned deformation front, creating a new wedge-top basin (Figure 4.5a). he upper pro-wedge thrust sheet and keystone block migrate onto the upper plate, forming a narrow thin-skinned retro-foreland fold-and-thrust belt in front (Figure 4.5b). he new pro-wedge thrust sheet at 80 km is quickly followed by another at 97.5 km of convergence. Gravitational sliding has exposed basement in the core of the orogen (Figure 4.5c). New pro-wedge thrusts form in pairs, alternating between a long interval (>60 km) and a short interval (<20 km), forming an average of ~45 km of convergence between thrusts. his forms a doubly vergent orogen with a thick-skinned pro-wedge that consists of alternating wide and narrow thrust sheets, and a well-developed retro-wedge. A very wide thinskinned pro-foreland fold-and-thrust belt is characterised by wide wedge-top basins separated by narrow zones where thin-skinned deformation is concentrated (Figure 4.5d).

Erosion is included in model M8. he early evolution is almost identical to other models with a salt décollement: symmetric inversion of the rit, reaching full inversion at 6.25 My. he cover slides of the keystone block as it is uplited, and at 17.5 km beyond full inversion a new thrust is created that determines the asymmetry and abandons the retro-wedge. At 10 My (37.5 km) sedimentation is started, illing the proforeland and triggering the thin-skinned deformation front to jump outward in the pro-foreland, creating a wide wedge-top basin between two zones of intense deformation (Figure 4.5e). At 80 km of convergence, As each pro-wedge thrust sheet is uplited, the cover slides down into the pro-foreland and upper thrust sheets are partially eroded. At 252.5 km of convergence, a new thick-skinned thrust is formed in the retrowedge. 300 km of convergence have created a narrower orogen compared to M7, with a relatively narrow thick-skinned pro-wedge with alternating long and short thrust sheets, and a well-developed retro-wedge (Figure 4.5h). A wide pro-foreland fold-and-thrust belt has wide basins separated by narrow zones of intense deformation. Erosion has exposed basement along a signiicant portion of the central orogen. (a) plots of the distribution of shortening between the lower (pro) and upper (retro) plate for models with no surface processes. Strain distributions for the High Atlas, Pyrenees, Western Alps and Central Alps from [START_REF] Teixell | Tectonic shortening and topography in the central High Atlas (Morocco)[END_REF]Beaumont et al., 2000;Schmid and Kissling, 2000;Schmid et al., 1996]. (b) plots of the distribution of shortening between the lower and upper plate for models with surface processes, and M1 as reference. (c) Distribution of shortening rates as percentage of the total convergence rate between the lower and upper plate. his plot shows the distribution for M4 as an example. All models feature a similar evolution of the distribution through time.

Strain distribution analysis

Here we describe the quantiied strain distribution between the pro-and retro-wedge in our models. hese results are plotted against convergence, not time, to eliminate any diferences in convergence rates when comparing to natural systems later on. Our irst observation is that the strain distribution varies through time, but eventually tends to stabilise ater 75 to 125 km of convergence (Figure 4.6a). he shortening distribution in M1, without rit inheritance, rapidly shits in favour of the pro-wedge, stabilising at only ~6% of shortening accommodated in the retro-wedge (Figure 4.6a). Every model that includes rit inheritance has accommodated more shortening in the retro-wedge than M1. his conirms that rit inheritance is essential to promoting retro-wedge shortening. he asymmetric décollement distribution we tested (M5) inluences the polarity of subduction, favouring a décollement in the lower plate. As a result, the retro-wedge does not have a décollement layer and shortening is decreased to ~11% of total convergence. his is lower than other models that also include rit inheritance. When surface processes are not included, the décollement rheology makes a signiicant diference in the shortening distribution. A shale décollement results in the retro-wedge accommodating ~18% of total convergence, while a salt décollement allows accommodating ~24% of total convergence in the retro-wedge (Figure 4.6a). However, when sedimentation and erosion are included, the diference in strain distribution between a shale décollement and a salt décollement is greatly reduced (Figure 4.6b). he strain distribution varies a few percent between these models, but they all accommodate 16-20% of total convergence in the retro-wedge. Erosion appears to systematically result in slightly lower retro-wedge shortening, but this diference is small and may be within the error margin.

We have also quantiied the shortening rates of the pro-and retro-wedge in our models (Figure 4.6c). Shortening rates of the pro-and retro-wedge are not constant, even during the later stages of orogeny when the shortening distribution has stabilised. On a larger timescale, the shortening rates reveal a twophase evolution in pro-retro strain partitioning: During the symmetric rit inversion phase, shortening is partitioned roughly 50/50 between both sides. he retro-wedge is then temporarily abandoned during the transition to the asymmetric collision phase. Later in the asymmetric collision phase, the retro-wedge is reactivated but the shortening rate remains low. Shortening rates also vary on a shorter timescale: Every time a new pro-wedge thrust is formed, the pro-wedge shortening rate is increased briely. he retro-wedge shortening rate lowers accordingly at these times. Events in the pro-wedge thus inluence the shortening rate of the retro-wedge and vice versa. he two are inextricably linked. 1 Individual factors Rit inheritance enables an orogen to become doubly vergent, by creating a weak shear zone through the retro-wedge that facilitates shortening [Jammes and Huismans, 2012;Erdős et al., 2014; see also Chapter 3]. Our quantiied results conirm that rit inheritance has a irst order inluence on creating a doubly vergent orogen. he inherited retro-shear zone is easily reactivated during the inversion phase, and later easily reactivated again during the collision phase. his results in a doubly vergent orogen, rather than a singly vergent orogen without rit inheritance (Figure 4.7a).

Our quantiied results show that the décollement rheology can make a diference in the pro-retro strain partitioning and crustal structure (Figures 4.6a and 4.7). A shale décollement enables thin-skinned foreland fold-and-thrust belts to form. However, because the shale décollement does not allow gravitational sliding, the sedimentary cover remains on top of the thick-skinned thrust sheets, promoting a wide thick-skinned orogen (Figure 4.7b). In contrast, a salt décollement promotes gravitational sliding of the sedimentary cover. his unloads the core of the orogen and increases overburden over the distal pro-and retro-wedge, resulting in a narrower thick-skinned orogen (Figure 4.7c). he salt décollement also reduces friction along the retro shear zone, which, combined with the weight of the pro-foreland fold-and-thrust belt results in more shortening being accommodated in the retro-wedge (Figure 4.6a). For maximum efect, the décollement must be weak enough for gravitational sliding to occur, but not so weak that the resulting foldand-thrust belt becomes very wide and thin. his is conirmed by models with a weaker salt décollement ( = 10 18 Pa•s) that we describe in Appendix C.

An asymmetric décollement distribution inluences the plate polarity, preferentially placing the décollement in the lower plate (Figure 4.3). his can be explained by comparing the integrated strength of the crust on both sides of the model. he décollement lowers the integrated strength of the crust. hus, all else being equal, when a new shear zone is formed ater inversion of the rit, it is highly likely that it forms on the décollement side. his new shear zone is what determines the plate polarity and the décollement thus ends up in the lower plate. he lack of a décollement in the upper plate results in increased friction along retrowedge shear zones compared to the pro-wedge, afecting the pro-retro strain partitioning (Figure 4.6a). Without a décollement the structure of the retro-wedge is diferent, with reduced fault throws and earlier frontal accretion compared to a retro-wedge with a décollement (Figure 4.3). With these results, we can speculate about the efects of other possible décollement distributions. he plate polarity is only inluenced by the décollement distribution around the plate boundary. A laterally limited décollement will alter the properties of both wedges once they reach the edge of the décollement, potentially inluencing the proretro strain partitioning, favouring the side with the largest proportion of décollement. Because thickness of a viscous décollement is a key parameter in determining the cover mobility, a thicker décollement will further facilitate gravitational sliding and reduce taper of the resulting thin-skinned wedge while a thinner décollement will have the opposite efect.

he inluence of surface processes on the crustal structure is well-documented [e.g., Buiter, 2012;Erdős et al., 2014Erdős et al., , 2015]]. Sedimentation results in longer thrust sheets by increasing the overburden of the distal pro-wedge and erosion unloads the core. Together these promote shortening of the hinterland, resulting in a narrower thick-skinned orogen (Figure 4.7c, d) [Erdős et al., 2015].

Interactions

Although surface processes do not appear to inluence the pro-retro strain partitioning when combined with a shale décollement, they do have an inluence when combined with a salt décollement. In fact, combining surface processes and a salt décollement results in a very similar pro-retro strain partitioning as a shale décollement model. Surface processes thus efectively counteract the efect of a salt décollement on the pro-retro strain partitioning. he results show that sedimentation without erosion already has this efect. Sedimentation stabilises the sedimentary cover in the salt décollement models while still allowing gravitational sliding (Figure 4.5). One would expect the greater mass of sediment in the pro-foreland to have a greater efect on the strain partitioning than the relatively little amount of sediment in the retro-foreland, but this is not the case. It is unclear why sedimentation results in a reduction of retro-wedge shortening relative to a salt décollement without sedimentation instead. he evidence that the thick-skinned structure can be inluenced by surface processes changing and redistributing the loads on top of an orogenic wedge has been steadily growing [e.g., Sinclair et al., 2005;Fillon et al., 2013]. hin-skinned deformation can also redistribute signiicant loads, but the relationship between thin-skinned and thick-skinned deformation is usually considered a one-way street, with thickskinned deformation driving thin-skinned deformation without any feedback [e.g., [START_REF] Bergh | Interaction of basement-involved and thin-skinned tectonism in the Tertiaru Fold-and-thrust belt of central Spitsbergen Svalbard[END_REF]Molinaro et al., 2005]. Our models show that the décollement rheology and resulting thin-skinned structure have a clear efect on thick-skinned structure. hin-skinned deformation is still driven by uplit of thick-skinned thrust sheets, but if the décollement is weak enough to enable signiicant gravitational sliding, thin-skinned deformation can inluence the thick-skinned structure. Gravitational sliding of the sedimentary cover results in a narrower thick-skinned pro-wedge.

Our results allow us to link the evolution of the pro-and retro-wedge in two ways. According to the original deinition of the pro-and retro-wedge [Willett et al., 1993], material migrates from the pro-wedge to the retro-wedge. Using our deinition instead, the hinterland of the pro-wedge migrates onto the upper plate during the asymmetric collision phase. his migration reactivates the retro-wedge. hus, with either deinition, growth of the pro-wedge drives retro-wedge shortening during the asymmetric collision phase, but not the other way around. Shortening rates of both wedges are also linked, but this link can work in both directions. he shortening rates for the pro-and retro-wedge are complementary, thus if the shortening rate of one wedge increases, the shortening rate of the opposing wedge decreases by an equal amount (Figure 4.6c). In our models, the most common cause for this is creation of a new pro-wedge thrust. Every new pro-wedge thrust causes a temporary increase in the pro-wedge shortening rate and an equal decrease in the retro-wedge shortening rate. A similar relationship has been observed in analogue models of bivergent wedges [Hoth et al., 2007]. In addition to new pro-wedge thrusts, new retro-wedge thrusts can have the same efect. Phrased more generally, any event or condition that afects the shortening rate of one wedge has an immediate, equal and opposite efect on the shortening rate of the opposing wedge (assuming shortening rate is measured as a percentage of the total convergence rate). his relationship is a natural consequence of the force balance that must always exist within the orogen as a whole.

Natural systems

Our models consistently reproduce a two-phase evolution when rit inheritance is involved. his allows us to propose a generic evolutionary model for inverted rit systems: Phase 1 is the roughly symmetric inversion of the rit. Phase 2 is the asymmetric main collision phase, where continental subduction forces ~80% of total convergence to be accommodated in the pro-wedge. Here we compare several natural systems to this proposed model.

High Atlas

he High Atlas is an ENE-trending intracontinental mountain belt in North Africa. he Middle Atlas branches of the High Atlas in a NE direction. Here we will discuss a cross section through the Central High Atlas to the east of this branch intersection (Fig. 4.8a) [Arboleya et al., 2004;[START_REF] Ayarza | Crustal structure under the central High Atlas Mountains (Morocco) from geological and gravity data[END_REF]. he High Atlas evolved from a rit that opened during the Mesozoic and was inverted during several Cretaceous and Cenozoic compressional phases [e.g., [START_REF] Mattauer | Sur les mecanismes de formation des chaines intracontinentales; l' exemple des chaines atlasiques du Maroc[END_REF][START_REF] Jacobshagen | Structure and post-Palaeozoic evolution of the central High Atlas[END_REF][START_REF] Beauchamp | Intracontinental riting and inversion: Missour basin and Atlas mountains, Morocco[END_REF][START_REF] Frizon De Lamotte | he two main steps of the Atlas building and geodynamics of the western Mediterranean[END_REF][START_REF] Teixell | Tectonic shortening and topography in the central High Atlas (Morocco)[END_REF]. Overall shortening is very low, only ~26 km, accommodated on several partially inverted normal faults that dip towards the centre of the orogen [START_REF] Teixell | Tectonic shortening and topography in the central High Atlas (Morocco)[END_REF]. Using a boundary to separate north-verging from south-verging structures, shortening is partitioned equally between the north and the south (~13 km each). Given that the High Atlas is an inverted rit system, and shortening is very low and partitioned equally, we propose that the High Atlas is a natural analogue for phase 1 of our evolutionary model. Distributions of shortening were obtained with a basic area balancing that assumes a constant-thickness crust before orogenesis. (a) Central Pyrenees ECORS Pyrenees section, ater Muñoz [1992] and Beaumont et al. [2000]. (b) High and Middle Atlas, ater Arboleya et al. [2004] and [START_REF] Ayarza | Crustal structure under the central High Atlas Mountains (Morocco) from geological and gravity data[END_REF]. (c) Western Alps ECORS CROP section, ater Schmid and Kissling [2000]. (d) Central Alps NFP-20 East section, ater Schmid et al. [1996]. 4.4.3.2 Pyrenees he Pyrenees is a WNW-ESE trending mountain belt that formed due to the Alpine collision of the Iberian microplate and the European continent. his Cenomanian to Oligocene collision inverted a hyperextended rit system that exhumed mantle between the two plates [e.g., Jammes et al., 2009;Lagabrielle et al., 2010]. he present-day structure of the Pyrenees consists of a narrow northern retro-wedge, a south verging antiformal stack of Iberian basement thrust sheets, and a zone of south Pyrenean cover thrust sheets (Fig. 4.8b). Estimates of shortening for the Central Pyrenees vary, from ~90 km to ~165 km [Beaumont et al., 2000;Mouthereau et al., 2014] Of this higher estimate, ~128 km is accommodated in the pro-wedge and ~37 km in the retro-wedge. Estimates for the Eastern Pyrenees are lower, but show a similar strain partitioning: ~111 to ~125 km overall, with the lower estimate partitioned as ~92 km in the pro-wedge and ~19 km in the retro-wedge [Vergés et al., 1995[Vergés et al., , 2002;;Grool et al., 2018]. A recent reconstruction of the pro-retro strain partitioning for the Eastern Pyrenees revealed a roughly symmetrical rit inversion phase, followed by an asymmetric main collision phase [Grool et al., 2018]. hese observations correspond very well to a transition from phase 1 to phase 2 of our evolutionary model. 4.4.3.3 Alps he Alps formed due to southward oceanic subduction of the Tethys Ocean followed by collision of the Apulian (south) and European (north) plates. Total convergence is estimated to be ~600 km over several phases [Schmid et al., 1996;2008;[START_REF] Handy | Reconciling plate-tectonic reconstructions of Alpine Tethys with the geological-geophysical record of spreading and subduction in the Alps[END_REF]. Here we compare only the post-suture or collisional phase (since ~35 Ma) in the Western and Central Alps. For a compilation of shortening estimates in the Alps, see Figure 4.9.

In terms of overall structure, the Western Alps show a wide thick-skinned pro-wedge and a narrow thick-skinned retro-wedge (Fig. 4.7c) [Schmid and Kissling, 2000]. he Ivrea mantle has reached a very shallow position between the upper and lower plate. he Jura mountains form a distal thin-skinned proforeland fold-and-thrust belt detached along an evaporitic décollement. Shortening ater 35 Ma is ~124 km, partitioned as ~102 km in the pro-wedge and ~22 km in the retro-wedge. hese numbers appear to it with the strain partitioning predicted by our model. However, alternative shortening estimates result in a less compatible strain partitioning: One such alternative north of the Penninic front and to the north of this transect is 76-96 km [START_REF] Burkhard | Evolution of the western Swiss Molasse basin: structural relations with the Alps and the Jura belt[END_REF]], but it is unclear how much of this shortening is the result of pre-35 Ma convergence. Even further northeast along the Alpine chain, Rosenberg and Kissling [2013] ind signiicantly lower 'post-nappe' shortening of 71 km and 17 km north and south of the Insubric line.

he Central Alps comprise a large stack of basement-involved thrust sheets north of the Insubric line that largely predates the collision phase (Figure 4.8d) [Schmid et al., 1996] he southern half of this stack rests on top of the lower crust of the Apulian plate. his corresponds to the hinterland of the pro-wedge that migrated onto the upper plate in our models. he Apulian upper crust forms a relatively narrow thick-skinned stack with thin-skinned thrusts of the Milan thrust system linking back into it. he strain partitioning does not correspond to those predicted by any of our models. Shortening since 40 Ma is partitioned as ~108 km and ~56 km north and south of the Insubric line, respectively [Schmid et al., 1996]. When accounting only for shortening since 32 Ma, shortening is partitioned as ~63 km and ~56 km north and south of the Insubric line [Schmid et al., 1996]. An alternative estimate of 'post nappe' shortening is ~41 km and ~56 km, respectively [Rosenberg and Kissling, 2013]. None of these estimates are close to the 80/20 split predicted by our models. In fact, it appears that more recent shortening was distributed equally. Based on these estimates, we conclude that our evolutionary model for inverted rit systems cannot be applied to the Alps. he high convergence, multiple orogenic phases, and pre-existing asymmetry due to oceanic subduction inluence the pro-retro strain partitioning to such a degree that a diferent evolutionary model is required to explain the strain partitioning of the Alps. Other, higher shortening estimates for the area around this section exist [Roeder, 1992;[START_REF] Schönborn | Kinematics of a transverse zone in the Southern Alps, Italy, in hrust Tectonics[END_REF]Pifner et al., 2000], but either the timing of these estimates is unclear or the estimates only cover one half of the section (Fig. 4.9). 4°F igure 4.9. Summary of available shortening estimates in the Western and Central Alps [START_REF] Laubscher | he problem of the deep structure of the Southern Alps: 3-D material balance considerations and regional consequences[END_REF]Roeder, 1992;[START_REF] Schönborn | Kinematics of a transverse zone in the Southern Alps, Italy, in hrust Tectonics[END_REF]Schmid et al., 1996;[START_REF] Burkhard | Evolution of the western Swiss Molasse basin: structural relations with the Alps and the Jura belt[END_REF]Schmid and Kissling, 2000;Rosenberg and Kissling, 2013]. 4.4.4 Limitations he models presented in this paper are simpliied systems that have some inherent limitations. Firstly, all models with extensional inheritance have a very simple rit structure. Natural rit systems generally have more complexity such as asymmetry or a higher number of normal faults that may inluence the orogenic structure when inverted. Secondly, our models do not feature full crustal breakup during the riting phase. However, earlier tests have shown that this has no noticeable efect on the orogenic structure [Erdős et al., 2014]. hirdly, our models feature no precursor subduction. As discussed above, this limits the applicability to orogens that did have oceanic subduction before collision. Fourthly, the décollement distribution and crustal strength are completely uniform in our models. his is a necessary simpliication to eliminate some variables and emphasise the fundamental behaviour and mechanisms. However, it prevents the models from exactly duplicating the structures and evolution of particular natural systems. hese models are a tool that helps us understand how natural systems work.

Due to these limitations in the models, comparing them to natural systems also brings some limitations. For the Alps, we were only able to compare post-suture shortening. However, during oceanic subduction, the geometry of the Alpine domain already evolved to a geometry that is signiicantly diferent from our models [e.g., Schmid et al., 1996]. In addition, oblique convergence may limit the amount of reactivation of inherited structures that are not optimally oriented. Our 2D models cannot account for this.

Conclusions

We have used high resolution numerical models to investigate the inluence of several factors, décollement rheology in particular, on the general crustal structure and strain partitioning between the pro-and retrowedge. From our results we draw the following conclusions:

• he evolutionary pattern of an inverted rit system predicted by our models can be applied to natural orogens that also started as inverted rit systems. he High Atlas represents phase 1, roughly symmetrical inversion. he Pyrenees represent phase 2, asymmetric collision.

• he rheology of a décollement layer in the sedimentary cover inluences the crustal structure of the orogen, and may also inluence the strain distribution between the pro-and retro-wedge.

• he distribution of a décollement layer inluences the plate polarity and promotes shortening of the side where the décollement is situated (or thicker). A décollement that is far away from the area of initial deformation does probably not inluence this, as such a décollement would remain uninvolved until ater the plate polarity is established.

• he absence of a décollement altogether in the retro-wedge promotes early thick-skinned propagation of deformation into the upper plate.

• Sedimentation and erosion tend to negate the efects of the décollement rheology on the pro-retro strain partitioning, but not the efects on the crustal structure (width of the thick-skinned orogen).

• he pro-and retro-wedge are inextricably linked. Events and conditions on one side have an immediate efect on the other side, as a result from the force balance that must exist in the system.

In Chapter 1 of this thesis, a number of questions related to the dynamic relationship of pro-and retrowedges were outlined. In this chapter, the new insights gained from this work are summarised, and an attempt is made to answer these questions.

How does a doubly vergent orogen evolve?

Detailed analysis of the Eastern Pyrenees has revealed that shortening was distributed roughly equally across the European and Iberian margins during early inversion of the rit. From the end of the Maastrichtian onwards, shortening was distributed asymmetrically in the Eastern Pyrenees. Around 80% of shortening was accommodated in the southern Pyrenees and the northern Pyrenees were temporarily abandoned during the Paleocene. From the hanetian onwards, the retro-wedge was reactivated, but the pro-wedge remained dominant. Because the change in shortening distribution occurred coeval with the estimated onset of main collision and subduction of the Iberian lower crust, it appears that this change is the result of an internal threshold (i.e. collision) being reached that is intrinsic to inverted rit systems. his change may thus have occcurred independent from external forcing (e.g., plate kinematics).

his evolution from symmetric inversion to asymmetric collision is reproduced by our numerical models that include extensional inheritance. Quantitatively, the models also agree with the evolution as observed in the Eastern Pyrenees: early shortening is distributed 50/50, and shortening during main collision is distributed roughly 80/20 between both plates. A temporary abandonment of the retro-wedge is also observed in our models, but the amount of shortening accommodated by the lower plate during the upper plate's quiescence is much higher in the models than reconstructed for the Eastern Pyrenees.

Do the pro-and retro-wedge interact?

Based on the data from the Eastern Pyrenees, the pro-and retro-wedge do interact. Firstly, the change in shortening distribution between the pro-and retro-wedge that is observed in the Pyrenees can be regarded as interaction between both wedges. Secondly, the reactivation of the retro-wedge occurred around the same time as the onset of exhumation in the Axial Zone is recorded. According to the reconstruction in Chapter 2, shortening and subsidence of the northern Pyrenees during the Eocene was mainly driven by backthrusting of the Axial Zone onto the European plate. he pro-wedge thus drives shortening in the retro-wedge during this phase.

he numerical models also reproduce this. However, an even more direct interaction between the pro-and retro-wedge can be seen in the numerical models: he pro-and retro-wedge do not grow at a constant rate, even at geological timescales. he shortening rate of the retro-wedge changes as soon as a new crustal thrust is formed in the pro-wedge. Unfortunately, interactions at these timescales are unlikely to be preserved in a natural system. However, theoreticaly, such interactions should indeed exist, because of the force balance that must exist in the system as a whole. Any event in one wedge thus has an immediate efect on the other wedge.

What factors inluence the shortening distribution, and how?

In this work, extensional inheritance is the only factor that can dramatically alter the shortening distribution. he numerical models show that, of the investigated factors, only extensional inheritance has enough inluence to cause an orogen to become singly or doubly vergent. Surface processes and thinskinned deformation can signiicantly alter the crustal structure, but the shortening distribution between the pro-and retro-wedge remains largely unafected.

Secondary questions

What causes the Paleocene quiescence in the Pyrenees remains debatable. he hypothesis developed on the basis of the East Pyrenean study states that such a quiescent period is part of the transition from equal to asymmetric shortening, and is to be expected for all inverted rit systems. he viability of this mechanism was conirmed by the numerical models, which reproduced a similar temporary abandonment of the retrowedge. A key implication of this mechanism is that shortening of the lower plate continues while the retrowedge is inactive. his is in direct contrast to explanations that implicate external factors (plate kinematics, intraplate shortening) as the cause for the Pyrenean Paleocene quiescence, which require shortening to slow down or stop in both wedges simultaneously. Unfortunately, data on Paleocene shortening of the southern Pyrenees is sparse and oten ambiguous, and plate kinematic reconstructions already predict a Paleocene phase of slow convergence. Both mechanisms thus remain valid for the time being. A hybrid solution could also have been the case: through sheer (and literal) coincidence, the transition from symmetric to asymmetric shortening may have occurred around the same time as a slowing in convergence.

hin-skinned deformation can inluence orogenic crustal sructure. he numerical models described in Chapters 3 and 4 show that thin-skinned deformation can inluence the crustal structure. When combined with syn-tectonic sedimentation and erosion, eicient decoupling of the sedimentary cover promotes formation of a crustal antiformal stack. he sedimentary décollement creates a system of two wedges that share a common top surface. he mid-crustal décollement can support a higher taper than the sedimentary décollement. his results in the cover wedge constantly collapsing to keep taper low, while the crustal wedge constantly deforms internally to increase taper. he high internal deformation of the crust is what creates the antiformal stack.

Perspectives

he timing of the change in shortening distribution coeval with the onset of subduction of the lower plate, both in the Pyrenees and numerical models, suggests that this transition is intrinsic to inverted rit systems, and not driven by external factors. hat implies that a similar change should be observed in other inverted rit systems. his hypothesis should be tested further by performing a similar analysis in other natural systems. Perhaps a broader model can be developed, that can also be applied to orogens that inherited an asymmetry from oceanic subduction prior to collision.

he method used in Chapter 2 to reconstruct a high resolution shortening rate history should be tested for sensitivities. On the whole it appears that the reconstruction for the Eastern Pyrenees is fairly representative for the whole orogen. It would be interesting to apply the same method to other cross sections further west to test this. If the method returns representative results along strike, it will allow more detailed results for the same amount of work as the commonly used method of stepwise restoration.

Should the shortening reconstruction method prove robust enough, it could be used to prove gravitational sliding in convergent settings. To do this, the rate of horizontal displacement of the cover and basement should be compared. Cover thrust sheets move only if they are pushed from behind, or if the subhorizontal component of the normal force vector is large enough on its own (assuming an inclined surface, of course). hus, in a standard setting with no gravitational sliding, a cover thrust sheet can never outpace the basement. However, if the cover does move faster than basement, only gravity can be responsible. hat is the theory, at least. he practical reality of constraining shortening rates of basement thrust sheets will, no doubt, make this very diicult. In this work, movement of the cover thrust sheets was used to constrain basement thrusting, which precludes an analysis as described above.

he dual approach of detailed study of a natural system in combination with testing factors and mechanisms with numerical models has revealed the crustal-scale dynamics of a complex orogenic system in more detail than before. he beneit of this dual approach goes further than simply testing ield hypotheses, a task models are not always suited for. Working with numerical models allows one to visualise deformation, see processes occur on a human timescale that makes them much more relatable and easier to understand. hese insights can then be applied to the interpretation of geological data. Vice versa, knowledge of natural systems and their characteristics informs the models. A similar approach of close collaboration between ield studies and modelling should thus continue, also in other tectonic settings. 2.8b, c in Chapter 2. he method used to obtain this data is described in Chapter 2. As a result of assuming a constant average rate for each structure, rather than gradually increasing and decreasing, some artefacts may be present in the data where two structures briely overlap in time. he spike around 47 Ma is probably such an artefact.

Contents

Table A.1 shows a detailed correlation of stratigraphic units used in the paper and the BRGM 1:50,000 geological maps.

Tables A.2 to A.4 present values used for, and results of the subsidence analysis of the northern and southern foreland basins. he methodology of the subsidence analysis is explained in Chapter 2.

Table A.5 presents all ield data collected and used during this study. he measurements are divided into three categories, each shown in a separate tab: plane measurements (bedding, cleavage, etc.), fault measurements (normal, reverse, etc.) and lineation measurements (fold axis, striae, etc.). mantle. At the center of the model, a weak seed of 6x6 km, centered around 28 km depth has been fully strain weakened. he side boundaries are thermally insulated, while the bottom and top boundaries are kept at 1330° C and 0° C, respectively. he initial temperature at the Moho is 550° C. he top boundary of the model is a true free surface, while the bottom allows free slip. On the let and right boundaries, we introduce material at 5 cm/y in the upper 132 km (10 cm/y in total). Below 172 km, material is extracted at ~0.15 cm/y to balance the volume of introduced material. Between 132 and 172 km, the velocity is interpolated linearly to create a smooth transition. During the extensional phase, where applicable, the velocity boundary conditions are inverted.

Eastern Pyrenees Central Pyrenees (ECORS) Western

Model parameters he material properties used in the models are listed in Table B.1.

Shortening distribution measurement method he distribution of shortening between the pro-and retro-wedge is quantiied by tracking the horizontal position of a column of points that originates at the centre of the model. his column will be displaced towards the side that accommodates the most deformation. he amount of displacement is a measure for the ratio of shortening accommodated by both sides. his method does not distinguish between the proand retro-wedge as originally deined by Willett et al. [1993], where the boundary between the pro-and retro-wedge is a topographic drainage divide. Instead we track the tectonic boundary between pro-and retro-verging structures, which starts aligned with the plate boundary, or point S in Willett et al. [1993]. DR2: no inheritance + shale + surface processes Initial convergence is accommodated by distributed deformation until a single shear zone forms ater 6 My. his shear zone reaches down into the lithospheric mantle, creating an asymmetry that determines the plate polarity (Figure B.3a). No strain-weakened retro-shear zone exists at this point. In the footwall of this thrust, a small thin-skinned wedge develops (Figure B.3b). 33 km of convergence is accommodated along the single shear zone before a thrust sheet is accreted into the pro-wedge. Both shear zones remain active, although the frontal shear zone is dominant. Meanwhile, the pro-foredeep is illed with sediment and the sedimentary cover of the uplited upper plate begins to erode. As new thrust sheets are stacked in the prowedge, the upper, older thrust sheets migrate very slowly onto the upper plate along a new retro-shear zone (Figure B.3c). his process continues until 215 km of convergence, where the oldest and uppermost thrust sheets have been deeply eroded (Figure B.3d). he resulting orogen is highly asymmetrical, with a wide pro-wedge and only a single retro-wedge shear zone that did not accommodate much shortening.

Figure 1 . 1 .

 11 Figure 1.1. Example of a natural fold-and-thrust belt taking on a wedge shape. Ater Suppe [1980].

Figure 1 .

 1 Figure 1.2. a) Cross-section through a critically tapered wedge. b) stability ield of the wedge. Ater Buiter [2012].

Figure 1 . 3 .

 13 Figure 1.3. he delection of an elastic plate under an applied load, forming a lexural foreland basin. he diferent lines show how the delection depends on the lexural rigidity. Ater Beaumont [1981].

Figure 1 . 4 .

 14 Figure 1.4. S-point model resulting in a doubly vergent geometry with a pro-and retro-wedge. Ater Willett et al. [1993].

Figure 1 . 5 .

 15 Figure 1.5. Pro-and retro-foreland basins record diferent subsidence histories. Ater Naylor and Sinclair [2008].

Figure 1 . 6 .

 16 Figure 1.6. A balanced and restored cross section through the Zagros mountains. Ater McQuarrie [2004].

Figure 1 .

 1 Figure 1.7. a) he generic porosity-depth relationship. b) Schematic of decompaction. Ater Allen and Allen [2005].

Figure 1 . 8 .

 18 Figure 1.8. Total versus tectonic subsidence. Ater Steckler and Watts [1978].

Figure 1 .

 1 Figure 1.9. a) Example of a inite element model. Ater Jammes and Huismans [2012]. b) Example of a discrete element model. Ater Naylor et al. [2005].

  1.3.1 Present-day structure he Pyrenees form a WNW-ESE striking orogen at the boundary between Spain and France that was created by Late Cretaceous to Oligocene collision of the Iberian and European plates. he orogen can be divided into the following tectonostratigraphic zones from north to south (Figure 1.10):

Figure 1 .

 1 Figure 1.10. Crustal structure of the Central Pyrenees based on the ECORS Pyrenees deep seismic section. Ater Muñoz [1992].

Figure 1 .

 1 Figure 1.11. Location of lherzolite outcrops in the Pyrenees. Ater Lagabrielle et al. [2010].

Figure 1 .

 1 Figure 1.12. hree competing plate kinematic reconstructions of the Pyrenean domain. Ater Mouthereau et al. [2014].

Figure 1 .

 1 Figure 1.13. Schematic cross section of the Pyrenean rit that exhumes mantle. Ater Lagabrielle et al. [2010].

Figure 1 .

 1 Figure 1.14. Lateral stratigraphic section through the northeastern Pyrenees showing the westward prograding geometry of Upper Cretaceous Flysch. Ater Ricateau and Villemin [1973] and Bilotte [1985].

Figure 1 . 15 .

 115 Figure 1.15. Estimates of shortening rate through time across the Pyrenees.

  Fig.6a

Fig. 6b

  Fig.6b

  Fig. 6c

2. 4 . 1

 41 Stratigraphy of the North Pyrenean ForelandAcross the Saint-Barthélémy -Fougax study area a Mesozoic to Cenozoic succession overlies Variscan basement, that comprises Paleozoic metasediments and crystalline rocks (Figure2.3b). Triassic to lower Cenomanian marine strata record rit-related subsidence. Upper Cenomanian to Santonian strata record post-rit subsidence. On a regional scale the syn-orogenic succession (Campanian to Oligocene) records a consistent along-strike transition from continental to marine facies to the west with facies interingering in the study area. No strata older than Cenomanian overlie Variscan basement north of the Benaix Fault implying that the foreland was largely unafected by Mesozoic riting. No strata younger than Campanian are preserved south of the thrust front.

FluvialFigure 2 . 3 .

 23 Figure 2.3. Chronostratigraphic charts for the north and south Pyrenean foreland basins. (a) South Pyrenean stratigraphic chart modiied ater Vergés [1993] and Vergés et al. [2002]. (b) North Pyrenean stratigraphic chart. Boreholes Dr3, Dr4 and Bx1 are located on Figure 2.2. Note double horizontal scaling compared to (a). For detailed north Pyrenean formation names and BRGM unit codes see Table A.1. (c) Timing of main tectonic events. Phases 1 to 4 represent the Pyrenean orogeny.

  Fig.4c

Figure 2 . 4 .

 24 Fig.4d

  Fig.6

Figure 2

 2 Figure 2.8.(a, c) Reconstruction of the distribution of shortening through time in the southern Pyrenees (J3 section; Fig.2.9) fromVergés et al. [1993] and northern Pyrenees (S3 section; Figures 2.4 and 2.5). Each line represents either the horizontal position of the leading edge of a principal fault or the position of the hinge zone at the surface of a major anticline. Displacement and fault length are assumed to increase linearly for each fault. For emergent faults, we assumed the fault reached maximum length during deposition of youngest cut strata. Fault timings for the southern Pyrenees fromfault-strata relationships[START_REF] Vergés | Corte compensado del Pirineo oriental: Geometría de las cuencas de antepaís y edades de emplazamiento de los mantos de corrimiento[END_REF] Vergés et al., 1995 Vergés et al., , 1998 Vergés et al., , 2002;; Ford et al., 1997; Ramos et al., 2002;[START_REF] Costa | Closing and continentalization of the South Pyrenean foreland basin (NE Spain): magnetochronological constraints[END_REF][START_REF] Carrigan | Fault-related fold kinematics recorded by terrestrial growth strata, Sant Llorenç de Morunys, Pyrenees Mountains, NE Spain[END_REF]. (b, d) Horizontal position of the deformation front, the main depocentre(s) and foreland pinchout of the basin as a function of time in the southern and northern Pyrenees, respectively. he deformation front tracks the location of the most frontal structure through time. Depocentres were measured from balanced cross sections. Pinchout locations were determined from borehole and seismic data. Shortening rates through time are shown by grey bars, derived from the fault activity in (a, c). Shortening rates for the south include internal shortening in the Axial Zone (AZ). Dashed bars have a high uncertainty for the early convergence.

Figure 2

 2 Figure 2.9. Balanced and restored cross section J3 in the south Pyrenean pro-foreland. Simpliied aterVergés [1993]. Boreholes and stratigraphic sections used for subsidence calculations are indicated as semitransparent white boxes with blue letters. M, Montserrat; C, Castellfollit; S, Santpedor; P-r, Puig-reig; J, Jabalí; G, Gombrèn (split in two); Pf, Pedraforca. Formation names are indicated by numbers in circles as in Figure2.3a.

  2.5.3 Subsidence of the South Pyrenean Foreland he four boreholes and two stratigraphic sections along section J3, analysed byVergés et al. [1998], record the main Eocene basin history(Figures 2.7a, b, 2.9; Table A.4). Of those, only the Gombrèn stratigraphic section records Paleocene subsidence. To represent the Late Cretaceous and Paleocene subsidence record for the northernmost part of the foreland basin, a new subsidence curve was constructed for the Lower Pedraforca thrust sheet, based on data from[START_REF] Vergés | Mapa Geológico de España[END_REF]. Due to uncertain dating of strata, Maastrichtian and Paleocene subsidence are shown as maximum and minimum rates, respectively(Figure 2.7a, b). If these are assumed to represent true tectonic subsidence, the Lower Pedraforca thrust sheet appears to record very slow subsidence during the Paleocene (<0.01 mm/y; Figure2.7b; Table

  .7a, b) [Gómez-Paccard et al., 2012].

Figure 2 . 10 .

 210 Figure 2.10. Balanced stepwise evolution of the Eastern Pyrenees. Shortening amounts are cumulative. Red lines indicate faults that have been active since the preceding step. hick black lines indicate abandoned faults. Dashed lines indicate faults that may have been inherited but have not yet been reactivated. (a) Moho and subduction contact are those along the ECORS line ~50 km to the west, ater Chevrot et al. [2015], projected to align along the North Pyrenean Fault. R, Rialp; O, Orri; N, Nogueres; SB, Saint-Barthélémy Massif; NPF, North Pyrenean Fault; 3MF, 3M Fault; NPFT, North Pyrenean Frontal hrust.

Model 1 Figure 3 . 1 .

 131 Figure 3.1. Evolution of model 1, with extensional inheritance and a salt décollement level. Δx c is total convergence. Plot windows were chosen so the upper plate appears stationary. he white line in the keystone block marks the column of points tracked for measuring the shortening distribution. Points on this line that are above topography (red) are ignored. Strain rate plots (let) show second invariant of the strain rate tensor. Grey dashed lines in strain rate plots indicate fully strain weakened areas. (a) rit phase. (b) Symmetric inversion. (c) Onset of collision and asymmetry, retro-wedge abandoned. (d) Retro-wedge reactivated and wide thin-skinned pro-foreland fold-and-thrust belt. (e) Final coniguration.Well-developed retro-wedge, narrow thick-skinned pro-wedge and wide thin-skinned pro-foreland foldand-thrust belt.

Figure 3 . 2 .

 32 Figure 3.2. (a) Distribution of shortening between the retro-wedge and pro-wedge, plotted against total convergence to eliminate efects of variable convergence rates. Central Pyrenean data (black box) from Beaumont et al. [2000]. East Pyrenean data (red line) from Grool et al. [2018]. (b) Distribution of shortening rates, plotted against total convergence. East Pyrenean data from Grool et al. [2018].

Figure 3 . 3 .

 33 Figure 3.3. (a) Cartoon of Model 1 ater 225 km of convergence, showing the combined efects of rit inheritance, an eicient décollement layer and surface processes. (b) cartoon of Central Pyrenees, modiied ater Muñoz [1992] and Beaumont et al. [2000].

  b

Figure 4 . 1 .

 41 Figure 4.1. (a) Initial geometry and boundary conditions. (b) Implementation of strain sotening of frictional materials. (c) Initial strength proile and rheological layering. Ater Erdős et al. [2014].

a

  Shale ϕ ef = 2°, Salt η ef = 10 19 Pa•s, Weak salt η ef = 10 18 Pa•s. b base level = 0 m, start at 10 My. c E f = 4.35•10 -15 s -1 . d Erdős et al., [2014]. e Erdős et al., [2015].

4. 3 . 1 Figure 4 . 2 .

 3142 Figure 4.2. Model results of set 1 ater 36 My and 300 km of convergence. Convergence (Δx C ) was measured starting from formation of the irst shear zone in M1, or starting from restored crustal thickness in models with extensional inheritance. he white line indicates the tracked points that measure the distribution of shortening. Points above topography are ignored and coloured red.

Figure 4 . 3 .

 43 Figure 4.3. Model results of M5, which has an asymmetric shale décollement layer. (a) and (b) are zooms of (c). Convergence (Δx C ) was measured from restored crustal thickness onwards. he white line indicates the tracked points that measure the distribution of shortening.

Figure 4 . 4 .

 44 Figure 4.4. Model results of M6, which has a viscous salt décollement. (b) and (c) are zooms of (d). Convergence (Δx C ) was measured from restored crustal thickness onwards. he white line indicates the tracked points that measure the distribution of shortening.

Figure 4 . 5 .

 45 Figure 4.5. Model results of M7 and M8, which are part of set 2 with a viscous salt décollement. (a) and (b) are zooms of (c). (e) and (f) are zooms of (g). Convergence (Δx C ) was measured from restored crustal thickness onwards for both models. he white line indicates the tracked points that measure the distribution of shortening. Points above topography are ignored and coloured red.

  Figure 4.6.(a) plots of the distribution of shortening between the lower (pro) and upper (retro) plate for models with no surface processes. Strain distributions for the High Atlas, Pyrenees, Western Alps and Central Alps from[START_REF] Teixell | Tectonic shortening and topography in the central High Atlas (Morocco)[END_REF] Beaumont et al., 2000; Schmid and Kissling, 2000; Schmid et al., 1996]. (b) plots of the distribution of shortening between the lower and upper plate for models with surface processes, and M1 as reference. (c) Distribution of shortening rates as percentage of the total convergence rate between the lower and upper plate. his plot shows the distribution for M4 as an example. All models feature a similar evolution of the distribution through time.

Figure 4 . 7 .

 47 Figure 4.7. Summary of model behaviour as a result of the tested factors.

Figure 4 . 8 .

 48 Figure 4.8. Summary of natural systems. he lower plate is shown in green and the upper plate in blue. Distributions of shortening were obtained with a basic area balancing that assumes a constant-thickness crust before orogenesis. (a) Central Pyrenees ECORS Pyrenees section, ater Muñoz [1992] and Beaumont et al. [2000]. (b) High and Middle Atlas, ater Arboleya et al. [2004] and Ayarza et al. [2005]. (c) Western Alps ECORS CROP section, ater Schmid and Kissling [2000]. (d) Central Alps NFP-20 East section, ater Schmid et al. [1996].

Figure A. 1 .

 1 Figure A.1. A comparison of shortening rates through time from diferent authors and diferent parts of the Pyrenees.

Figure A. 1 .

 1 Figure A.1. Shortening rates for the Western, Central and Eastern Pyrenees. West Pyrenean shortening rate from Teixell et al. [2016]. Central Pyrenean shortening rates from Beaumont et al. [2000], and Mouthereau et al. [2014]. East Pyrenean shortening rate from this work (Chapter 2).

Figure B. 1 .

 1 Figure B.1. Model set up. (a) Initial geometry of the models, including velocity boundary conditions and initial thermal proile. (b) Implementation of strain sotening in frictional materials. (c) Initial strength proile, showing the strong coupling of the lower crust and lithospheric mantle. Ater Erdős et al. [2014].

  Figure B.2. Evolution of model DR1, with extensional inheritance and a shale décollement level. Δx c is total convergence. Plot windows were chosen so the upper plate appears stationary. he white line in the keystone block marks the column of points tracked for measuring the shortening distribution. Points on this line that are above topography (red) are ignored. Strain rate plots (let) show second invariant of the strain rate tensor. Grey dashed lines in strain rate plots indicate fully strain weakened areas. (a) Symmetric inversion. (b) Onset of collision and asymmetry, retro-wedge abandoned. (c) Retro-wedge reactivated and narrow thin-skinned pro-foreland fold-and-thrust belt. (d) Final coniguration. Well-developed retrowedge, wide thick-skinned pro-wedge directly overlain by thin-skinned pro-foreland fold-and-thrust belt.

Figure B. 3 .

 3 Figure B.3. Evolution of model DR2, with no extensional inheritance and a shale décollement level. Δx c is total convergence. Plot windows were chosen so the upper plate appears stationary. he white line in the keystone block marks the column of points tracked for measuring the shortening distribution. Points on this line that are above topography (red) are ignored. Strain rate plots (let) show second invariant of the strain rate tensor. Grey dashed lines in strain rate plots indicate fully strain weakened areas. (a) initial asymmetry, no conjugate shear zone. (b) Continued deformation only in pro-wedge. (c) Stack of thickskinned pro-wedge thrust sheets builds and begins to activate a retro-shear zone. (d) Final coniguration, highly asymmetrical orogen with very little deformation of the single retro-shear zone, a wide thickskinned pro-wedge overlain by a thin-skinned pro-foreland fold-and-thrust belt.

ia n R a n g e Western Pyrenees C at al an C oa st al R an ge s

  . South of the North Pyrenean Frontal hrust is the North Pyrenean

						°2°1°0°1°2°3°4°3
			Santander a	Atlantic Ocean Bay of Biscay	France	Aquitane Basin	Central Massif	4°4
										Toulouse
		Zaragoza Spain Pamplona Jaca Huesca 2°4 Lleida Bilbao Burgos Logroño Vitoria San Sebastian Pau Tarbes Carcassonne Narbonne Perpignan Ripoll Girona Tremp Andorra PYRENEES J3 S3 S3b Lyon Central and Eastern Pyrenees South Central Unit 42°4 Ib e r ECORS Bay of 3°4 3°4 Fig. 2
				Duero Basin				Ebro Basin
										Barcelona
		Late Cretaceous-Paleocene Eocene-Oligocene Eocene-Miocene Neogene-Quaternary 1°3	thrust			Tarragona	Mediterranean Sea València Trough	1°4
				Early Cretaceous		blind thrust		50	100 km
				Basement			normal fault
					°2°1°0°1°2°3°4°0
			b	Tectonic zones:	Ebro Basin	SPZ	Axial Zone	NPZ	Sub-PZ	Aquitaine Basin
	Depth (km)	0 -10 20 10 30	S	Iberian plate			O r r i R ia lp N o g u e r e s	NPF	European plate N NPFT	-10 0 10 20 30	Depth (km)
		40		Sediments					40
		50		Upper Crust Pyrenean basement				0	50 km	50
				Lower Crust
	Figure 2.1. (a) Simpliied geologic map of the Pyrenees. Red lines locate the cross sections (Figs. 2.4, 2.5,
	2.9). Black line indicates the trace of the ECORS Pyrenees deep seismic section shown in (b). Modiied
	ater Vergés et al. [2002] and Angrand et al. [2018]. (b) Interpretation of the ECORS Pyrenees deep seismic
	section, with major structural zones indicated. Modiied ater Muñoz [1992]. SPZ, South Pyrenean Zone;
	NPZ, North Pyrenean Zone; Sub-PZ, Sub-Pyrenean Zone; NPF, North Pyrenean Fault; NPFT, North
	Pyrenean Frontal hrust.

  Roquefeuil block have a wavelength of ~1.5 to ~2.5 km (Figures 2.2, 2.5c). he southward twofold increase in fold wavelength is abrupt, and is interpreted to indicate a change in depth to the basal thrust. Because Mirande Group limestones are clearly absent below the northern Fougax block, we propose that they are limited to the north by a blind inverted extensional basin margin fault (Roquefeuil Fault; Figure 2.5b) sealed by the Black Flysch Group.

he Saint-Barthélémy Massif is a large block of Paleozoic metasediments and crystalline basement

[de Saint Blanquat et al., 1990]

. he top of basement drops eastward to a deeper level across an east-dipping lateral ramp to lie below the Metamorphic Internal Zone

[Souquet and Peybernès, 1987]

and the Roquefeuil block, as evidenced by the small tectonic windows around Camurac

(Figures 2.2, 2.6)

. South of Montségur, a syncline of the Fougax block plunges steeply eastward (Figure

  .4.4.1 Early Cretaceous Rit and Late Cretaceous Post-rit Phases he mildly metamorphic (up to 200° C;

				Southern Pyrenees						Northern Pyrenees
					Time (Ma)							Time (Ma)
		100	90	80	70	60	50	40	30	100	90	80	70	60	50	40	30
		a	Total subsidence						c Total subsidence
			sea level								sea level
		0															0
	Depth (km)	1 2	Montserrat Santpedor													1 2	Depth (km)
		3	Castellfollit Puig-Reig							Dr4						3
			Jabalí								T1
		4	Gombrén Lower Pedraforca equivalent					syn1 MRL1						4
		b	Tectonic subsidence					d	Tectonic subsidence
	Depth (km)	0 1															0 1	Depth (km)
		2															2
			post-rift	Phase 1 Ph.2	Phase 3	Ph.4		post-rift	Phase 1 Ph.2	Phase 3	Ph.4
			Late Cretaceous		Paleoc.	Eocene	Olig.		Late Cretaceous		Paleoc.	Eocene	Olig.
		100	90	80	70	60	50	40	30	100	90	80	70	60	50	40	30
					Time (Ma)							Time (Ma)

Figure 2.7. Total subsidence and tectonic subsidence curves for the southern (a, b) and northern (b, c) Pyrenees. All depths are relative to present-day sea level. (a, b) Due to uncertain dating of strata, the solid line of the Lower Pedraforca equivalent shows Maastrichtian and Paleocene subsidence as maximum and minimum rates, respectively. he dashed line shows an average subsidence rate. Other south Pyrenean subsidence data from Vergés et al. [1998]. Detailed supporting data are available as Tables A.2, A.3, and A.4 in supplementary information.

2

  Table 2.1). herefore, while we cannot rule out a Paleocene quiescence in the southern Pyrenees, sparse data appear to indicate slow shortening during the Paleocene. he irst and second orogenic phases were characterised by inversion of rit structures. Although poorly constrained, we estimate ~10 km of shortening were accommodated between ~84 and ~56 Ma, yielding a rate of 0.4 mm/y during the Late Cretaceous and Paleocene (Figure 2.8b; Table 2.1). 2.5.4.2 hird Phase: 59 -34 Ma, Major Convergence During early Eocene, shortening in the South Pyrenean foreland reached a maximum rate of 4.0 mm/y (Figure 2.8b). South of the rapidly advancing thrust front this was accompanied by an increase in tectonic subsidence (0.53 mm/y maximum, Gombrèn; Figure 2.7b; Table A.4), resulting in a deep marine basin (Figure

Table 2 . 1 .

 21 Tectonic subsidence rates and shortening rates in the Eastern Pyrenees a

		Tectonic subsidence rate (mm/y)	Shortening rate (mm/y)
	Orogenic phase	South b	North c	South b	North c
	84-66 Ma	0.06	0.09	0.4	0.6
	66-59 Ma	0.05	0.00	0.4	0
	59-34 Ma	0.12	0.03	2.7	0.3
	34-28 Ma	-	0.01	2.2	0

a All rates were averaged over their respective phases. b Combines Axial Zone, South Pyrenean Zone, and Ebro foreland basin. c Combines North Pyrenean Zone, Sub-Pyrenean Zone, and Aquitaine foreland basin. Chapter 2. Case study of the Eastern Pyrenees 52

Table 4 . 1 .

 41 Mechanical and thermal material properties.

	Parameter	sediment	shale	salt	upper crust lower crust lithospheric	astheno-
	(units)						mantle	sphere
	Flow law	wet qtz • 1	wet qtz • 1 wet qtz • 1 wet qtz • 1 wet qtz • 100 dry olivine wet olivine
	Type	visco-plastic visco-plastic viscous visco-plastic visco-plastic visco-plastic visco-plastic
	Plastic rheology
	c (MPa) a	20 → 4	2	-	20 → 4	20 → 4	20 → 4	20 → 4
	ϕ a	15° → 2°2° → 1°-	15° → 2°15° → 2°15° → 2°15° → 2°V
	iscous rheology
	n	4	4	4	4	4	3.5	3.0
	η (Pa•s)	power-law	power-law	10 19	power-law	power-law	power-law	power-law
	A (Pa -n •s -1 ) b	8.574 • 10 -28 8.574 • 10 -28	-	8.574 • 10 -28 8.574 • 10 -36 2.4168 • 10 -15 1.393 • 10 -14
	Q (kJ mol -1 ) 222.815 • 10 3 222.815 • 10 3	-	222.815 • 10 3 222.815 • 10 3 540.41 • 10 3 429.83 • 10 3
	V (cm 3 mol -1 )	0	0	-	0	0	25 • 10 -6	15 • 10 -6
	Density & thermal parameters
	ρ (kg m -3 )	2300	2300	2300	2800	2800	3360	3300
	α (K -1 )	3.1 • 10 -5	3.1 • 10 -5	3.1 • 10 -5	3.1 • 10 -5	3.1 • 10 -5	0	0
	T 0 (°C)	273.15	273.15	273.15	273.15	273.15	273.15	273.15
	c p (J kg -1 K -1 )	803.57	803.57	803.57	803.57	803.57	681.82	681.82
	k (W m -1 )	2.25	2.25	2.25	2.25	2.25	2.25	2.25
	H (µW m -3 )	0.8 • 10 -6	0.8 • 10 -6	0.8 • 10 -6	0.8 • 10 -6	0.8 • 10 -6	0	0

a he arrow indicates the efect of strain weakening, going from the initial value to maximum weakening.

Table 4 .2. List of models

 4

	Décollement

Table A .

 A 1. A stratigraphic table that correlates the PYRAMID stratigraphic groups used in Chapter 2 and the BRGM map units along the north Pyrenean cross section. Table A.2. Material constants for each lithology used in decompaction calculations Table A.3. Subsidence of the north Pyrenean foreland Table A.4. Subsidence of the south Pyrenean foreland Table A.5. Field measurements Figure A.1 presents the (minimum) shortening rate in the Pyrenees through time. he data used to plot the shortening rate in the Eastern Pyrenees is already presented in Figure

Table A . 1 .

 A1 Stratigraphy correlation of PYRAMID stratigraphic groups with BRGM map units in the northern Pyrenees along section S3.

Table A .

 A

	2. Values for decompaction calculations a
	Lithology	ρ (kg/m3) b	c (1/km) c	φ 0 (%) d
	Shale	2500	0.51	63
	Siltstone	2530	0.45	59
	Sandstone	2600	0.27	49
	Conglomerate	2600	0.30	50
	Marl	2500	0.59	58
	Limestone	2670	0.70	50
	Gypsum	2310	0.90	50

a hese represent the 'pure' lithologies. Bulk values for each stratigraphic unit were assumed to be a weighted average based on the lithological composition of the stratigraphic unit. Values taken from Vergés et al. [1998]. b Density c Compaction coeicient d Initial Porosity

Table A . 3 .

 A3 Subsidence of the north Pyrenean foreland a

	Strat. BRGM Age hickness Lithology e Water Depth Sea Level ρ c φ0 Total Subs. Tectonic Subs. Rate 0.180 Group b Unit c (Ma) d (km) (km) d,f (km) d,f (kg/m3) g (1/km) h (%) i (km) f Subs. (km) f (mm/y) f,j 2Shl; 26Slt; 30Sst; 7Cgl; 20Mrl; 15Gps DR4 0 0.16 2517 0.48 54 0.194 0.007 0.020 CG AVG e2bM 56 0.118 31Sst; 47Mrl; 22Lst 0 0.17 2569 0.51 53 4.043 1.771 0.031 e3-4 51.9 0.149 5Shl; 26Slt; 40Sst; 8Cgl; 20Mrl; 1Lst RG e2bC 56.5 0.052 100Lst 0 0.16 2670 0.70 50 3.987 1.756 -0.006 0 0.17 2557 0.40 54 0.001 -0.130 0.024 basement RG e2a 57.5 0.165 100Lst 0.005 0.16 2670 0.70 50 3.971 1.762 0.008 56 0.005 0.17 -0.160 -0.230	AVG e1 59.2 0.058 9Sst; 55Mrl; 36Lst 0 0.15 2571 0.60 54 3.884 1.749 0.002 Syn1 AVG c7b-e1 61.6 0.038 100Lst 0 0.15 2670 0.70 50 3.856 1.745 0.006 CSG AVG c7b 66 0.576 8Sst; 87Mrl; 5Lst 0 0.16 2517 0.57 57 3.819 1.719 0.062 g1-2 28.1 0.014 10Shl; 40Cgl; 40Mrl; 10Lst 0 0.04 2557 0.48 55 1.935 1.015 0.010	PG c6bG 69 0.295 54Sst; 46Mrl 0 0.18 2554 0.42 53 3.405 1.534 0.018 CSG e7 33.9 0.179 14Shl; 14Mrl; 69Lst; 0 0.08 2611 0.67 53 1.891 0.957 0.026	PPG c6 75 2.588 7Sst; 92Mrl; 1Lst 0 0.20 2509 0.57 57 3.173 1.426 0.144 3Gps	unconf. 83.6 --0.15 0.23 ---0.492 0.188 0.052 CSG e6 38 0.550 67Sst; 13Cgl; 13Mrl; 0.01 0.12 2591 0.35 50 1.739 0.850 0.073	GFG c3-4 86.3 0.190 100Mrl 0.01 0.23 2500 0.59 58 0.352 0.048 0.028 7Lst	GFG c1-3 93.9 0.128 4slst; 18Sst; 14Cgl; 0.01 0.22 2642 0.56 50 0.019 -0.163 0.020 CSG e5b 41.2 0.407 8Shl; 83Cgl; 8Lst 0.01 0.14 2598 0.35 51 1.296 0.615 0.054	65Lst CSG e5a 45 0.262 6Sst; 76Mrl; 18Lst 0 0.16 2537 0.59 56 0.922 0.411 0.025	basement 100.5 0.005 0.21 -0.205 -0.296 CG e3a-b 51.9 0.213 12Sst; 55Mrl; 33Lst 0 0.17 2569 0.59 54 0.624 0.237 0.033	RG unconf. 56 --0.005 0.17 ---0.374 0.103 -0.007	T1 AVG e2a 57.5 0.162 100Lst 0.005 0.16 2670 0.70 50 0.381 0.113 0.065	CG e3 51.9 0.310 6Sst; 76Mrl; 18Lst 0.005 0.17 2519 0.53 56 0.858 0.376 0.060 PG unconf. 59.2 --0 0.15 ---0.167 0.002 -0.003	CG e3a-b 54 0.087 12Sst; 55Mrl; 33Lst 0.005 0.17 2509 0.56 57 0.630 0.249 0.017 c7b-e1 61.6 0.182 9Sst; 91Mrl 0 0.15 2509 0.56 57 0.172 0.010 0.034 a unconf. 56 --0.005 0.17 ---0.560 0.215 -0.007 Measured in boreholes Dr4, T1, MRL1 (Bureau Exploration-Production des Hydrocarbures; beph.net), and synthetic borehole Syn1. Locations in Figure 2.4. b Using ANR PYRAMID harmonised stratigraphic framework for central and eastern north Pyrenees (Figure 2.3b) [Ford et al., RG e2a 57.5 0.211 100Lst 0.005 0.16 2670 0.70 50 0.567 0.225 0.055 2016; Rougier et al., 2016]. GFG, Grey Flysch Group; PPG, Petites Pyrénées Group; PG, Plantaurel Group; AVG, Aude Valley Group; RG, Rieubach Group; CG, Coustouge Group; CSG, unconf. 59.2 --0 0.15 ---0.371 0.132 0.002 Carcassonne Group.	AVG c7b 66 0.192 9Sst; 91Mrl 0 0.16 2569 0.59 54 0.361 0.117 0.054	PG c6b-7a 69 0.223 19Sst; 81Mrl 0 0.18 2537 0.59 56 0.128 -0.045 0.040	MRL1	CSG e6 37.8 0.072 2Shl; 16Slt; 29Sst; 0.01 0.12 2557 0.41 54 0.415 0.183 0.022	24Cgl; 30Mrl	CSG e5c 41.2 0.119 16Slt; 64Sst; 7Cgl; 0.01 0.14 2573 0.35 52 0.329 0.109 0.020	16Mrl

c

From BRGM 1:50000 geological maps

[Cavaillé et al., 1975a;[START_REF] Cavaillé | Carte géologique de la France au 1/50 000[END_REF] Bilotte et al., 1988b]

. d At top of unit. e Composition as percentages of 'pure' lithologies. Shl, shale; Slt, siltstone; Sst, sandstone; Cgl, conglomerate; Mrl, marl; Lst, limestone; Gps, gypsum. f Relative to present-day sea level. g Bulk density h Bulk compaction coeicient i Initial porosity j Tectonic subsidence rate

Table A . 4 .

 A4 Subsidence in the south Pyrenean foreland a Rough estimate based on stratigraphic section 'La Nou' on Spanish 1:50,000 geological map 255 'La Pobla de Lillet'[START_REF] Vergés | Mapa Geológico de España[END_REF]. A compaction coeicient of 0.6 and initial porosity of 55% were used to estimate decompaction.

	Stratigraphic Name Paleocene Maastrichtian Campanian Santonian Cenomanian pre-Cenom. Bellmunt Coubet Beuda Campdevánol 46.85 Age (Ma) b 56 66 72.1 83.6 86.3 100.5 42.65 45.82 46.18 Armáncies 49.36 Corones 50.69 Sagnari 52.05 Paleocene 55.9 Basement 65 Solsona 33.77 Igualada-Tossa 36.81 Stratigraphic Name Age hickness (km) 0.39 1.07 0.64 0.06 0.03 1.2034 0.1268 0.17 0.7232 1.0766 0.2256 0.465 0.49 0.698 1.41 (Ma) b hickness Water Depth (km) b,c Lower Pedraforca equivalent f Sea Level (km) b,c ρ (kg/m3) d 0 0.165 2600 0 0.160 2550 0.01 0.185 2550 0.01 0.230 2550 0.01 0.230 2550 0.01 0.210 Gombrén g -0.05 0.148 2560 0.015 0.165 2520 0.05 0.166 2390 0.2 0.170 2540 0.1575 0.173 2520 0.005 0.171 2600 0.0375 0.170 2590 -0.05 0.165 2590 0 0.157 Jabalí g -0.05 0.074 2500 0.09 0.105 2530 (km) Water Depth (km) b,c Sea Level (km) b,c ρ (kg/m3) d Total Subs. (km) c 2.025 1.800 0.835 -0.060 -0.160 -0.200 4.283 3.577 3.537 3.552 2.924 1.587 1.385 0.629 -0.157 3.580 3.260 Total Subs. Tectonic Subs. (km) c 0.928 0.839 0.354 -0.220 -0.280 -0.300 1.842 1.593 1.607 1.702 1.456 0.750 0.691 0.240 -0.225 1.660 1.580 (km) c Tectonic Subs. Rate (mm/y) e 0.009 0.080 0.049 0.025 0.001 0.078 -0.038 -0.142 0.098 0.531 0.044 0.117 0.051 0.028 0.182 Subs. (km) c Subs. Rate (mm/y) e Santpedor g Solsona 36.15 0.15 -0.05 0.098 2530 1.592 0.823 -0.141 Igualada-Tossa 36.81 0.92 0.09 0.105 2440 1.637 0.916 0.158 Collbas 40 0.16 0.025 0.131 2580 0.796 0.413 0.072 Pontils 41.5 0.36 -0.01 0.142 2520 0.621 0.305 0.033 Orpí 52.7 0.15 0.015 0.169 2610 0.094 -0.065 0.054 Basement 55.9 0 0.165 -0.165 -0.237 Castellfollit g Solsona 34.31 0.574 -0.05 0.079 2520 1.806 0.964 0.021 Igualada-Cardona 36.81 0.683 0.09 0.105 2410 1.583 0.912 0.133 Collbas 40 0.108 0.025 0.131 2520 0.885 0.488 0.059 Pontils 41.5 0.521 -0.01 0.142 2530 0.750 0.399 0.051 Orpí 52.7 0.049 0.015 0.169 2650 -0.070 -0.174 0.056 Basement 53.9 0 0.168 -0.168 -0.241 Montserrat g (17.1n) 37.2 0.37 -0.1 0.109 2600 1.688 0.505 0.271 37.47 0.05 -0.08 0.112 2600 1.440 0.432 -0.219 (17.2.3n) 37.6 0.4325 -0.04 0.113 2580 1.448 0.460 0.141 38.11 0.028 0.025 0.117 2550 1.152 0.388 0.271 (1.8n) 38.42 0.3995 -0.04 0.120 2570 1.049 0.304 0.110 Planes type waypoint lat long dip azimuth strike stratigraphy comment Bedding 1.3 43.419968 1.651450 51.3 202.4 112.4 c7ag best measurement hard sandstone surface and relatively flat. Bedding 1.3 43.419968 1.651450 66.9 187.9 97.9 c7ag measured in marls not so reliable Bedding 1.3 43.419968 1.651450 52.3 194.8 104.8 c7ag silt/sandstone in tracks average reliability Bedding 1.3 42.928951 1.917953 29.4 200.7 110.7 c6bm unreliable measured in small stream more sandy level of Bedding 1.6 42.928951 1.917953 64.4 204.3 114.3 c7ag very good quality measurement Bedding 1.6 42.928951 1.917953 56.7 189.5 99.5 c7ag quite good quality Bedding 1.6 42.928951 1.917953 53.3 206.0 116.0 c7ag very good Bedding 1.8 42.940868 1.898085 55.1 6.9 276.9 c6bg Bedding 1.8 42.940868 1.898085 64.9 34.5 304.5 c6bg Bedding 1.8 42.940868 1.898110 56.2 20.2 290.2 c6bg Bedding 1.8 42.940868 1.898110 59.6 21.9 291.9 c6bg Bedding 1.8 42.940392 1.897982 58.4 25.1 295.1 c6bg Bedding 1.8 42.940392 1.897982 60.0 16.6 286.6 c6bg Bedding 1.9 42.920753 1.897063 24.2 195.9 105.9 e2 base of thanetian limestones contact with black marls/shales Bedding 1.10 42.958481 1.862358 52.1 6.6 276.6 e3 base of turitella malrs Bedding 1.11 42.958294 1.860599 35.4 29.5 299.5 e3 ilerdian basal limestones Bedding 2.5 42.947819 1.905628 51.7 6.9 276.9 e2 thanetian limestones Bedding 2.5 42.947826 1.905752 60.6 16.5 286.5 e2 thanetian limestones less reliable Bedding 2.5 42.947998 1.905791 55.8 12.5 282.5 e2 thanetian limestones Bedding 2.5 42.947998 1.905791 59.5 11.9 281.9 e2 thanetian limestones less reliable Bedding 2.5 42.949249 1.904099 21.3 32.0 302.0 e3 ilerdian basal limestones not very reliable but best we can Bedding 2.6 42.947330 1.906368 57.2 21.9 291.9 e2 beautiful surface thanetian limestones Bedding 2.6 42.947330 1.906368 56.7 25.6 295.6 e2 same surface different spot Bedding 2.6 42.947918 1.906671 57.2 17.0 287.0 e2 probable bedding plane slightly steeper than s of here but Bedding 2.7 42.951885 1.996932 28.0 5.0 275.0 e3 Bedding 2.7 42.951885 1.996932 29.3 3.2 273.2 e3 Bedding 2.7 42.951885 1.996932 36.7 4.5 274.5 e3 Bedding 2.7 42.951885 1.996932 24.7 6.5 276.5 e3 Bedding 2.7 42.951885 1.996932 24.3 6.8 276.8 e3 Bedding 2.8 42.986580 2.004328 6.2 34.6 304.6 e3c-5 Bedding 2.8 42.986580 2.004328 2.6 6.9 276.9 e3c-5 Bedding 2.8 42.986580 2.004328 21.8 38.3 308.3 e3c-5 Bedding 2.8 42.986580 2.004328 10.5 7.0 277.0 e3c-5 Bedding 2.8 42.986572 2.004330 6.8 7.6 277.6 e3c-5 Bedding 2.8 42.986572 2.004330 12.2 297.0 207.0 e3c-5 Bedding 2.12 42.909618 2.005900 39.5 197.2 107.2 e2 probably thanetian limestone bed base of small gully dug Bedding 2.14 42.909748 2.006697 28.5 185.8 95.8 e2 best plane visible medium reliabiliy Bedding 2.14 42.909687 2.006691 22.2 168.1 78.1 e2 best plane beautiful Planes type waypoint lat long dip azimuth strike stratigraphy comment Bedding 4.5 42.864925 1.903605 49.7 167.0 77.0 n5-6 n7a Bedding 4.6 42.865078 1.910591 46.9 330.6 240.6 n5-6 n7a dip uncertain strike correct Bedding 4.6 42.865070 1.910575 49.4 336.9 246.9 n5-6 n7a dip slightly more certain strike correct Bedding 4.6 42.865051 1.910786 50.3 329.1 239.1 n5-6 n7a good surface but irregular. Bedding 4.7 42.865253 1.911923 86.4 167.5 77.5 n5-6 n7a the steepest bed in the outcrop slightly overturned. other beds aren't but are still subvertical Bedding 4.8 42.864990 1.913026 79.6 161.0 71.0 n5-6 n7a Bedding 4.8 42.864788 1.913069 79.9 155.8 65.8 n5-6 n7a Bedding 4.8 42.864861 1.913131 76.5 154.3 64.3 n5-6 n7a Bedding 4.8 42.864861 1.913131 78.1 156.2 66.2 n5-6 n7a Bedding 4.8 42.864624 1.913384 82.6 164.0 74.0 n5-6 n7a Bedding 4.9 42.864182 1.926726 70.3 167.5 77.5 n5-6 n7a Bedding 4.9 42.864182 1.926726 74.7 164.8 74.8 n5-6 n7a Bedding 4.9 42.864182 1.926726 67.3 161.6 71.6 n5-6 n7a Bedding 4.9 42.864182 1.926726 74.3 165.7 75.7 n5-6 n7a Bedding 4.10 42.864925 1.935388 63.6 164.4 74.4 n5-6 n7a Bedding 4.11 42.814030 1.967165 64.0 356.4 266.4 jD unclear could be this. matches with map Bedding 4.13 42.817894 1.984360 50.8 162.6 72.6 n1-3 weathered surface Bedding 4.13 42.817894 1.984360 38.6 181.4 91.4 n1-3 weathered surface Bedding 4.14 42.823914 1.995807 59.1 147.3 57.3 U1 marty most recurring planar feature possible bedding plane. that would mean overturned northern flank. Bedding 4.18 42.886559 1.953603 31.1 275.2 185.2 n4a no good surfaces but brdding is approximately correct Bedding 5.1 43.044670 1.933478 12.5 189.6 99.6 e5b-1 Bedding 5.1 43.044670 1.933478 20.9 200.7 110.7 e5b-1 Bedding 5.1 43.044670 1.933478 18.5 193.7 103.7 e5b-1 Bedding 5.2 43.031769 1.948481 23.4 194.5 104.5 e5b-1 most reliable bed in a sandstone outcrop with a lot of troughs. Bedding 5.3 43.059387 1.985749 14.2 28.9 298.9 e5c2 footwall syncline? measured in marly bed Bedding 5.4 43.065891 1.980308 28.6 35.4 305.4 e5c2 best approximation of bedding taken in sandstone. Bedding 6.4 42.843102 1.877044 83.0 309.5 219.5 n4b estimated possible bedding irregular. Bedding 6.5 42.843792 1.877202 58.1 357.7 267.7 n4b Bedding 6.6 42.845997 1.875002 67.5 1.8 271.8 n4b measured next to waterfall Bedding 6.7 42.845997 1.875002 66.8 355.1 265.1 n5-6 n7a not very reliable Bedding 6.7 42.846085 1.875109 79.7 9.2 279.2 n5-6 n7a probably slightly wrong strike Bedding 6.7 42.845943 1.874590 70.0 1.2 271.2 n5-6 n7a reasonably ok Bedding 6.7 42.846455 1.874623 77.2 353.3 263.3 n5-6 n7a good Bedding 6.8 42.849926 1.875373 82.1 178.5 88.5 n5-6 n7a Bedding 6.8 42.849926 1.875373 82.6 179.1 89.1 n5-6 n7a Planes type waypoint lat long dip azimuth strike stratigraphy comment Bedding 9.2 42.893276 1.854544 30.4 201.4 111.4 c5b-6 undulating Bedding 9.3 42.891582 1.851803 46.6 160.7 70.7 c5a2 irregular but ok Bedding 10.1 42.951637 1.922155 21.2 350.6 260.6 e3a1 embedded in road Bedding 10.1 42.951637 1.922155 25.4 332.0 242.0 e3a1 idem best Bedding 10.1 42.951637 1.922155 19.6 323.7 233.7 e3a1 idem Bedding 10.1 42.951763 1.922121 16.2 4.4 274.4 e3a1 idem Bedding 10.1 42.951763 1.922121 17.2 7.1 277.1 e3a1 idem Bedding 10.1 42.951763 1.922121 13.4 17.0 287.0 e3a1 idem Bedding 10.2 42.949974 1.922069 22.0 12.5 282.5 e3a1 approximation Bedding 10.2 42.949974 1.922069 17.5 354.9 264.9 e3a1 approximation Bedding 10.2 42.949974 1.922069 22.8 0.3 270.3 e3a1 approximation in marls Bedding 10.2 42.949974 1.922069 13.7 30.8 300.8 e3a1 approximation Bedding 10.2 42.949974 1.922069 14.5 356.3 266.3 e3a1 approximation Bedding 10.3 42.959412 1.900889 11.4 326.2 236.2 e3a1 approximate Bedding 10.4 42.962368 1.903509 12.0 358.4 268.4 e3b approximated strike correct and dip estimated. dip difficultnto see in outcrop Bedding 10.4 42.962296 1.903571 7.2 317.4 227.4 e3b much better Bedding 10.4 42.962296 1.903571 10.1 300.9 210.9 e3b much better Bedding 10.4 42.962296 1.903571 11.2 345.1 255.1 e3b dip estimated Bedding 10.5 42.962135 1.889214 19.4 8.3 278.3 e3b Bedding 10.5 42.962135 1.889214 18.3 17.8 287.8 e3b Bedding 10.5 42.962135 1.889214 13.8 341.4 251.4 e3b approximated Bedding 10.5 42.962135 1.889214 19.0 22.2 292.2 e3b Bedding 10.5 42.962135 1.889214 15.9 348.0 258.0 e3b Bedding 10.5 42.962063 1.889216 18.4 7.2 277.2 e3b Bedding 10.6 42.971687 1.911127 10.3 251.7 161.7 e3b Bedding 10.6 42.971687 1.911127 3.9 331.5 241.5 e3b not as good as previous Bedding 10.6 42.971561 1.911532 9.8 17.3 287.3 e3b very good underside of bed Bedding 10.6 42.971561 1.911532 5.8 355.1 265.1 e3b Bedding 10.6 42.971561 1.911532 10.9 20.4 290.4 e3b Bedding 10.6 42.971561 1.911532 13.2 29.0 299.0 e3b Plane Type 11.1 42.961346 1.775601 62.7 35.3 305.3 c7b-e1 recurring structural plane in danian limestone outcrop Plane Type 11.1 42.961346 1.775601 56.2 55.2 325.2 c7b-e1 idem planes very planar on outcrop scale but not on fieldbook-scale Plane Type 11.1 42.961346 1.775601 73.8 270.7 180.7 c7b-e1 possible conjugate? sense: normal Plane Type 11.1 42.961369 1.775589 56.6 75.3 Planes type waypoint lat long dip azimuth strike stratigraphy comment Bedding 18.10 42.856155 1.877836 35.9 213.2 123.2 n5-6 n7a Bedding 18.10 42.856159 1.877890 33.9 183.0 93.0 n5-6 n7a approx 3d Bedding 18.11 42.864189 1.886069 29.6 145.3 55.3 n5-6 n7a could be loose block Bedding 18.12 42.865780 1.888044 71.0 358.7 268.7 n5-6 n7a Cleavage 18.12 42.865791 1.888031 13.3 213.7 123.7 n5-6 n7a Bedding 18.13 42.865807 1.888121 66.5 351.4 261.4 n5-6 n7a Cleavage 18.13 42.867519 1.889999 35.9 152.8 62.8 n5-6 n7a Bedding 18.14 42.892696 1.927176 68.7 321.0 231.0 n5 Plane Type 18.17 42.892105 1.932804 80.0 161.3 71.3 n5-6 n7a Plane Type 18.17 42.892071 1.932886 78.2 167.0 77.0 n5-6 n7a Plane Type 18.17 42.891884 1.933066 63.7 170.6 80.6 n5-6 n7a Plane Type 18.17 42.891884 1.933066 60.0 180.0 90.0 n5-6 n7a Cleavage 18.17 42.891884 1.933067 85.4 352.4 262.4 n5-6 n7a Bedding 18.18 42.891178 1.934045 29.5 357.7 267.7 n5-6 n7a question mark Bedding 18.18 42.891216 1.934088 25.8 12.9 282.9 n5-6 n7a sedimentological features visible on eroded surface confirm this plane is bedding Plane Type 19.1 42.766727 1.850349 89.0 345.5 255.5 meso meta marmetamorphic foliation in marble Bedding 20.9 42.784637 1.975004 63.2 196.8 106.8 d4-7 overturned devonian close to hinge Bedding 20.10 42.785149 1.976014 11.6 140.6 50.6 d4-7 top flank 1 Bedding 20.10 42.785160 1.976074 8.0 176.2 86.2 d4-7 top flank 2 Bedding 20.10 42.785145 1.976202 79.3 188.7 98.7 d4-7 bottom flank 1 Cleavage 20.10 42.785114 1.976191 45.7 215.8 125.8 d4-7 axial plane cleqvage Bedding 20.13 42.776936 2.140635 62.6 32.4 302.4 turonian Bedding 20.13 42.777752 2.139368 40.5 355.7 265.7 turonian Bedding 20.13 42.777370 2.138630 47.5 3.0 273.0 turonian Bedding 20.13 42.778290 2.138091 56.5 343.4 type waypoint lat long dip azimuth strike stratigraphy comment Sinistral Strike S1.6 42.928951 1.917953 54.9 282.7 192.7 c7aG slickenside 1 Sinistral Strike S1.6 42.928951 1.917953 71.2 283.3 193.3 c7aG slickenside 2 Unknown 1.8 42.940353 1.897991 65.0 108.0 18.0 c6bG slickenside 3 Unknown 2.6 42.947918 1.906671 73.4 21.5 291.5 e2 plane with striations Dextral Strike S 4.1 42.867420 1.897573 64.3 245.7 155.7 n5-6 n7a with slickensides still visible under coating of mud. sense: dextral Dextral Strike S 4.4 42.866379 1.900385 68.0 54.0 324.0 n5-6 n7a very small and irregular but with slickensides so measured it Normal Fault 6.4 42.843102 1.877044 43.7 17.9 287.9 n4b extremely flat recurring multiple times on other side stream in western cliff. has striations. Reverse Fault 11.4 42.965450 1.772800 32.4 300.3 210.3 e2a with slickensides reverse Unknown 12.7 42.883137 2.038856 36.1 308.5 218.5 l1-4 no signs of movement direction Unknown 12.9 42.881886 2.038496 71.2 77.3 347.3 n1-3 probable bedding but moves around a lot through outcrop and is hard to recognise Unknown 12.9 42.881886 2.038496 70.7 83.8 353.8 n1-3 idem Unknown 12.9 42.881886 2.038496 66.6 84.2 354.2 n1-3 idem Unknown 12.9 42.881886 2.038496 77.3 76.0 346.0 n1-3 idem Unknown 12.9 42.881886 2.038496 69.1 79.1 349.1 n1-3 idem Unknown 16.1 42.856220 2.101801 55.4 306.2 216.2 n4b not completely flat has striations 1 Unknown 16.1 42.856270 2.101821 41.6 239.0 149.0 n4b good surface but small so could only fit 2 corners of ipad on fault. has striations 2 Dextral Strike S 16.1 42.856293 2.101766 48.2 208.2 118.2 n4b bigger fault plane. with striations 4 Dextral Strike S 16.2 42.856438 2.102143 47.0 285.0 195.0 n4b has striations 5 dextral Dextral Strike S 16.2 42.856457 2.102132 38.9 245.0 155.0 n4b has striations 6 dextral fault along which the bedding and cleavage have been Lineations type waypoint lat long plunge azimuth stratigraphy comment Slickenside 1.6 42.928951 1.917953 2.0 7.1 c7ag slickenside 1 sense unclear maybe sinistral Slickenside 1.6 42.928951 1.917953 30.4 201.9 c7ag slickenside 2 sinistral normal Slickenside 1.8 42.940353 1.897991 46.0 50.3 c6bg slickenside 3 Slickenside 2.6 42.947918 1.906671 73.8 54.6 e2 striations in thanetian limestones Slickenside 3.2 42.882740 1.809537 68.2 127.4 n7c-d northern block up contrary to regional sense Slickenside 4.1 42.867420 1.897573 21.4 317.5 n5-6 n7a with slickensides still visible under coating of mud. sense: dextral Fold Axis 4.4 42.866379 1.900385 1.7 265.2 n5-6 n7a possible cleavage bedding intersection Slickenside 4.4 42.866379 1.900385 23.2 332.8 n5-6 n7a dextral Slickenside 6.4 42.843102 1.877044 41.9 6.9 n4b measured on fault plane sense: looks like normal but feels the same as reverse. unclear but guessing normal. striations have been partly dissolved. Slickenside 7.1 42.916245 1.982443 28.7 263.9 c7aG no growth of mineral just striation in sandstone. sense: dextral Slickenside 11.4 42.965450 1.772800 10.8 237.3 e2a top to ene Line Type 1 15.5 42.879311 1.800301 2.0 175.5 h2 intersection of probable bedding with joint Line Type 2 15.5 42.879337 1.800309 21.8 187.7 h2 intersection of probable cleavage with joint. Fold Axis 15.6 42.877052 1.798257 47.4 121.3 d4-7 on bedding plane of limestone Fold Axis 15.11 42.849213 1.753293 7.9 76.8 k-o bedding-cleavage intersection not easily measured best location Slickenside 16.1 42.856251 2.101811 26.2 19.5 n4b striations 1. direction impossible to tell here. Slickenside 16.1 42.856205 2.101872 3.5 329.8 n4b striations 2. direction unknown Slickenside 16.1 42.856277 2.101780 76.7 45.1 n4b most common direction of striations on irregular bedding plane. direction unknown 253.4 turonian Faults Normal Fault 16.13 42.864033 2.114037 22.0 25.8 295.8 n6a2 deformed in bands Reverse Fault 17.8 42.812855 1.979467 50.8 228.8 138.8 M1-4 marty Slickenside 16.1 42.856285 2.101823 1.1 125.3 n4b striations 4 dextral. (top to wnw) fault contact with metamorphic. bright white marble. Reverse Fault 17.8 42.812824 1.979453 34.0 200.0 110.0 M1-4 marty Slickenside 16.2 42.856430 2.102158 2.2 206.1 n4b striations 5 Mary's measurement Reverse Fault 17.8 42.812874 1.979467 36.9 221.2 131.2 M1-4 marty Slickenside 16.2 42.856392 2.102188 2.7 332.6 n4b striation 6 dextral i think measurement three ipad Reverse Fault 17.8 42.812977 1.979425 55.6 112.4 22.4 M1-4 marty Slickenside 17.1 42.823875 1.995799 51.5 181.0 U1 marty on bedding plane 5 might indicate movement smaller fault and actually more inclined than fault but a Slickenside 17.1 42.823853 1.995754 17.3 68.7 U1 marty direction 2 of lineation on same plane repeating surface Reverse Fault 17.8 42.812901 1.979399 33.7 55.2 325.2 M1-4 marty lineation on joint measured on slightly loose hanging piece. less inclined than second fault so dip azimuth different but it Slickenside 17.1 42.823978 1.995820 17.2 184.2 U1 marty actual plunge a few degrees steeper is a big plane Reverse Fault 17.8 42.812851 1.979373 30.4 91.9 1.9 M1-4 marty bedding cleavage intersection in cleavage heavy outcrop if best approximation of actual fault because no really Fold Axis 17.6 42.816528 1.987627 60.6 249.7 m6 marty bedding is corrrect representative planes could be found Slickenside 17.9 42.811829 1.988332 29.4 169.9 meso meta marslickensides indicate normal movement 345.3 c7b-e1 Plane Type 11.1 42.961372 1.775583 85.0 296.2 206.2 c7b-e1 conjugate? Bedding 11.1 42.961681 1.776178 14.4 339.8 249.8 c7b-e1 best bedding Normal Fault 17.9 42.811928 1.988314 26.8 224.6 Slickenside 17.9 42.811825 1.988193 30.4 181.1 meso meta marattempt 2 of slickensides same place 134.6 meso meta marsmall fault cuts through beds reducing dip pf beds below Slickenside 17.10 42.813526 1.986720 31.7 337.9 meso meta maron fault sinistral
	Folgueroles Banyoles Tavertet (1.9n) Mediona-La-41.5 40 43 40.13 0.155 0.58 0.417 0.068 41.25 0.12 41.52 0.4238 Bedding 2.15 42.909534 2.009521 Bedding 3.1 42.889126 1.792310 Bedding 3.2 42.882740 1.809537 Bedding 3.3 42.892540 1.806664 Bedding 6.8 42.849892 1.875331 0.125 0.075 0.02 -0.05 -0.05 -0.05 55.0 253.7 36.4 201.0 77.2 3.4 37.4 219.0 78.3 Bedding 6.8 42.849892 1.875331 76.5 Bedding 6.8 42.849892 1.875331 86.1 Bedding 6.9 42.859928 1.879832 18.2 Bedding 11.1 42.962032 1.776134 21.4 14.6 Bedding 11.3 42.964359 1.773709 19.7 24.9 Bedding 11.3 42.964359 1.773709 16.5 7.5 Bedding 11.3 42.964359 1.773709 17.7 6.4 Normal Fault 17.9 42.811848 1.988303 46.2 Slickenside 17.11 42.813526 1.987253 56.5 110.8 meso meta maron previous fault plane direction ambiguous maybe normal 0.131 2590 2.040 1.000 0.074 0.142 2550 1.860 0.890 0.269 0.150 2650 1.140 0.480 0.093 0.132 2570 0.621 0.116 0.033 0.140 2580 0.544 0.078 0.210 0.142 2570 0.410 0.022 163.7 e2 111.0 n7c-d 273.4 n7c-d no clear bedding but probably subvertical and the strike is 129.0 c3-4 right on border not clear wich unit 180.6 90.6 n5-6 n7a 180.7 90.7 n5-6 n7a 180.4 90.4 n5-6 n7a srike best approached 38.8 308.8 n5-6 n7a not a very good outcrop but dip is consistent enough 284.6 c7b-e1 also relatively good 294.9 e2a good surface 277.5 e2a idem 276.4 e2a idem 229.3 139.3 meso meta mar another fault that steepens dip of strata above and below it but has slickensides indicating normal movement Normal Fault 17.9 42.811852 1.988248 30.9 223.9 133.9 meso meta mar on second fault plane on same block direction ambiguous attempt 2 of fault plane with slickensides actual plane with slickensides this time. (rough plane) Slickenside 17.11 42.813576 1.987176 36.5 maybe dextral 63.6 meso meta mar 0.018 Bedding 3.3 42.892563 1.806816 46.3 205.9 115.9 c3-4 same material north side of road everywhere to believe Bedding 11.3 42.964359 1.773709 14.6 14.4 284.4 e2a not as good Sinistral Strike S17.10 42.813530 1.986720 72.9 257.5 167.5 meso meta marwith slickensides sinistral Slickenside 18.5 42.843109 1.876790 41.5 358.5 n4b normal
	Penya Cadí 3.3 Bedding Bedding Bedding 11.4 Normal Fault 17.11 46.26 49 42.892563 1.806816 0.143 0.297 6.10 42.860176 1.880414 0.0125 0.015 41.2 42.965450 1.772800 21.6 42.813549 1.987239 Fold Axis 18.6 42.846214 1.875133 75.4 213.2 53.3 349.4 85.5 304.7 n5-6 n7a 0.167 0.174 123.2 c3-4 358.8 268.8 n5-6 n7a 2720 2680 259.4 e2a 190.6 100.6 meso meta marhas slickensides direction ambiguous maybe normal 0.560 0.180 0.039 0.350 0.070 idem bad quality fairly good bedding cleavage intersection Salut Basement 55.9 0 0.165 -0.165 -0.237 Bedding 3.4 42.892467 1.808852 53.1 201.6 111.6 c5a1 high quality measurement overturned Bedding 3.7 42.883301 1.843118 72.6 166.3 76.3 n7c-d or c1-3 sandstone bed interbedded in black marls or shales exposed Bedding 6.10 42.860153 1.880492 89.0 6.9 276.9 n5-6 n7a could be cleavage instead Cleavage 6.10 42.860153 1.880492 37.8 19.8 289.8 n5-6 n7a could be bedding instead not high quality measurement Bedding 11.4 42.965450 1.772800 21.2 3.6 273.6 e2a Slickenside 18.15 42.894432 1.929458 48.3 224.3 n5 striae 1 reverse medium Bedding 11.4 42.965450 1.772800 27.6 12.4 282.4 e2a idem Slickenside 18.15 42.894382 1.929474 30.1 175.1 n5 striae 2 reverse and normal? 0.045 Basement 55.9 0 0.165 -0.170 in path. dip estimated strike correct Bedding 11.4 42.965450 1.772800 21.9 21.5 291.5 e2a better Slickenside 18.15 42.894470 1.929421 12.4 18.1 n5 striae 3 sinistral -0.240 Bedding 3.7 42.883350 1.843082 76.2 169.0 79.0 n7c-d or c1-3 another sandstone bed dip estimated strike correct Bedding 6.11 42.861431 1.882061 63.1 193.1 103.1 n5-6 n7a strike approximate Bedding 11.5 42.966923 1.773761 19.4 350.1 260.1 e2a very good underside of bed but small surface making it Fold Axis 19.16 42.813480 1.852060 3.7 340.4 meso meta marty
	Bedding Bedding Fold Axis	3.8 20.10 6.11	42.883591 1.842013 42.861431 1.882061 42.785137 1.976164	47.0 22.6 166.9 75.5 142.6 d4-7 76.9 c1-3 183.9 93.9 n5-6 n7a different partnofnsame hinge outcrop of more marly part in path dip not precise strike may be cleavage difficult to keep ipad in same orientation
	Puig-reig g 0.080 0.105 0.131 173.5 83.5 n5-6 n7a 2540 2400 2540 16.1 6.9 276.9 c7aG 29.2 211.6 121.6 c7aG 346.3 256.3 e2a 201.8 111.8 c6 213.4 123.4 c6 80.6 4.0 274.0 n5-6 n7a 37.0 327.2 237.2 marble breccia?bedding undulates a lot 3.100 1.490 2.910 1.490 1.430 0.750 dip varies strike correct vey high quality approx 20 m n along road from previous waterfall) topside of bed bedding contacts much more chaotic here. 0 0.076 0.231 correct limestone beds turning in anticline core. probable southern limb bedding measured behind travertine waterfall (not directly behind water more next to but strike reaches behind may be bedding. 27.9 343.0 253.0 marble breccia?bedding undulates a lot 43.9 310.6 220.6 marble breccia?bedding undulates a lot 31.9 200.1 110.1 c7aG 350.7 260.7 e2a very good topside of thin bed 326.8 236.8 e2a 2.8 approximation of hinge newr other measurements 108.7 d4-7 relatively good underside approx 10 m n along road from 45.9 -0.05 0.09 0.125 62.7 42.917164 1.983906 0.5632 1.6628 0.0975 42.869225 1.894742 34.36 Igualada-Tossa 36.81 Solsona Folgueroles 40 Bedding 4.1 Bedding 7.2 Bedding 7.1 42.916470 1.982203 Bedding 11.5 42.967197 1.773628 31.9 Bedding 3.10 42.909325 1.864726 30.4 Bedding 3.11 42.908852 1.864850 47.2 Cleavage 6.11 42.861465 1.882128 Cleavage 6.12 42.821552 1.859688 Cleavage 6.12 42.821552 1.859688 Cleavage 6.12 42.821602 1.859777 Bedding 7.1 42.916245 1.982443 Bedding 11.5 42.966923 1.773761 20.5 Bedding 11.5 42.966923 1.773761 14.5 Fold Axis 20.10 42.785091 1.976126 Lineation 20.14 42.776527 2.141088 21.3 black dolomite breccia previous measurement Bedding 11.5 42.967197 1.773628 20.1 19.1 289.1 e2a Lineation 20.14 42.776524 2.141097 30.7 353.0 black dolomite normal inside very massive bed missing piece of small layer. approx 10 m north along road from previous Lineation 21.6 42.767666 2.299687 73.6 119.3 black marble lineation in marble
	Banyoles Bedding 4.1 Bedding Bedding 11.5 Tavertet Bedding 4.1 Bedding 4.1 Bedding Bedding Bedding 11.5 Penya Plane Type 4.1 Bedding Plane Type 4.1 Bedding Bedding 11.5	7.2 7.2 7.2 7.2 7.2	41.5 42.869225 1.894742 0.218 42.917130 1.983954 0.075 75.4 42.967197 1.773628 12.1 43 0.127 0.02 42.867401 1.897466 82.7 42.867401 1.897466 77.1 42.917168 1.984024 42.917183 1.984108 42.967422 1.773374 21.0 46.26 0.0967 0.0125 42.867302 1.897581 79.5 42.917091 1.984259 42.867302 1.897581 42.0 42.917198 1.984292 42.967506 1.773338 18.4	165.5 11.9 357.8 356.8 356.3 16.8 16.2 15.6 67.9 17.4 351.5 18.7 67.9	0.142 75.5 n5-6 n7a 2520 354.6 264.6 c7aG 267.8 e2a 0.150 2640 266.8 n5-6 n7a 266.3 n5-6 n7a 345.0 255.0 c7aG 23.2 293.2 c7aG 285.6 e2a 0.167 2670 337.9 n5-6 n7a 3.5 273.5 c7aG 261.5 n5-6 n7a 353.1 263.1 c7aG 337.9 e2a	1.260 dip varies strike correct underside of bed slightly chaotic. same location as previous 0.640 0.141 0.930 0.430 not very reliable relatively good underside approx 10 m along road from 0.029 0.760 0.330 bedding not clear many planes. plane 1 previous 0.019 bedding not clear many planes plane 2 good underside of bed with thin loose layers. approx 14 m
	Cadí Plane Type 4.1 Bedding Basement Plane Type 4.1 Bedding Bedding 11.5 Plane Type 4.1 Bedding	7.3 7.4 7.7	49 42.867302 1.897581 0.4605 42.920551 1.989191 0.015 31.3 55.9 0 42.866760 1.899530 70.5 42.932060 1.998419 42.967506 1.773269 16.6 42.866760 1.899530 84.4 42.938812 1.997894	185.5 69.7 73.6 13.7 4.4 357.2 13.6	0.174 95.5 n5-6 n7a 2590 32.4 302.4 c7aG 0.165 343.6 n5-6 n7a 344.3 254.3 c7b 274.4 e2a 267.2 n5-6 n7a 12.5 282.5 e2a	0.650 bedding not clear many planes plane 3 0.280 north along road -0.170 another underside very good quality approx 5 m north along 0.075 -0.240 probable bedding similar strike as other measurements and very irregular surface road from previous.
	Bedding Bedding	11.5	7.7	42.938812 1.997894 42.967506 1.773269	26.7	41.4 357.3	32.2 267.3 e2a 302.2 e2a	what's visible down in riverbed from distance although hat better but steeper than general outcrop underside slightly undulating. approx 10 m north from
	Bedding		7.7	42.938812 1.997787		13.8	327.6	237.6 e2a	might also be joint previous. limestone more evenly bedded (30-50 cm) and
	Plane Type 4.1 Bedding	7.7	42.866760 1.899530 42.938812 1.997787	58.5	284.0 13.9	194.0 n5-6 n7a 356.5 266.5 e2a	darker beige in colour.
	Bedding Bedding Bedding	4.4 11.5	9.1	42.866379 1.900299 42.899036 1.850392 42.967506 1.773269	87.0 24.4	352.4 24.8 355.3	262.4 n5-6 n7a 174.4 84.4 c5b-6 265.3 e2a	bedding very approximate same location underside not so good
	Bedding Bedding Bedding	4.4 11.5	9.2	42.866379 1.900299 42.893642 1.854856 42.967506 1.773269	82.5 21.2	356.3 18.3 358.8	266.3 n5-6 n7a 167.2 77.2 c5b-6 268.8 e2a	might be loose underside same location again slightly better than previous
	Bedding Bedding next page: 4.4 Bedding 4.4 Bedding Bedding 11.5	9.2 9.2	42.866379 1.900299 42.893593 1.854873 42.866432 1.900370 42.893444 1.854515 42.967655 1.773161	83.7 85.1 19.6	353.3 50.6 358.7 41.8 351.4	263.3 n5-6 n7a 158.9 68.9 c5b-6 268.7 n5-6 n7a 177.9 87.9 c5b-6 261.4 e2a	very good fixed in streambed gutter underside approx 5 m north from previous. wet.
	Cleavage Bedding Table A.5. Field data 4.4 42.866226 1.900417 9.2 42.893444 1.854515 Bedding 9.2 42.893276 1.854544	49.3	150.7 43.1 35.4	60.7 n5-6 n7a 194.5 104.5 c5b-6 181.7 91.7 c5b-6	possible cleavage measured in more strong possibly limey very good plane part. no really clear plane ver good plane

a Borehole locations in Figure 2.9. b At top of unit. c Relative to present-day sea level. d Density e Tectonic subsidence rate f g Converted from Vergés et al. [1998].

Jammes, S., Manatschal, G., Lavier, L., and Masini, E., 2009, Tectonosedimentary evolution related to extreme

f See Chapter 3. g See Appendix C for details, igures, and descriptions.

Acknowledgements References

Acknowledgements

his study was funded by the ANR (France) PYRAMID research project. he French-Norwegian Foundation (13-06 PYR-FFTP; Sedimentary basin and North Pyrenean foreland fold and thrust belt formation) supported study visits to the University of Bergen, Norway. Collaboration with CSIC Barcelona, Spain was funded by the project ALPIMED (PIE-CSIC-201530E082). A stratigraphic correlation of the BRGM 1:50,000 geological maps supporting Figure 2.3 is available as Table A.1 in the supplementary information. Detailed subsidence data supporting Figure 2.7 are available as Tables A.2, A.3, and A.4 in the supplementary information. Field data used to constrain the structure in the northern Pyrenees (Figures 2.2, 2.4, 2.5 and 2.6) are available as Table A.5 in the supplementary information. We thank the PYRAMID team, in particular M. de Saint Blanquat for fruitful discussions on the structure and nature of the Metamorphic Internal Zone in the Saint-Barthélémy area. We thank Andrew Leier and one anonymous reviewer for their insightful comments that helped to signiicantly improve the manuscript.

Acknowledgements

Supplementary methods and model descriptions are available in Appendix B. Computing hours are part of SIGMA2 high performance computing allocation project NN4704K: 3D forward modeling of lithosphere extension and inversion. Ritske Huismans thanks and is indebted to Josep Anton Muñoz for sharing his insights into Pyrenean geology and for highly fruitful discussions.

is the stress tensor, the spatial dimensions, is density and the gravitational acceleration. In the viscous regime, materials follow a power-law rheology, meaning strain rate depends on stress to a power .

is the strain rate tensor, is a pre-exponential material constant, is the deviatoric stress tensor, the exponent, is the activation energy, is pressure, is the activation volume, is the gas constant and is temperature. Material constants , , , and are determined from laboratory data [Karato and Wu, 1993;Gleason and Tullis, 1995].

Frictional (brittle) behaviour is approximated by choosing an efective viscosity that places the state of stress at yield.

is the second invariant of the strain rate tensor, deined as . Yield stress is determined using a Drucker-Prager yield criterion, which is equivalent to the Coulomb failure criterion in 2D. (B.4) is the second invariant of the deviatoric stress, cohesion, and the angle of internal friction.

Strain sotening of frictional materials is simulated by reducing and linearly with increasing strain in the range 0.5 < < 1.5, where and is the second invariant of the deviatoric strain rate tensor. he initial value of = 15° for the upper crust has been chosen to represent the strength of the crust with hydrostatic pore luid pressure. he fully weakened value = 2° has been experimentally determined to best approximate behaviour of natural examples [START_REF] Lavier | Factors controlling normal fault ofset in an ideal brittle layer[END_REF]Huismans and Beaumont, 2007]. his ensures that all shortening accommodated by the upper and lower plates will be correctly accounted for, even though our tracked boundary does not necessarily represent the suture zone. Our method allows easy comparison to quantiied shortening in natural systems, where shortening on either side of a tectonic boundary can be more reliably reconstructed back in time than the location of the drainage divide. To further facilitate comparison of our results to natural systems, we eliminated the efect of early distributed deformation by only counting shortening accommodated since deformation began to localise in the irst crustal shear zone. When shear zones are already present at the onset of convergence (Models 1 and DR1), shortening was counted from the onset of convergence.

Supplementary model descriptions

We compared our results to two models that have been previously published [Erdős et al., 2014[Erdős et al., , 2015]]. he main characteristics of all three models are listed in Table B.2. For completeness, we describe the full evolution of the two supplementary models below.

DR1: rit inheritance + shale + surface processes he pre-orogenic rit phase of model DR1 is practically identical to that of model 1 described in the main paper (phase 0). A narrow, symmetric rit is bounded by two conjugate frictional-plastic shear zones that root in the weak middle crust. Two similar shear zones in the lithospheric mantle accommodate extension in the mantle. he only diference is a lack of gravitational sliding of the sedimentary cover in model DR1.

Phase 1 is the symmetric inversion of the rit. Both inherited shear zones are preferentially reactivated simultaneously during the irst 25 km of convergence (Figure B.2a). A keystone block is uplited, and the cover does not slide into the foredeeps, but remains on top.

Phase 2 begins ater 25 km of convergence, when a single new shear zone is created through the crust (Figure B.2b). his new shear zone creates an asymmetry that allows subduction of the lower crust and lithospheric mantle, while the other, inherited shear zones are temporarily abandoned, as shown in the strain rate plot (Figure B.2b). At 10 My (50 km of convergence), sedimentation is started, illing the proforedeep. he single pro-wedge thrust sheet steepens as it overthrusts this new basin, reactivating the retro-wedge very slowly ater 55 km of convergence. A new thrust sheet is accreted to the pro-wedge ater 80 km of convergence. his uplits the older, upper thrust sheet and increases shortening rates in the retro-wedge (Figure B.2c). he pro-foreland basin is uplited on top of the thick-skinned thrust sheet and a new thin-skinned thrust is formed in the footwall (Figure B.2c). More convergence results in an outward growing thick-skinned pro-wedge, while the sedimentary cover is mostly uplited and eroded along with the thick-skinned thrust sheets (Figure B.2d). Ater 225 km of convergence, the orogen comprises a wide thick-skinned pro-wedge and well-developed retro-wedge. hin-skinned thrusts always link back to the nearest thick-skinned shear zone. Supplementary models hree supplementary models are described here. hese models are similar to M6, M7 and M8 described in Chapter 4, but feature an even weaker salt décollement. he efective viscosity of these supplementary models is 10 18 Pas. Here we only describe the evolution during the irst 125 km of convergence, as opposed to 300 km. he salt décollement in these models is so weak that deformation reaches the edge of the model domain, which invalidates the model result ater this point.

127

Model S1 he overall development of model S1 is very similar to model M6. During the irst 5 My, the weak salt décollement allows the sedimentary cover to slide into the rit. he rit is symmetrically inverted along both inherited shear zones, past full inversion at 6.5 My. As soon as the central block is uplited, the cover slides down towards both sides. At 15 km beyond full inversion, a new shear zone is created that establishes an asymmetry and the retro-wedge is abandoned. A large thin-skinned pro-foreland fold-andthrust belt forms as the sedimentary cover of the lower plate is stacked against the orogen. A new prowedge thrust sheet is accreted around 77.5 km ater full inversion. his uplits the keystone block and older pro-wedge thrust sheet. he rigid top of the keystone block is separated from its ductile root as two large blocks slide downward onto the upper plate. he retro-foreland fold-and-thrust belt is transported over the weak salt décollement in front of these basement blocks. Meanwhile, in the pro-foreland, a wide thinskinned fold-and-thrust belt has formed, and crustal thrust sheets in the pro-wedge almost completely overlap each other. Ater 125 km of convergence, the crustal structure is extremely similar to model M6. A narrow thick-skinned orogen lanked by a wide thin-skinned pro-foreland fold-and-thrust belt and slightly narrower retro-foreland fold-and-thrust belt (Figure C.1a). he thin-skinned foreland fold-andthrust belts are a bit wider.

Model S2

Model S2 includes sedimentation, similar to model M7. he evolution prior to the onset of synorogenic sedimentation is the same as S1. Ater 10 My, sedimentation starts and the thin-skinned deformation front in the pro-foreland rapidly migrated outward. his creates wide, relatively stable zones of wedgetop basins separated by narrow zones of intense deformation. Ater 75 km of convergence, a new prowedge thrust sheet is accreted. his uplits the keystone block and older pro-wedge thrust sheet while the sedimentary cover of the new pro-wedge thrust sheet slides of the thrust sheet and into the foredeep under the inluence of gravity. Ater 125 km of convergence, the geometry remains similar to model M7, except the pro-foreland fold-and-thrust belt is wider (Figure C.1b).

Model S3

Erosion is included in model S3. he irst 10 My are almost identical to S1 and S2: Symmetric inversion of the rit, full inversion at 6.5 My. he cover slides of the keystone block and the asymmetry is established at 15 km of shortening beyond full inversion. Similar to model s2, the pro-foreland establishes wide wedgetop basins separated by narrow zones of intense deformation. he cover of each newly accreted crustal thrust sheet slides of into the foreldeep. Erosion of the central orogen causes deformation to focus there. Ater 125 km of convergence, the crustal structure of the central orogen is more chaotic than models M8 and S2 (Figure C.1c).

Appendix D Raman data

Appendix D. Raman data 130

Raman spectra of carbonaceous material

Raman spectra of carbonaceous material (RSCM) analysis is a paleothermometry technique for rocks that have experienced temperatures from ~200° C to ~600° C [START_REF] Beyssac | Raman spectra of carbonaceous material in metasediments: A new geothermometer[END_REF][START_REF] Lahid | Evolution of the Raman spectrum of carbonaceous material in low-grade metasediments of the Glarus Alps (Switzerland)[END_REF]. It is used to measure the peak temperature a given sample has experienced and gives no age. Raman spectra of carbonaceous material have several distinct peaks at determined Raman shits. he precise location and amplitude of these peaks is a function of the temperature experienced by the sample. Because preparation of the sample into a thin section involves lots of abrasive processes, the temperature experienced by the surface of the thin-section is an unreliable thermometer. herefore, it is necessary to ind carbonaceous material that is not at the surface of the sample, but underneath a transparent mineral (usually calcite). Because the analysis and data processing involve highly subjective steps, a large number of measurements per sample is needed to give a statistically meaningful result. 10 spectra per sample is a good minimum value [START_REF] Lahid | Evolution of the Raman spectrum of carbonaceous material in low-grade metasediments of the Glarus Alps (Switzerland)[END_REF].

Unsuccessful analysis

93 rock samples were collected in the ield in the North Pyrenean Zone, in particular around the Metamorphic Internal Zone and the contact with the non-or less metamorphic North Pyrenean Zone. he goal was to establish the peak temperatures at various locations and produce a map that might help unravel the structure of the Metamorphic Internal Zone and its contacts. 41 Samples were processed into thin sections, and 13 samples were analysed using the RSCM technique. Most samples were white or pinkish, highly recrystallized marbles. he organic content proved insuicient to ind enough good spots to analyse, oten only resulting in a single, two, or three spectra per sample. Several samples returned very large diferences between measurements. he analysis was therefore not statistically reliable enough to use. he sample names, locations and unsuccessful Raman results are added here. Following this failed analysis, the samples were handed to Corentin Gouache, a master student at the time, for reanalysis. His results (among other things) can be found in his MSc thesis [Gouache, 2017]. 132

Stellingen

It is a Dutch tradition to place a list of propositions ('stellingen') at the end of the thesis. Originally, these were statements related to the topic of research that the PhD candidate had to successfully defend publicly. Over time, the emphasis shited more and more towards judging the written work of a PhD candidate, and the propositions lost their original purpose. Nowadays, the list of stellingen should still include some related to the topic, but may also contain unrelated topics that the PhD candidate happens to be interested in or feels strongly about. he last stelling is usually an attempt at humor.

• Using a balanced and fully restored cross section in combination with fault timings to calculate shortening gives better results with equal or less work than stepwise restoration.

• he rate of displacement of a suprasalt cover unit in relation to its basement can be used to recognise gravitational sliding in convergent settings.

• A lux steady state (accretionary inlux = erosional outlux) of orogens or orogenic wedges is iction.

• here should be a half-open access science publishing platform analogous to Spotify or Youtube.

• he ratio of references cited in a study that are also cited by other references cited in the same study to 'fresh' references can be used as a measure of quality.

• Isolation and motivation are inversely correlated.

• Successful procrastination is a ine art that requires lots of practice.

Du système de rit à l'orogène à double vergence : Un modèle évolutif basé sur l'étude de cas des Pyrénées Orientales et une étude des facteurs de contrôle à partir des modèles numériques. From rit system to doubly vergent orogen: An evolutionary model based on a case study of the Eastern Pyrenees and controlling factors from numerical models.

he doubly vergent nature of some natural orogens is classically understood as two opposing thrust wedges (pro and retro) that comply with critical taper theory. he evidence that retro-wedges and their associated basins behave diferently from their prowedge counterparts has been steadily increasing over the past few decades. However, what causes an orogen to become doubly vergent is currently not well understood. Nor is the relationship between the pro-and retro-wedge during the evolution of a doubly vergent orogen. It is the aim of this work to improve our understanding of: 1) how the pro-and retro-wedges relate to each other during the orogenic process, 2) what factors control the evolution of a doubly vergent orogen and 3) a possible link between the pro-and retro-wedge. Answering these questions requires an improved knowledge of the evolution of a doubly vergent orogen. We focussed on the Eastern Pyrenees as a type example of a doubly vergent orogen, due to the large amount of available data. We performed a detailed tectonostratigraphic study of the retro-foreland of the Eastern Pyrenees (European plate), updating the interpretation based on recent insights into its hyperextended rit origins. We link the evolution of the retro-foreland to that of the pro-foreland (Iberian plate) in order to derive insight into the crustal scale dynamics. Based on cross section restoration, reconstructed shortening rates and subsidence analysis, we subdivide the East Pyrenean evolution into four phases. he irst (Late Cretaceous) phase is characterised by closure of an exhumed mantle domain between the European and Iberian rited margins, and simultaneous inversion of a salt-rich, thermally unequilibrated rit system. Shortening was distributed roughly equally between both margins during this early inversion phase. Following inversion, a quiescent phase (Paleocene) was apparently restricted to the retroforeland. his phase may record the period of transition between inversion and full collision in the Eastern Pyrenees. he main collision phase (Eocene) records the highest shortening rates, which was predominantly accommodated in the pro-wedge.

Retro-wedge shortening rates were lower than during the rit inversion phase. During the inal phase (Oligocene) the retrowedge was apparently inactive and shortening of the pro-wedge slowed. his demonstrates that the relationship between the pro-and retro-wedges changes through time.

We used lithosphere-scale thermo-mechanical numerical models to simulate the evolution of a doubly vergent orogen. Our results show a similar evolutionary pattern as observed in the Pyrenees: A roughly symmetrical rit inversion phase is followed by an asymmetric collision phase. Rit inheritance was found to be essential for enabling double vergence. Other factors, such as surface processes and thin-skinned deformation, were found to have a signiicant efect on the crustal structure and strain partitioning between both wedges. A salt décollement layer in the sedimentary cover promotes the formation of a crustal antiformal stack such as observed in the Pyrenees and Alps by forming a wide and low-taper thin-skinned fold-and-thrust belt that forces crustal deformation to focus in the hinterland. Finally, we show that the evolution of the pro-and retro-wedges is inextricably linked: events or conditions on one side of the doubly vergent orogen have an immediate efect on the other side of the orogen. his is clearly demonstrated in our models by constant variations in shortening rates of the pro-and retro-wedge in response to accretion of new pro-wedge thrust sheets. he High Atlas (Morocco) and Pyrenees can be seen as examples of symmetric rit inversion and later asymmetric collision phases, respectively. PhD hesis, Grool, A.R., 2018, Université de Lorraine, Nancy, France & University of Bergen, Norway